Rocket-Borne In-Situ Measurements in the Middle Atmosphere

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Rocket-borne in-situ measurements in the middle atmosphere

Jonas Hedin

AKADEMISK AVHANDLING för filosofie doktorsexamen vid Stockholms Universitet att framläggas för offentlig granskning den 20:e februari 2009

Rocket-borne in-situ measurements in the middle atmosphere Doctoral thesis Jonas Hedin

Cover photo: Launch of the HotPay-I sounding rocket at Andøya Rocket Range, Norway on July 1, 2006 (photo by Jonas Hedin).

ISBN 978-91-7155-813-8, pp. 1 - 59 © Jonas Hedin, Stockholm 2009

Stockholm University Department of Meteorology 10691 Stockholm Sweden

Printed by Universitetsservice US-AB Stockholm 2009

Contents

Abstract 5

List of papers 7

1 Introduction 9 1.1 The Earth’s atmosphere 9 1.2 Important mesospheric phenomena and their study 12 1.2.1 Ice particles 13 1.2.2 Meteoric smoke 17 1.2.3 Nightglow 23

2 Atmospheric science with sounding rockets 26 2.1 The need for sounding rockets 26 2.2 Challenges 27

3 The PHOCUS project 30 3.1 Science questions 31 3.2 Instrumentation 35

4 Results of this thesis 39

5 Outlook 43

Acknowledgements 45

Bibliography 47

Abstract

The Earth's mesosphere and lower thermosphere in the altitude range 50- 130 km is a fascinating part of our atmosphere. Complex interactions between radiative, dynamical, microphysical and chemical processes give rise to several prominent phenomena, many of those centred around the mesopause region (80-100 km). These phenomena include noctilucent clouds, polar mesosphere summer echoes, the ablation and transformation of meteoric material, and the Earth’s airglow. Strong stratification and small scale interactions are common features of both these phenomena and the mesopause region in general. In order to study interactions on the relevant spatial scales, in situ measurements from sounding rockets are essential for mesospheric research.

This thesis presents new measurement techniques and analysis methods for sounding rockets, thus helping to improve our understanding of this remote part of the atmosphere. Considering the need to perform measurements at typical rocket speeds of 1 km/s, particular challenges arise both from the design of selective, sensitive, well-calibrated instruments and from perturbations due to aerodynamic influences. This thesis includes a quantitative aerodynamic analysis of impact and sampling techniques for meteoric particles, revealing a distinct size discrimination due to the particle flow. Optical techniques are investigated for mesospheric ice particle populations, resulting in instrument concepts for accessing smaller particles based on Mie scattering at short ultraviolet wavelengths. Rocket- borne resonance fluorescence measurements of atomic oxygen are critically re-assessed, leading to new calibration concepts based on photometry of O 2 airglow emissions.

The work presented here also provides important pre-studies for the upcoming PHOCUS rocket campaign from Esrange in July 2010. PHOCUS (Particles, Hydrogen and Oxygen Chemistry in the Upper Summer mesosphere) will address the interaction between three major mesospheric players: meteoric smoke, noctilucent clouds and gas-phase chemistry.

List of papers

This thesis consists of an introduction and the following five papers

I Hedin, J., Gumbel, J., and Rapp, M.: On the efficiency of rocket-borne particle detection in the mesosphere, Atmospheric Chemistry and Physics, 7, 3701-3711, 2007.

II Hedin, J., Gumbel, J., Waldemarsson, T., and Giovane, F.: The aerodynamics of the MAGIC meteoric smoke sampler, Advances in Space Research, 40, 818-824, 2007.

III Rapp, M., Hedin, J., Strelnikova, I., Friedrich, M., Gumbel, J., and Lübken, F.-J.: Observations of positively charged nanoparticles in the nighttime polar mesosphere, Geophysical Research Letters, 32, L23821, doi:10.1029/2005GL024676, 2005.

IV Hedin, J., Gumbel, J., Khaplanov, M., Witt, G., and Stegman, J.: Optical studies of noctilucent clouds in the extreme ultraviolet, Annales Geophysicae, 26, 1109-1119, 2008.

V Hedin, J., Gumbel, J., Stegman, J., and Witt, G.: Use of O 2 airglow for calibrating direct atomic oxygen measurements from sounding rockets, manuscript for submission to Atmospheric Measurement Techniques, 2009.

Chapter 1

Introduction

1.1 The Earth’s atmosphere

The Earth’s atmosphere (Greek ‘atmos’ = ‘vapour’ + ‘sphaira’ = ‘ball’) can be divided into regions and boundaries, or “spheres” and “pauses”, in several ways based on the properties the atmosphere is described in [Brasseur and Solomon , 1986]. The most common way is a division according to the temperature change with altitude (figure 1). The lowermost region called the troposphere (Greek ‘tropos’ = ‘turning’ or ‘change’) is where our daily weather takes place. It extends to about 7 km at the poles and about 17 km at the equator and is characterised by rather strong vertical mixing, hence its name. Molecules can traverse the entire depth in periods between a few days and a few minutes depending on the weather situation. The mixing is due to solar heating of the surface that in turn heats the air masses which then become less dense and rise. While rising, it cools adiabatically and latent heat can be released as water vapour condenses, leading to a vertical temperature decreases of typically ~ 6.5 K/km up to the tropopause , where a local temperature minimum of about 200 - 230 K is reached depending on latitude. In the stratosphere (Latin ‘stratus’ = ‘layered’) the temperature increases with altitude due to the absorption of shortwave solar ultraviolet (UV) radiation by ozone (O 3) ( λ ~ 200 - 300 nm). The main cooling agent in the stratosphere is the ν2 band (bending vibrational state) of carbon dioxide (CO 2) at infrared (IR) wavelengths of 15 µm. The stratosphere extends up to the stratopause at about 50 km and is characterised by very small vertical mixing with transport time-scales of the order of years. The stratosphere is very dry since only about 3 ppmv of water vapour is transported from the troposphere past the ‘cold trap’ at the tropopause. An additional ~3 ppmv of water vapour is produced in the

9 stratosphere from the oxidation of methane (CH 4), transported from the Earth’s surface.

120 Thermosphere Heterosphere

100 turbopause

mesopause 80 Mesosphere Summer profile Vinter profile 60 stratopause

Altitude [km] 40 Stratosphere Homosphere

20 tropopause Troposphere 0 100 150 200 250 300 350 Temperature [K]

Figure 1. The temperature structure of the atmosphere.

In the mesosphere (Greek ‘mesos’ = ‘middle’) ozone levels and, hence, direct shortwave heating at 200 – 300 nm is low. As a consequence, adiabatic temperature decrease with altitude is dominant again, as in the troposphere. In the upper mesosphere, between about 80 and 100 km, the mesopause region forms a boundary to the thermosphere above. In summer this boundary is more sharply defined, with temperature minimum that at polar latitudes can reach extremely low temperatures (down to 100 K) around 90 km [ Höffner and Lübken , 2007], while during other seasons the minimum (~180 - 200 K) is less well defined and can occur over a range of several 10 km. Above the mesopause, in the thermosphere (Greek ‘thermos’ = ‘heat’), the absorption of extreme ultraviolet (EUV) and X-ray radiation at wavelengths below 180 nm, mostly by O 2, leads to a rapid warming with altitude.

The summertime polar mesopause is the coldest place on Earth and it is here that ice particles form and can be seen optically as noctilucent clouds in twilight conditions. It is also in this region where meteoroids deposit much of their material as they burn up entering the atmosphere, and excited atoms and molecules emit radiation known as nightglow (or more general airglow) (see chapter 1.2.3 and paper V). The reason why the upper

10 mesosphere is so cold in summer and warmer in winter is the wave-driven general circulation. A detailed review of gravity wave dynamics and its effects in the middle atmosphere has been given by Fritts and Alexander [2003] Gravity waves generated in the lower atmosphere (e.g. in flow over topography, convection, wind shear, geostrophic adjustment and wave- wave interactions) propagate upwards and eventually become unstable and break, depositing momentum and energy in the mesosphere. This wave forcing drives the atmospheric flow velocity into the direction of the wave phase velocity. In summer this acts as a “wave drag” on the prevailing zonal winds, slowing down the mean wind near the mesopause and completely overwhelms the radiative forcing which otherwise would maintain very high easterly wind speeds at mesopause heights. The wave drag thus forces a reversal of the zonal winds from easterly at lower heights to westerly in the lower thermosphere. It also forces a mean equatorward wind of the order of 10 m/s in the summer and a similar poleward flow in the winter hemisphere. The equatorward moving air, in combination with the radiative forcing, induces strong upward motion (of the order of several cm per second) over the summer pole as a result of mass continuity. The rising air cools adiabatically, resulting in the very low temperatures of the summer mesopause while the opposite circulation occurs in the other hemisphere, and downwelling warms the wintertime mesosphere.

Another important way of describing the atmosphere is to divide it according to how the atmospheric composition changes with altitude. Composition changes occur with altitude both because of transport properties and because of chemical properties. As for transport, the altitude range from the Earth’s surface up to the turbopause at about 100 km is called the homosphere as it features more or less uniform conditions due to turbulent mixing of the atmosphere. However, above the turbopause, in the heterosphere , the composition starts to vary with altitude. This is because in the absence of turbulent mixing the rate of fall-off in density of a species with increasing altitude depends on the molecular mass. Hence, heavier atmospheric constituents such as molecular oxygen and nitrogen fall off more quickly than lighter constituents such as atomic oxygen, hydrogen and helium. The absence of turbulence above ~100 km is closely related to the increasing molecular mean free path with altitude. This increase of the mean free path does not only have a major effect on atmospheric composition but also on aerodynamics, as will be discussed in papers I and II. As for chemistry, composition changes with altitude as sources and sinks for various species are highly variable. In the middle and upper atmosphere, this is in particular true for the flux of solar UV radiation that provides much of the energy for atmospheric photochemistry. This is the basis e.g. for the production of odd oxygen, the partition between atomic oxygen and

11 ozone, and the generation of excited airglow species that will be discussed in paper V.

The ionosphere is the altitude region of the atmosphere where ionisation of atoms and molecules by shortwave solar UV radiation occurs. It includes both the mesosphere and the thermosphere. The ionosphere can be divided up into several layers depending on the main ionised species. In the mesosphere, NO is ionised by the very intense solar Lyman-α line at 121.57 nm, giving rise to the ionospheric D layer. Photolysis by Lyman-α is also the main loss process of water vapour. The E-layer between 90 and 120 km is formed by soft X-ray and far UV ionisation of O 2 and the F-layer above 120 km by EUV ionisation of O. The interaction of ionospheric charges with atmospheric particles will be discussed in paper III.

Figure 2. Schematic overview of the middle atmosphere processes (water vapour in ppm(v)) .

1.2 Important mesospheric phenomena and their study

The mesosphere and lower thermosphere (MLT) region of the Earths atmosphere between 50 and 130 km is a transition region that is controlled by a combination of radiative and dynamic forcings from above and below (figure 2). Prominent phenomena related to the complex interactions in this region are noctilucent clouds (NLC) [ Rapp and Thomas , 2006], polar

12 mesosphere summer echoes (PMSE) [ Rapp and Lübken , 2004], the ablation and transformation of meteoric material [ Plane , 2003] and the Earth’s nightglow [ Meriwether , 1989]. Systematic observations of NLC, PMSE, meteoric smoke or the nightglow provide important clues to our understanding of the middle atmosphere as a whole. However, in order to draw robust conclusions from such observations, we need a more robust understanding of the physical and chemical processes behind these phenomena.

Figure 3. Noctilucent clouds observed over Stockholm. (Photo by Nathan Wilhelm)

1.2.1 Ice particles

NLC’s (figure 3) and PMSE’s are related to ice particles in the polar summer mesopause region. Existing just at the edge of feasibility, mesospheric ice clouds are expected to be extremely sensitive to changes in middle atmospheric conditions and have thus gained interest as indicators of global variability and trends. It has been argued that even small long- term changes of mesospheric water vapour (e.g. due to anthropogenic methane emissions or changes in lower atmospheric circulation patterns) or mesospheric temperatures (e.g. due to anthropogenic carbon dioxide emissions) should lead to prominent long-term changes of the observed

13 properties of mesospheric ice clouds [ Thomas et al. , 1989]. The question whether such long term changes are already evident in the experimental records has been debated [ von Zahn , 2003; Thomas et al. , 2003]. In addition, the occurrence of mesospheric ice particles has been discussed in connection with space traffic exhaust [ Stevens et al. , 2003, 2005], observed differences between the Arctic and Antarctic mesosphere [ Hervig and Siskind , 2005], and the middle atmosphere coupling between the two hemispheres [ Becker et al., 2004; Karlsson et al ., 2007].

We know today that mesospheric clouds are composed of water ice particles, possibly with small nucleation cores. A detailed review about current knowledge of the microphysics of mesospheric ice particles has been given by Rapp and Thomas [2006]. Ice particle nucleation takes place at the altitudes with the largest saturation ratios S=p H2O /p Sat (where p H2O is the water vapor partial pressure and p Sat is the equilibrium vapor pressure of water vapor over ice). This is the case is at altitudes close to the temperature minimum between ~86 and 90 km. Lübken [1999] found that typical saturation ratios in the mesopause region are on the order of ~100 (although significantly higher supersaturations can be expected for wave- disturbed conditions). Thus, homogeneous nucleation should be negligible, and pre-existing ice nuclei are required for the formation of ice particles. Once formed, the ice particles then continue to grow by direct deposition of water vapor onto their surface. They settle down due to gravity and consume the available water vapor on their way down through the atmosphere. During this motion, the particles are further subjected to transport by mean winds and small scale motions, i.e. waves and turbulence. Once the particles have reached equivalent spherical radii in excess of ~30 nm, they scatter light efficiently such that they may eventually be observed optically with the naked eye or from ground based or space borne optical instruments as NLC or polar mesospheric clouds (PMC).

All these ice particles are also immersed in the plasma of the ionospheric D-region. Hence, electrons attach to the ice surfaces so that the particles become charged. In addition, the 80 – 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 persistent small scale structures in the ionospheric charge distribution. This causes irregularities in the radio refractive index (which is effectively determined by the electron number density at these altitudes) which can be observed from the ground by radars as PMSE [ Rapp and Lübken , 2004]. Important is that the lifetime of the irregularities against diffusion is determined by the radius of the ice

14 particles. Typical values for the lifetime are just a few minutes for small radii at altitudes above ~85 km and up to hours for the larger radii below. As a consequence, PMSE at altitudes above typically 85 km can only exist during phases of active neutral air turbulence. In contrast, the long lifetime below this altitude “freezes” structures in the plasma structures so that PMSE also can occur long after the disappearance of neutral air turbulence.

One central open question is the actual origin of the ice particles. Homogeneous nucleation is very unlikely to occur at typical mesospheric supersaturation levels, so there has to be some kind of nucleus that the water molecules can nucleate on. Candidates that have been proposed in the past are: large proton hydrate ion clusters [ Witt , 1969], soot [ Pueschel et al ., 2000], sulfuric acid aerosol particles [ Mills et al. , 2005], sodium bicarbonate [ Plane , 2000], sodium hydroxide [ Petrie , 2004] and meteoric smoke particles [ Rosinski and Snow , 1961; Hunten et al. , 1980]. The meteoric smoke particles have long been considered as the most likely candidate but there are uncertainties: The influx, or the amount of meteoric material that enters the Earth’s atmosphere per year, has an uncertainty of one to two orders of magnitude. The processes of meteoric ablation and re- condensation into smoke particles are poorly understood. And, last but not least, it has recently been recognized that the global circulation efficiently transports meteoric material away from the polar summer mesopause region [ Megner at al ., 2008a]

The initial step of ice particle formation, the nucleation process, introduces the most significant uncertainty into the current theoretical understanding of mesospheric ice microphysics. The nucleation feasibility and rate of ice particle formation is linked to the type of ice nuclei, including composition and corresponding properties like their crystal structure, their number density and size distribution, as well as their charge state. The actual shape of the particles is uncertain and this affects the growth rate and sedimentation speed. In most modeling efforts the particles are considered to be spherical, but as they are most probably non-spherical [ Baumgarten et al. , 2002; Eremenko et al. , 2005], the sedimentation speed will decrease and hence the particle will spend more time in the altitude range with favorable conditions for growth. Also the equilibrium vapor pressure of water vapor over ice, or the saturation vapor pressure, is uncertain for the temperatures relevant to the polar summer mesopause region. In addition, the very nucleation process on nanometer-size nuclei is a major source of uncertainty since the commonly applied liquid drop theory is of limited validity under these conditions [ Kessee , 1989; Gumbel and Megner , 2008]. Finally, the accommodation coefficient, i.e. the efficiency of collisional energy transfer from ice particles and the surrounding air molecules, is

15 another major uncertainty that affects both ice particle temperature and sedimentation speed.

The size range of visible NLC particles has been studied by optical measurements with satellites [e.g. Rusch et al ., 1991; Merkel et al ., 2003; Karlsson and Rapp , 2006; Bailey et al ., 2008], sounding rockets [ Gumbel et al ., 2001], and lidars [ von Cossart et al. , 1999; Baumgarten et al. , 2007]. The size retrievals are based on the comparison of measured scattering properties for sunlight (spectral dependence or phase function) to theoretical simulations of the size dependence of Mie scattering. According to these measurements the particle radii generally remain smaller than 100 nm. However, from normal optical measurements it is not possible to access the smaller particles in the mesospheric ice population. This has two reasons: (1) The scattering is strongly dominated by the larger particles in the population because of the strong size dependence of the scattering. (2) Smaller sizes become indistinguishable because they all fall into the Rayleigh scattering regime, i.e. the particle radius is much smaller than the light’s wavelength. The current lack of knowledge on the smaller particles is a major limitation when studying important questions regarding the nucleation and growth processes governing the ice particles in the mesosphere.

Optical measurements are today the only means of obtaining information about the size of NLC particles. New information about the properties of the particles has become available from measurements in the infrared (Hervig et al. , 2001; Eremenko et al. , 2005). However, in order to access particle sizes efficiently, the deviation of the particle’s scattering/absorption properties from the pure Rayleigh regime needs to be studied. This requires wavelengths of the order of the particle size and, hence, measurements in the visible and, in particular, the ultraviolet part of the spectrum. Previous NLC measurements have been made at wavelengths down to 355 nm with ground-based instruments (e.g. von Cossart et al. , 1999) and down to about 200 nm with rockets and satellites (e.g. Rusch et al. , 1991; Gumbel et al. , 2001). In paper IV we argue that measurements at even shorter wavelengths in the extreme ultraviolet (EUV) are necessary to extend our knowledge of the NLC particle size distribution. In particular, Lyman-α radiation at 121.57 nm is well suited for such studies as NLC are well illuminated by solar Lyman-α during daytime conditions. At 121.57 nm even small mesospheric ice particles come well into the Mie scattering regime and start to influence the total NLC scattering. Therefore, an investigation of the scattering and extinction properties of ice particles at Lyman-α can yield important complementary information about the polar mesospheric ice layer.

The SLAM instrument (Scattered Lyman-Alpha in the Mesosphere) presented in paper IV is a new development and a test of the measurement concept for NLC scattering of solar Lyman-α. The instrument is designed as a double monochromator. A first flight opportunity for measuring Lyman-α in the vicinity of NLC was in July 2006 with SLAM as part of the eARI HotPay-1 sounding rocket from Norway. The SLAM instrument was launched together with four other scientific instruments on 1 July 2006. While this test flight was intended to be the first measurement of the altitude- and angle-dependent scattering of Lyman-α by NLC in the mesosphere, a malfunction of the payload shortly after launch unfortunately made it impossible to collect any mesospheric data. A second opportunity to measure the scattering of Lyman-α by NLC will be on the PHOCUS sounding rocket (section 3).

1.2.2 Meteoric smoke

The Earth’s atmosphere is constantly bombarded by meteoric material. Most of the incoming meteoroid dust is in the size range 10-300 µm [Ceplecha et al. , 1998]. Depending on entry speed and inclination, smaller particles (< 20 µm) may survive almost intact as micrometeorites, while larger particles (>100 µm) are believed to ablate and vaporize entirely in the 70-130 km altitude region. Estimates of the total amount of incoming material are still uncertain and vary between 10 and 100 tons per day [ Love and Brownlee , 1993; Mathews et al. , 2001; von Zahn , 2005]. Also our knowledge about the fate of the vaporized material is rather limited. It is well recognized that meteoroids are the source of the metal atom layers that are observed by lidar [Kane and Gardner , 1993] and satellite [ Gumbel et al., 2007], and much progress has been made in understanding the chemistry of these metals [ Plane , 2003]. It is further thought that this vapour recondenses into secondary nanometre-size particles, which are then subject to coagulation, sedimentation and global transport [Rosinski and Snow , 1961; Hunten et al ., 1980; Gabrielli at al ., 2004; Megner et al., 2006]. Processes of relevance for this “meteoric smoke” are schematically summarized in figure 4.

Figure 4. Schematic of the fate of meteoric material in the atmosphere.

Once formed from the ablated meteoric material, meteoric smoke particles are thought to play a major role in mesospheric processes. Examples are:

• Smoke particles are today recognized as the major candidate for condensation nuclei for ice particles and, hence, phenomena like noctilucent clouds and polar mesosphere summer echoes (section 1.2.1). • Heterogeneous chemistry has been suggested to take place on the surface of smoke particles. An example is the proposed catalytic recombination of O and H 2 as a local source of water in the mesosphere [ Summers and Siskind , 1999]. By scavenging various gas-phase products of meteoric ablation, smoke is also thought to act as (ultimate) sink in mesospheric metal chemistry [ Plane , 2004]. • Smoke can play a significant role in the D-region charge balance [Rapp and Lübken, 2001]. While the particles efficiently scavenge free electrons and positive ions, the charging may in turn strongly affect smoke coagulation processes.

In the stratosphere meteoric particles are thought to constitute an important part of the background aerosol. Murphy et al . [1998] showed that a high proportion of stratospheric aerosol is meteoric. The importance of meteoric smoke in the middle atmosphere relates to the fact that input from the troposphere is small. The only dominant dust sources from below are major volcanic eruptions and gas-to-particle conversion, mainly of sulphur

18 compounds complemented by the existence polar stratospheric clouds (PSC), featuring varying composition (water ice, supercooled ternary solutions (STS), nitric acid trihydrate (NAT)) and important effects on ozone chemistry. Due to the strong downward circulation in the winter polar vortex, meteoric smoke can be efficiently transported to the lower winter stratosphere [ Megner et al ., 2008b]. Here they become involved in the microphysics of PSC. In order to explain the observed composition of PSC particles, Voigt et al . [2005] have recently suggested meteoric smoke as a trigger of heterogeneous nucleation of NAT. Smoke particles have also been used as a tracer for atmospheric circulation [ Gabrielli et al ., 2004; Lanci and Kent , 2006].

Despite this potential importance of mesospheric particles very little is known about their properties. This is because the only means of accessing them is by in situ measurements from sounding rockets. These measurements are very difficult and there are a number of challenges inherent to sounding rocket experiments that need to be considered, such as aerodynamics [ Gumbel , 2001; Hedin et al., 2005] and charging processes (payload charging). The basic aerodynamic challenge lies in the size of the meteoric smoke particles. Smoke particles are so small that they tend to follow the airflow around the payload rather than reaching the detector. Since we want the particles to hit the detector surface, careful aerodynamic design is of critical importance for smoke particle experiments. The interpretation of the particle measurements requires also a detailed understanding of the specific detector response, and this is far from trivial. Basic instrumental questions are:

• How are measured particle concentrations related to the undisturbed particle concentrations in the atmosphere? • How are properties of detected particles related to the particle properties in the atmosphere? • How is the charge measured by particle detectors related to the charge of particles in the atmosphere?

Paper I and II focus on the aerodynamical aspects of these questions.

Measurement of charged smoke particles

Most data on meteoric smoke available today is indirect evidence for the existence of the particles from measurements of heavy charge carriers.

19 Recent examples of charged particle measurements have been reported by Gelinas et al . [1998], Horányi et al . [2000], Lynch et al. [2005], Rapp et al. [2005, paper III], and Rapp and Strelnikova [2008]. The major class of detectors aimed at the detection of charged meteoric smoke particles is based on a detector design originally developed by Havnes et al . [1996] for the study of ice particles in the polar summer mesosphere. The detector concept uses a Faraday Cup for the detection of heavy charge carriers in combination with biased grids that shield the electrode in the inner part of the detector against contamination by electrons and light ions from the ambient ionospheric D-region. Gelinas et al. [1998] were the first to apply a detector of this kind to the rocket-borne study of meteoric smoke particles and Lynch et al . [2005] further developed the design to allow the discrimination between positive and negative particles.

Figure 5. Left panel: The top section of the ECOMA payload for meso spheric particle studies. The ECOMA particle detector from IAP is in the centre and the MAGIC particle sampler from MISU/NRL in the foregr ound. Other instruments are tem perature probes from IAP (left, right) and ionospheric instrumentation from the Technical University Graz (booms). Right panel: The ECOMA detector.

By combining the original design of Havnes et al . [1996] with a xenon flash lamp, Rapp et al. [2005, paper III] provided a way to detect neutral particles by photoionization. This detector is the central part of the ECOMA sounding rocket project (figure 5). The ECOMA project (Existence and Charge state Of Meteor smoke particles middle Atmosphere) comprises a series of 4 campaigns from Andøya Rocket Range. The project is led by the Leibniz Institute of Atmospheric Physics (IAP) and the Norwegian Defence Research Establishment (FFI). One campaign in dark autumn conditions in September 2006 and two in polar summer conditions in August 2007 and July 2008, respectively, have already taken place. Results are described by Rapp and Strelnikova [2008] and in paper III. Additional instruments onboard the payload concern temperature, turbulence, ionospheric parameters, optical NLC properties (MISU) and particle sampling (see below), thus providing a very

20 comprehensive characterisation of neutral and ionospheric conditions. The campaigns are also supported by the extensive ground-based equipment located at ALOMAR and Andøya Rocket Range, in particular the ALWIN VHF radar, the SAURA MF wind radar, the SKiYMET meteor radar, and the ALOMAR RMR and sodium lidar. In addition it is supported by the EISCAT incoherent scatter radar facility.

In addition to Faraday cup type detectors other techniques have been applied in the study of meteoric smoke. These include mass spectrometry [Schulte and Arnold , 1992], Gerdien condenser [ Croskey et al ., 2001] and magnetically shielded probes [ Robertson et al ., 2004; Smiley et al ., 2006]. Signatures of charged particles have recently also been reported from incoherent scatter radar measurements [ Rapp, et al ., 2007a]. It is however important to note that none of the above measurements of heavy charge carriers provides a definite proof that the detected species really are smoke particles of meteoric origin. In order to provide such a proof and more detailed studies of the particles, instruments are needed that directly sample the smoke and this is the idea of the MAGIC meteoric smoke samplers discussed next and in paper II.

Paper I analyses the aerodynamics of smoke particle measurements by Faraday cup detectors. The paper provides a response function that specifies the fraction of atmospheric particles that is actually detected by a given instrument design as a function of particle size and charge, altitude, and flow conditions.

Rocket-borne sampling of smoke particles

Because of the lack of knowledge about meteoric smoke particles, they have less seriously been called “magic dust”. This is the origin of the name of the MAGIC rocket project (MAGIC: Mesospheric Aerosol - Genesis, Interaction and Composition). The major objective of the MAGIC project was the development of a sampling instrument and analysis techniques that for the first time would allow smoke particles to be brought back from the mesosphere to the laboratory and to study their properties in detail. The MAGIC detector concept has been developed in collaboration between the Naval Research Laboratory in Washington D.C. and German and Swedish scientists [Gumbel et al ., 2005]. As nanometre-size particles tend to follow the airflow around payload structures, a basic idea of MAGIC is to minimise aerodynamical perturbations in the measurement volume so that the particles could be sampled. This resulted in an instrument design with sampling surfaces of 3 mm diameter located on top of pins that are ejected

21 from the payload top (see figure 6). The aerodynamics of the MAGIC meteoric smoke sampler is the topic of paper II. The sampling surfaces are standard transmission electron microscopy (TEM) grids and can be taken directly from the measurement pin to the microscope for analysis. The basic scientific questions addressed in this project are:

• Do re-condensed smoke particles of cosmic origin exist in the mesosphere? • What is their number density and size distribution? • What is the spatial distribution of the particles and how are they transported? • What is the elemental and molecular composition of the particles? • What is their charge state and how do they interact with their mesospheric and ionospheric environment?

Figure 6. a) The MAGIC collectio n unit with extended sampling pin. b) The top of MAGIC sounding rocket payload with a sampling probe extended from one of the three MAGIC sampling units surrounding the Hygrosonde instrument [ Khaplanov et al ., 1996; Lossow et al , 2008a] in the centre. c) T EM grid mounted with copper cap on top of an extended sampling pin.

Earlier rocket campaigns with the MAGIC instruments were conducted from Esrange, Sweden, in January 2005 [Gumbel et al ., 2005], from Wallops Island, USA, in May 2005, and as part of the German/Norwegian ECOMA project from Andøya, Norway, in September 2006, August 2007 and July 2008. However, these flights have not yet resulted in any quantitative results. The September 2006 campaign featured launches of two payloads, each carrying two MAGIC’s. Analysis of the first launch did not give quantitative results due to contamination of the sampling grids, most likely from the payload itself, while the second payload was lost due

22 to problems with the recovery system. Recovery problems also led to the lack of results in 2007. In July 2008 three payloads were successfully launched, this time with one MAGIC particle sampler on each payload. Samples from these flights are currently being analysed, but quantitative results are not yet available.

The analysis of the MAGIC smoke samples after recovery is primarily based on detailed TEM studies at Naval Research Laboratory and, since 2008, at a new Centre for advanced electron microscopy studies in the Department of Physical, Organic, Inorganic and Structural Chemistry at Stockholm University. The TEM analysis is intended to provide particle numbers, sizes and shapes, while the combination of the TEM technique with energy-filtered imaging and energy-dispersive x-ray spectroscopy can provide the elemental composition of the sampled particles.

MAGIC is currently scheduled for one future rocket flight as a central part of the PHOCUS project, focusing on the interaction between meteoric smoke, noctilucent clouds and gas-phase chemistry (section 3).

1.2.3 Nightglow

The atmosphere is constantly emitting radiation. This radiation called airglow is emitted by excited atoms, molecules or ions generated in a variety of production and excitation processes following the interaction of sunlight with atmospheric constituents during the day. The airglow can be divided into several types depending on the time of day they are observed and include dayglow, twilightglow and nightglow. The nightglow is rather faint and can be compared in intensity to the light of a candle at a distance of 100 m. Although the dayglow intensities are several orders of magnitude larger, it is not detectable by eye because of the strong scattering of sunlight in the atmosphere. The airglow is driven by photons from the sun, whereas auroras are excited by the impact of energetic particles. Outside the auroral region or in times of quiet geomagnetic conditions, the terrestrial night sky spectrum is dominated by emission features related to oxygen photochemistry. Molecular oxygen has relatively low dissociation energy and O 2 photolysis, mainly in the Schuman-Runge bands and continuum, is therefore a very efficient way of converting solar radiation into chemical energy.

The energy is deposited during daytime and the part not directly re-radiated or converted into heat is stored in the form of excitation of metastables and

23 dissociation energy. Atomic oxygen is the major carrier of this chemical energy in the upper atmosphere and it is the recombination of atomic oxygen that generates molecular oxygen in a number of metastable states. With the energy available in this recombination reaction, high vibrational 3 − 1 levels of the ground state (X Σ g ) or any of the five bound states (a ∆ g , 1 + 1 − 3 + 3' 5 b Σ g , c Σu , A Σu , A ∆u ) may be formed. Another weakly bound state, ∏g, has been theoretically predicted but not yet observed and has been proposed as a likely precursor to the above lower-lying light-emitting states. Through a number of energy transfer and quenching processes the excited O 2 then gives rise to nightglow emissions covering the spectral range from the ultraviolet to the infrared [ Meriwether , 1989; Mlynczak , 1999; Wayne , 2000]. These processes cause a complex and strongly altitude-dependent photochemistry between about 85 and 105 km. Other emissions that contribute to the nightglow are emissions from atomic oxygen [Barth and Hildebrandt , 1961], and vibrationally excited OH molecules [ Bates and Nicolet , 1950]. A fourth source of nightglow emissions is the emissions of metallic atoms such as sodium, calcium, potassium and magnesium. Also for these emissions, atomic oxygen is essential as its reaction with metal oxides is the major production channel of metal atoms in excited states [Plane , 2003].

Although it is generally accepted that the energy of excitation originates from the recombination of oxygen atoms, the exact mechanisms and reaction rates took long time to be established, and some have still not been established today [ Slanger and Copeland , 2003]. Much of today’s knowledge is based on the Energy Transfer in the Oxygen Nightglow (ETON) campaign that was conducted in March 1982 from South Uist, U. K. [Greer et al ., 1986]. A series of seven payloads was launched to measure altitude profiles of emission features and atmospheric composition. There have been several rocket experiments in the past with the objective to simultaneously study atomic oxygen density and oxygen airglow intensity [e.g. Thomas, 1981; Sharp , 1980; McDade et al ., 1984]. However, the most extensive of these multi-instrumented rocket studies was the set of campaigns with OXYGEN in March 1981 [ Witt et al. , 1984] followed by the ETON venture. The analysis of the ETON database resulted in detailed knowledge about nightglow excitation and quenching mechanisms [ Greer et al ., 1986; McDade et al ., 1986a; 1986b; 1987a; 1987b; Murtagh et al. , 1986]. This was not only interesting for atmospheric scientists but brought new understanding also to the basic physics of excited molecular states. In this sense, ETON used the mesosphere and lower thermosphere as a laboratory that provided experimental conditions that could hardly be achieved in conventional laboratory experiments.

Today, basic studies of airglow mechanisms are still an active field of research. But more importantly, airglow emissions are today applied as a versatile tool for atmospheric research. Remote sensing of airglow emissions from the ground or space is being used for studies of middle atmospheric composition, structure and dynamics over a wide range of spatial and temporal scales [e.g. Mlynczak et al. , 2000; French et al., 2005; Taylor et al ., 1997; Degenstein et al ., 2005]. The rocket studies originally relating the airglow processes to the state of the atmosphere have been the basis for these techniques. Measurements of the emission rates yield information about atmospheric composition, while studies of the Doppler shifts and the ratios of the airglow lines provide information about winds and other motions, and temperature, respectively [e.g. Shepherd et al. , 1993; Espy et al., 1997]. Measurements of line shifts and shapes are mainly performed from satellites or from ground-based instruments while the emission rates are readily measured also from sounding rockets. Ground- based imaging of nightglow patterns is used to study dynamics in terms of gravity wave activity in the upper mesosphere [e.g. Taylor et al. , 1997]. Airglow emissions have also been shown to be fundamental measures of energy deposition from which rates of atmospheric heating can readily be derived [ Mlynczak , 1999].

In paper V we present another use of O 2 airglow emissions. We demonstrate how airglow measurements can be used to calibrate direct atomic oxygen measurements by the resonance fluorescence technique from sounding rockets. The resonance fluorescence technique is sensitive and provides good altitude resolution, but is in itself notoriously difficult to calibrate [ Gumbel et al ., 1998; Dickinson et al. , 1980; Sharp, 1991]. As atomic oxygen is the major carrier of chemical energy in the MLT region accurate knowledge about the vertical distribution of O is crucial for any comprehensive study of chemical and dynamical processes in this part of the atmosphere [ Ward and Fomichev , 1993; Mlynczak and Solomon , 1993].

Chapter 2

Atmospheric science with sounding rockets

2.1 The need for sounding rockets

Sounding rockets are crucial as tools for atmospheric research. It is the features of the mesosphere and lower thermosphere that define the need for sounding rockets. As mentioned earlier, complex interactions between radiative, dynamical, microphysical and chemical processes give rise to many prominent phenomena such as NLC, PMSE, the ablation and transformation of meteoric material, or the Earth’s nightglow. Interactions on small spatial scales are a common feature of these phenomena and it is here that fundamental physics and new science can be revealed in the mesosphere today. When aiming at an understanding of these processes, we need to address interacting parameters simultaneously and on the small spatial scales of interest. Rocket-borne in situ studies with dedicated sets of experiments meet both these requirements. This cannot be provided by remote sensing techniques from the ground or from space.

For half a century sounding rocket experiments have thus been crucial as a source of knowledge about the mesosphere and thermosphere. In Sweden, the first scientific rocket was launched in 1961. Five years later Esrange was established. In Norway, the first scientific rocket was launched in 1962 from Andøya Rocket Range. The complexity of scientific questions has generally increased throughout the history of these rocket projects. In the early years, a new altitude profile of a single atmospheric parameter was certainly well worth a publication. Today, the strength of rocket projects

26 often lies in the study of complex interactions that involve many atmospheric parameters. In many cases, this increase in complexity has been based on a development towards larger multi-instrument payloads. But also the combination of in situ and remote sensing techniques has contributed to the growing complexity of rocket-borne studies.

Sounding rockets keep their important role in a world of global satellite measurements. Scientifically, the relationship between rockets and satellites is defined by the different nature of their measurements. The local 'snapshot' measurement makes sounding rockets a natural complement to the large-scale data provided by satellites. This can be used in different ways such as validation of satellite measurements by sounding rockets or real co-analysis of local and large-scale datasets. In addition to direct data validation and co-analysis, satellite research and rocket research have been inspiring one another in many other and fascinating ways. For example, in situ and ground-based results from the MaCWAVE/MIDAS rocket campaign indicated that unusual planetary wave activity in the southern hemisphere winter stratosphere could control the thermal conditions in the northern hemisphere summer mesosphere [ Goldberg et al. , 2004]. Subsequently, the existence of such an interhemispheric coupling was also suggested by general circulation modelling [ Becker and Fritts , 2006]. Based on these ideas, global noctilucent cloud data from the Odin satellite were then analysed and, indeed, showed a significant interhemispheric coupling on a year-to-year basis [ Karlsson et al , 2007]. This is one of many examples of rockets and satellites jointly contributing to the development of new scientific ideas.

Last but not least, an important connection between rocket and satellite projects is the exchange of technology. Many of today's remote sensing instruments on satellites build on experience from many years of rocket- borne optical instruments. Also the opposite way is possible: The experience with the submillimetre/millimetre radiometer onboard the Odin satellite [ Murtagh et al. , 2002] is today the basis for a new type of rocket- borne water vapour radiometer proposed for the PHOCUS rocket project (section 3).

2.2 Challenges

In situ experiments in the mesosphere are challenging. Many scientific and technical requirements need to be fulfilled by sounding rocket instruments.

27 The scientific skills behind rocket instruments and the need for engineering experience and continuity are not to be underestimated.

Scientifically, a basic requirement is to obtain quantitative measurements. This is far from trivial. The obstacles on the way from a simple detection idea to a quantitative rocket experiment may in the best case be surprisingly many, in the worst case they are just impossible to overcome. The first steps involve the physical measurement principle, the instrument design, and the design of appropriate calibration techniques. Beyond these steps come the challenges of the actual rocket flight. Most notably, our understanding of interactions between the rocket payload and its neutral and charged environment is still too limited. After half a century of sounding rockets, there remain a remarkable number of unsolved questions. New experimental as well as theoretical approaches are required to improve this situation.

An instructive example is the history of rocket-borne measurements of atomic oxygen [ Gumbel , 1999]. Resonance fluorescence has been used by many groups and calibration techniques have been developed in the laboratory, in situ and by theoretical means [e.g. Dickinson et al ., 1980; Sharp, 1991; Gumbel and Witt , 1997]. Nevertheless, measurement results show a spread that is most likely not geophysical. Paper V discusses these issues further.

Also the particle measurements discussed in chapter 1.2.2 can illustrate these points. What particle sizes can we actually detect? A basic aerodynamic challenge lies in the fact that nanometre-size particles tend to follow the gas flow around the payload structure rather than reaching the detector. This leads to a size-dependent and altitude-dependent detection efficiency that needs to be quantified [papers I and II]. How do we actually measure charged particles? Many in situ measurements of mesospheric particles have been limited to the charged fraction of the particle population [Rapp et al ., 2007a; paper III]. However, it has recently been speculated that impact charge measurements may be strongly influenced by secondary charging effects like fragmentation or triboelectric charging [ Barjataya et al., 2006, Havnes and Naesheim 2007]. What processes influence the electric potential of the rocket payload? Payload charging can strongly influence measurements of ionospheric properties or particles [ Holzworth et al. , 2001]. Payload charging mechanisms are today well understood when it comes to satellite platforms at higher altitudes. However, the intimate interplay of aerodynamics and charge flows makes the problem much more complicated for sounding rockets. Charging effects have puzzled scientists after many rocket flights and we are today still far from the necessary

28 answers. An important step has recently been taken by combining numerical simulations of aerodynamics, charge flows and electric fields near a sounding rocket [ Sternovsky et al ., 2004]. But the complete picture is complex and more comprehensive efforts will be necessary to quantitatively understand the interaction of rocket payloads with their environment.

Chapter 3

The PHOCUS project

Aerosol particles are currently the hottest topic in mesospheric research. Particle species comprise ice particles, smoke particles of meteoric origin, and possibly other background particles formed by conversion from trace gases. Important questions concern both the properties of particle layers and their interaction with various phenomena in the mesosphere and lower thermosphere. This includes the relationship between smoke and ice, ice particle nucleation and evolution, and the possible influence of these particle species on neutral and ion chemistry. The scientific importance of mesospheric aerosols is reflected by the fact that there are currently two major rocket efforts in Europe focussing on this topic. One is the German/Norwegian ECOMA programme from Andøya Rocket Range (section 1.2.2), the other is the Swedish PHOCUS project (Particles, Hydrogen and Oxygen Chemistry in the Upper Summer mesosphere) from Esrange which is the topic for this chapter. Both aim at the characterisation of smoke and ice particles and their interactions. While ECOMA concentrates on interactions with ionospheric processes, PHOCUS concentrates on interactions with neutral chemistry. Important for the ECOMA and PHOCUS projects is the fact that most of the scientific groups involved are contributing with instruments to both projects.

PHOCUS is funded by the Swedish National Space Board and the sounding rocket campaign is scheduled for a launch from Esrange, Sweden, in the summer of 2010. The results of this thesis provide important input to the preparation and the scientific analysis of the project. In the following the science questions and details about the individual instruments on board PHOCUS and their development will be discussed.

3.1 Science questions

With PHOCUS we intend to study the interaction between three important mesospheric players that earlier have been addressed independently of each other: meteoric smoke, noctilucent clouds, and odd oxygen/odd hydrogen chemistry. The PHOCUS project brings three lines of scientific developments at MISU together. This includes the optical probing of mesospheric chemistry, the optical characterization of NLC particles, and the study of mesospheric smoke particles within the MAGIC programme. This is complemented by our partners focusing on particle/ionosphere interactions and the distribution of water vapour. The basic science questions related to these particle interactions are summarized in Figure 7.

Figure 7. Schematic of the science questions addressed by PHOCUS.

Smoke and NLC

The direct connection between NLC and their potential condensation nuclei has never been investigated experimentally. This is a critical shortcoming for our current understanding of NLC processes. Meteoric smoke particles have been suggested as dominant condensation nuclei for ice particles [e.g., Rapp and Thomas , 2006; Plane , 2000]. The role of ions and charged particles in ice nucleation remains unclear [ Gumbel and Witt, 2002; Gumbel and Megner , 2008]. Recently, also sulphate compounds have been suggested as a new actor in mesospheric particle processes [ Mills et al .,

31 2005]. Today we neither know typical number densities nor particles sizes of various aerosol types in the summer mesosphere. Both NLC nucleation rates and ice particle properties depend strongly on these parameters [e.g., Turco et al ., 1982]. Important open questions also concern the shape of NLC particles, with growing evidence that particles are not spherical [ Rapp et al. , 2007b; Eremenko et al ., 2005].

NLC and HO x

Earlier rocket-borne measurements of atomic oxygen and NLC showed distinct O minima near the cloud altitude, leading to the idea that ice particles might influence mesospheric chemistry [ Kopp et al ., 1985; Gumbel et al ., 1998]. These questions have recently been addressed by laboratory experiments [ Murray and Plane , 2005], showing that heterogeneous processes on NLC surfaces are not efficient enough to explain the observed oxygen loss. Rather, NLC particles can influence atomic oxygen indirectly by affecting the distribution of water vapour and, thus, gas-phase HO x chemistry. A shortcoming of earlier in situ NLC studies is that all rocket flights were carried out during twilight when the solar Lyman-α flux exhibits a strong gradient exactly at NLC altitudes. This has a strong influence on the altitude-dependent chemistry and makes it difficult to distinguish particle effects from radiation effects. This is one strong reason why the PHOCUS payload should be launched during mid- day rather than in twilight.

Smoke and HO x

Influences of particle layers on mesospheric chemistry have also been discussed in connection with meteoric smoke particles [ Summers and Siskind , 1999]. Recombination of H 2 and O on these particles has been suggested to explain observations of mesospheric HO x chemistry and unexpectedly high water concentrations. A conclusion from these observations has been that the odd hydrogen/odd oxygen chemistry in the middle atmosphere is not well understood and might not be explainable with conventional gas-phase chemistry alone. This problem has been denoted as the "HO x dilemma" [ Conway et al ., 2000].

Water vapour

A fourth atmospheric player intimately connected to the above three is water vapour. A sufficient supply of water vapour is the basic pre-condition for the formation of ice particles. The availability of water vapour at the summer mesopause is determined by a delicate balance with upward

32 transport as the source and photolysis in 24-hour sunlit conditions as an efficient sink. At the same time, being the source of odd hydrogen species (HO x), water vapour together with atomic oxygen controls the chemical environment in the mesopause region. Large effects of NLC on the (re-) distribution of water vapour have earlier been reported by Summers et al . [2001] and modelled e.g. by von Zahn and Berger [2002]. With Odin's measurements, we are now able to study the distribution of water vapour on a global scale [ Lossow et al ., 2008b]. However, in situ water vapour measurements in the daytime mesosphere have so far not been available. With the recent Swedish advances in space-borne microwave technology for the Odin satellite and related projects, such measurements are now feasible as a part of PHOCUS.

Temperature

The existence of ice particles in the mesopause region is a direct consequence of the extremely low temperatures in the altitude range between 80 and 90 km. On top of the mean temperature structure, local deviations occur due to gravity. Rapp et al . [2002] showed a substantial influence of this gravity wave activity on the formation of NLC. Hence, in situ temperature measurements with high spatial resolution are needed together with accurate measurements of water vapour when NLC are to be analysed in terms of ice condensation and sublimation processes.

Ionospheric parameters

Charging processes are closely connected to a variety of layered particle phenomena in the mesosphere. The most prominent example of this interaction is the existence of polar mesosphere summer echoes [ Rapp and Lübken , 2004]. In fact, for typical ionospheric conditions, efficient negative particle charging can be expected due to capture of electrons [ Rapp and Lübken , 2001]. Even positive particle populations have been observed [Havnes et al ., 1996; paper III], although so far in lack of a conclusive explanation. Interactions between ion chemistry and the existence of particles are mutual: Particles influence ion composition and reaction schemes by capture of electrons and ions. Charges captured by particles, on the other hand, can influence further particle growth by coagulation processes [ Jensen and Thomas , 1991]. In summary, a comprehensive characterisation of particle processes in the mesosphere is not possible without a simultaneous analysis of the state of the background ionosphere.

In addition to the in situ measurements planned for the PHOCUS rocket payload, the objectives described above require complementary ground-

33 based data (NLC and PMSE monitoring) as well as detailed laboratory and model studies. The schematic overview in Figure 8 summarizes the various parts of the mission. Details about the individual instruments and their development are discussed in the next section.

Figure 8. A schematic overview over the in situ, ground-based, laboratory and model efforts contributing to the PHOCUS project from Esrange in 2010.

Table 1. Instrumentation for the PHOCUS sounding rocket. Instrument Experimenter Objective O-probe MISU atomic oxygen density H-probe MISU atomic hydrogen density 2× IR photometers MISU oxygen and hydrogen related airglow MAGIC MISU/NRL smoke particle properties 2× polarisation photometers MISU NLC particle properties phase function photometer MISU NLC particle properties 2× microwave radiometer Chalmers water vapour density 2× Faraday cups IAP and UiT charged dust particles CDD LASP charged dust particles CONE IAP/FFI temperature, density, turbulence Faraday, ion probe TU Graz electron and ion densities

3.2 Instrumentation

The backbone of the PHOCUS rocket campaign is an instrumented payload for in situ measurements of particles, atomic and molecular species as well as atmospheric background conditions. Table 1 summarizes the proposed instrumentation. All experiment concepts are discussed briefly below.

Atomic oxygen probe: The atomic oxygen measurement is based on MISU’s experience with rocket-borne composition studies by VUV resonance techniques [ Gumbel and Witt , 1997]. Resonance fluorescence of the 3S-3P triplet at 130.6 nm is used for the detection of the O profile in the mesosphere and lower thermosphere. We now plan the development of a smaller probe, focusing entirely on the sensitive resonance fluorescence technique. This is a pre-condition for obtaining a light-weight and cost- effective instrument for small rocket payloads. The direct measurement will be calibrated by a simultaneous photometric dayglow measurement of the 1 3 − (0-0) band of the O 2 (a ∆ g → X Σ g ) IR Atmospheric system at 1.27 µm as described in paper V. This dayglow emission is related to the photolysis of ozone, which during the day is in steady state with atomic oxygen and a retrieval of atomic oxygen is possible [ Mlynczak et al ., 1993]. Other laboratory or in situ techniques for calibrating resonance fluorescence payloads are in principle available, but they are complicated and expensive and they exhibit inherent uncertainties [ Dickinson et al. , 1980; Sharp , 1991; Gumbel et al. , 1998]. With the photometric calibration proposed in paper V, many of these problems are avoided.

Atomic hydrogen and Lyman-α probe: The atomic hydrogen measurement uses a similar resonance fluorescence technique, employing the 121.6 nm H Lyman-α line [ Sharp and Kita , 1987]. This technique has been used by MISU during the NLC-91 campaign. Data from these measurements have been analysed. However, due to sensitivity of the detector to background radiation at longer wavelengths no useful results could be obtained at that time. The detector developments and measurements of the diffuse daytime Lyman-α fluxes described in paper IV were intended as an important input to the preparation of the PHOCUS project. Similar to the atomic oxygen measurements, also the direct measurement of atomic hydrogen will be calibrated by photometric airglow measurements. This will be based on photometry of OH Meinel emissions in the near infrared. As the OH Meinel emission is produced by the reaction

35 between mesospheric O 3 and H, and O 3 concentrations are available from the O 2 IR Atmospheric band photometry (paper V), H concentrations can be deduced. In addition to the resonance fluorescence of H atoms [Meriwether , 1989], the detector on PHOCUS will also measure solar Lyman-α scattered by NLC particles. As discussed in section 1.2.1, this will complement the longer wavelength scattering measurements by the PHOCUS polarisation photometers and the phase function photometer described below.

NLC photometry: NLC photometry is currently the best technique available for determining altitude ranges of NLC in situ. At the same time, UV photometry allows a study of particle properties like size and shape by analysing the spectral dependence, phase function, and polarisation of the scattering [ Witt , 1969; Gumbel et al ., 2001]. On PHOCUS, we will fly a unique photometer package that for the first time will investigate all three parameters simultaneously. Two forward-viewing photometers will measure at different wavelengths and will both be equipped with polarisors. A third photometer will be tilted for viewing the sky under varying scattering angles. The design is based on photometer developments earlier flown on the DROPPS and ECOMA/MASS rocket campaigns [Gumbel et al ., 2001; Megner et al. , 2008c].

The polarisation photometers will be equipped with fixed linear polarisors. They will be mounted in forward direction with their field of view parallel to the rocket spin axis. In order to gain maximum information, measurements are needed in the Mie regime and, hence, one photometer will measure in the UV at 220 nm. The second will be centred at 450 nm in order to study the spectral dependence. The payload spin will be utilised to scan through the polarisation direction, thus providing us with the Stokes vectors I, Q and U at both wavelengths. The phase function photometer will be mounted sideways, viewing the overhead sky at an angle of 30 ° from the rocket spin axis. Due to the payload spin, NLC will be observed under varying scattering geometries as the payload approaches the cloud layer [Gumbel et al ., 2001]. In addition, as mentioned above, the detector for the H resonance fluorescence measurement will also measure the Lyman-α light at 121.6 nm scattered by ice particles thus providing NLC measurements at different scattering angles and at even shorter wavelengths.

Particle impact measurements: As a large fraction of mesospheric particles can be expected to be charged, charge-sensitive detectors can provide information about particles and their distribution (section 1.2.2). The two Faraday cup detectors contributed by IAP and the University of

36 Tromsø (UiT) are based on the original design by Havnes et al. [1996] and will measure heavy charged particles. Discrimination against light ions is provided by electrostatic shielding. In addition, the UiT detector is designed as to distinguish the generation of secondary charges upon ice particle on the detector surface [ Havnes and Naesheim , 2007]. Faraday cup measurements are discussed further in paper I and III.

Also the Colorado dust detectors (CDD) provided by the Laboratory of Atmospheric and Space Physics (LASP) will measure charged particle concentrations. They are graphite patch collectors mounted flush with the skin of the payload and measure a net current deposited on the detectors by heavy aerosol particles [e.g. Horányi et al , 2000; Amyx et al. , 2008]. For the CDD’s discrimination against light ions is provided by magnetic shielding.

Particle sampling: Direct sampling of meteoric smoke particles will be made with the MAGIC meteoric smoke sampler (section 1.2.2). The aim is a comprehensive characterisation of nanometre-size meteoric particles. With a new facility for electron microscopy at Stockholm University, good opportunities for the laboratory analysis of the PHOCUS/MAGIC results are now available in Sweden. The sampling of meteoric smoke particles by the MAGIC meteoric smoke sampler is discussed further in paper II.

Water vapour measurements: The concept of a rocket-borne microwave radiometer for water vapour is being developed by Chalmers University of Technology. Based on substantial recent advances in microwave technology [ Murtagh et al., 2002], the measurement will be based on the 183 GHz and 558 GHz water lines. It is today fully feasible to build a suitable mm-wave instrument well within the size of a rocket module and without the need for active cooling. In fact, long-going design work for such an instrument is already available at Chalmers in preparation of the Atacama Large Millimeter Array (ALMA). The instrument for PHOCUS will utilise one forward-looking antenna, thus detecting water vapour emissions from the atmospheric column above the payload, and one side- looking antenna for high altitude resolution. It is interesting to note that already in the early 1990s a similar concept for mesospheric water measurements was proposed by Croskey et al . [1993]. However, only now has microwave technology advanced sufficiently to make an easy incorporation of such an instrument into a sounding rocket feasible. The perspective of obtaining local measurements of water vapour from the polar summer mesosphere is extremely attractive to the entire middle atmosphere science community.

Temperature measurements: The Combined Neutral and Electron sensor (CONE) is contributed by the Leibniz-Institute for Atmospheric Physics (IAP) and the Norwegian Defence Research Establishment (FFI). Basically, CONE is an ionisation gauge for the neutral density measurements. Temperatures are accurately derived from the neutral density via the hydrostatic equation. CONE is today the only instrument available that can resolve temperature structures in the mesopause region down to the level of gravity waves [ Rapp et al ., 2002]. It also provides measurements of turbulence on scales down to 10 cm [ Lübken et al ., 2002]. CONE has been flown before from Esrange on the NLTE payloads in 1998 [ Lübken et al ., 1999].

Ionospheric conditions: Ionospheric instrumentation will be provided by the Technical University Graz. The instrument package consists of a Faraday Radio Wave Propagation experiment (RWP) for high-accuracy electron density measurements and an ion probe for ion densities. The RWP uses four radio waves in the HF-range (1.3 to 7.8 MHz) that are transmitted from the ground to the flying rocket payload. Electron densities are inferred from the change of the polarisation without influences of payload charging or aerodynamics [ Jacobsen and Friedrich , 1979]. Positive ions are measured with good altitude resolution by a gridded sphere with a positively biased collector inside [ Thrane , 1986]. In combination, this instrument package provides both absolute calibration and high sensitivity/altitude resolution. By separately measuring number densities of electron and light positive ions, the charge balance is studied and detailed information on both ionospheric conditions and ion/particle interactions can be obtained.

Chapter 4

Results of this thesis

Paper I: This paper concerns the aerodynamics of rocket-borne meteoric smoke particle impact measurements. Basically, the small smoke particles tend to follow the gas flow around the payload rather than reaching the detector if aerodynamics is not considered carefully in the detector design. Quantitative ways are needed that relate the measured particle population to the atmospheric particle population. This requires in particular knowledge about the size-dependent, altitude-dependent and charge-dependent detection efficiency for a given instrument. In this paper, we investigate the aerodynamics for a typical electrostatic detector design by first quantifying the flow field of the background gas and then introduce particles in the flow field to determine their trajectories around the payload structure. We use two different models to trace particles in the flow field, a Continuous motion model and a new Brownian motion model. Detection efficiencies are determined for three detector designs, including two with ventilation holes to allow airflow through the detector. Results from this investigation show that Brownian motion is of basic importance for the smallest particles as they can be completely decelerated by the stagnating airflow inside the detector and their further motion is governed by Brownian diffusion. Results also show that the detection efficiency for meteoric smoke particles is very much size and altitude dependent for this kind of particle impact measurements. Particles have to be larger than 1 - 2 nm in radius to be detected at any altitude between 70 and 95 km. The altitude range 70 - 75 km is an aerodynamical lower detection limit for particles up to at least 5 nm. It is important to note that earlier measurements of sharp cut-offs in particle concentrations at these altitudes have been interpreted as a real geophysical feature rather than instrumental in nature. Providing venting holes allowing air to flow through the detector can improve the detection

39 efficiency at lower altitudes. However, the flow has to be relatively large in order to significantly improve the detection efficiency.

Paper II: This paper discusses rocket-borne sampling of neutral meteoric smoke particles and the detection efficiency of the MAGIC meteoric smoke samplers. The MAGIC project (Mesospheric Aerosol – Genesis, Interaction and Composition) aims to quantitatively answer fundamental questions about the properties of smoke particles in the atmosphere. This is pursued with the development of an instrument and analysis techniques that for the first time allow smoke particles to be brought back from the mesosphere into the laboratory for a detailed study of their properties. The idea with the MAGIC smoke samplers is to minimise the aerodynamic perturbations in the measurement volume by reducing the detector size to the order of the molecular mean free path. The sampling surfaces are 3 mm diameter carbon grids for transmission electron microscopy (TEM) mounted on top of pins that are extended from the payload top and reach outside the payload shock front. When the samples are returned to ground these grids can be moved directly into the TEM for analysis. In order to characterise the sampling process, we have performed simulations of the trajectories of nanometre- sized smoke particles towards the MAGIC detectors with the Brownian motion model described in paper I. As a result, we obtain the detection efficiency for the MAGIC detectors as a function of altitude and particle size. Our simulations confirm that particles of radii down to 0.75 nm impact on the sampling surface with an efficiency exceeding 80% over the entire mesospheric altitude range of interest.

Paper III: In this paper we present results of in situ measurements of charged nanoparticles, electrons, and positive ions obtained during a sounding rocket flight in October 2004 from Kiruna, Sweden, under nighttime conditions. The particle measurement reveals positive charge signatures in the altitude range between 80 and 90 km corresponding to peak charge number densities of ~100 e/cm 3 at around 86 km. Aerodynamical analysis of the sampling efficiency of our instrument using the Continuous motion model described in paper I reveals that the particles must have been larger than 2 nm assuming spherical particles with a density of 3 g/cm 3. The plasma environment of the observed particles is dominated by negative and positive ions, with only few free electrons. A calculation of the mean particle charge expected for particles in a plasma consisting of electrons and positive and negative ions shows that the presence of sufficiently heavy and numerous negative ions (i.e., m n > 300 amu and λ ≥ 50) can explain the observed positive particle charge.

Paper IV: This paper concerns the study of mesospheric ice particles by optical methods. In order to better understand noctilucent clouds (NLC) and their sensitivity to the variable environment of the polar mesosphere, more needs to be learned about the actual cloud particle population. Optical measurements are today the only means of obtaining information about the size of mesospheric ice particles. In order to efficiently access particle sizes, scattering experiments need to be performed in the Mie scattering regime, thus requiring wavelengths of the order of the particle size. We show that NLC measurements in the extreme ultraviolet, in particular using solar Lyman-α radiation at 121.57 nm, are an efficient way to further promote our understanding of NLC particle size distributions. Here, we present examples from recent rocket-borne studies that demonstrate how ambiguities in the size retrieval at longer wavelengths can be removed by invoking Lyman-α. We discuss basic requirements and instrument concepts for future rocket-borne NLC missions. In order for Lyman-α radiation to reach NLC altitudes, high solar elevation and, hence, daytime conditions are needed. Considering the effects of Lyman-α on NLC in general, we argue that the traditional focus of rocket-borne NLC missions on twilight conditions has limited our ability to study the full complexity of the summer mesopause environment.

Under the 6 th Framework Program of the European Union, the European Commission has promoted the utilization of the Arctic Lidar Observatory for Middle Atmosphere Research (ALOMAR) at Andøya Rocket Range, Norway. In particular, the enhanced ALOMAR Research Infrastructure (eARI) included two sounding rocket launches in 2006 and 2008. A rocket- borne instrument to measure the scattering of Lyman-α by ice particles was developed and launched on the first of these sounding rockets. The instrument is a double monochromator using two 3600 lines/mm holographic gratings to give a good selectivity of the Lyman-α line and provide a blocking of other wavelengths by 6 - 8 orders of magnitude. Unfortunately, a malfunction of the payload shortly after launch made it impossible to test the instrument concept and collect any mesospheric data.

Paper V: In this paper the possibility to calibrate direct atomic oxygen measurements by simultaneous airglow measurements is presented. Accurate knowledge about the distribution of atomic oxygen is crucial for many studies of the mesosphere and lower thermosphere. Direct measurements of atomic oxygen by the resonance fluorescence technique at 130.6 nm have been made from many sounding rocket payloads in the past

41 with good sensitivity and altitude resolution. However, accuracy is a problem as calibration and aerodynamics make the quantitative analysis challenging. In general, accuracies better than a factor 2 are not to be expected from direct atomic oxygen measurements. As an example, we present results from the NLTE sounding rocket campaign at Esrange, Sweden, in 1998, with simultaneous O 2 airglow and O resonance fluorescence measurements. O number densities are found to be consistent with the nightglow analysis, but only within the uncertainty limits of the resonance fluorescence technique. Based on these results, we describe how better atomic oxygen number densities can be obtained by calibrating direct techniques with complementary airglow photometer measurements and detailed aerodynamic analysis. Night-time direct O measurements can be 1 complemented by photometric detection of the O 2 (b Σ) Atmospheric Band 1 at 762 nm, while during daytime the O 2 (a ∆) Infrared Atmospheric Band at 1.27 µm can be used. The combination of a photometer and a rather simple resonance fluorescence probe can provide atomic oxygen profiles with both good accuracy and good height resolution.

Chapter 5

Outlook

As for the future, mesospheric particles will remain in the focus of research in the middle atmosphere. The research with sounding rockets will address the small-scale interactions and processes needed in order to understand the atmosphere as a whole. More needs to be learned about the particles microphysics and interactions. This is particularly important since mesospheric particle phenomena are considered to be a sensitive indicator for the variability and long-term trends of the middle atmosphere. In order to understand their role, the size spectrum and composition of both ice and smoke particles needs to be quantified. Further development of direct sampling techniques will hopefully open new ways to study these particles.

It is very important to be able to relate the measured particles to the actual particle population in the mesosphere. There are a number of improvements that can be made to the particle motion models described and used in papers I, II, and III. The Brownian motion description can be further developed to include a more realistic treatment of the molecule/particle collisions. Both inelastic collisions and non-spherical particles can be considered. The mass loss due to heating and subsequent sublimation in the shocked gas flow in front of a detector is negligible for smoke particles and large ice particles. However, for smaller ice particles this becomes important and must be included in the model if ice particles are to be traced. Also neglected in the model at this point are effects of payload charging by photons, ions, and particles, as well as the possibility that incident particle trajectories are influenced by such a payload charging. Another important development is to extend the model into 3 dimensions to be able to simulate rocket payloads with angles of attack other than 0° or payloads that are asymmetric. The Brownian motion model can be used to quantify the

43 possible contamination of the Hygrosonde water measurement [ Khaplanov et al. , 1996; Lossow et al. , 2008a] due to adsorption or outgassing from the rocket payload by studying how water vapour diffuses in the aerodynamically shocked region in front of the instrument.

For a better understanding of mesospheric ice particle phenomena it is important to develop new optical techniques to efficiently access the smaller particles in the ice particle population. In paper IV we show how this can be done by extending scattering experiments to be performed in the Mie scattering regime, using wavelengths of the order of the particle size. Due to a fortunate coincidence, the wavelength of the strong solar Lyman-α line at 121.57 nm coincides with a minimum in the absorption spectrum of O2 that otherwise efficiently shields the atmosphere from UV radiation in the wavelength range 100-180 nm. As a result, direct Lyman-α penetrates deep into the atmosphere during the day, well below the ice clouds in the mesosphere. Further development of the SLAM instrument and incorporation on sounding rocket payloads together with longer wavelength measurements would give new and important information about the polar mesospheric ice layer.

The new concept presented in paper V of calibrating direct atomic oxygen measurements with simultaneous photometric O 2 airglow measurements is promising. In order to further improve the accuracy of this method, it is important to further investigate the rate constants and excitation parameters related to the airglow emissions. The coefficients currently in use are to a large extent based on data from the single rocket campaign ETON. A future, independent sounding rocket project with dedicated instruments to study both atomic oxygen and oxygen related airglow is highly desirable.

For future rocket projects, the development of small and “cost-effective” instrumentation will be mandatory. The concepts described in this thesis are currently part of the preparation for the PHOCUS rocket campaign from Esrange in 2010. They are also part of the continuously ongoing process of providing improved experimental tools for the scientific understanding of our atmosphere.

Acknowledgements

When I first started my Ph. D. position at MISU, five years felt like a very long time. However, as you probably know, time flies when you do fun and interesting things. These past five years has not been an exception and the time has gone by very fast, maybe even too fast. I thank everyone at MISU for making this part of my life very pleasant and memorable.

I wish to express my deepest and most sincere appreciation towards my supervisor Jörg Gumbel for the inspiration and guidance through the different projects and for being a very good friend. Many thanks also to my assistant supervisor Markus Rapp at the Leibniz Institute of Atmospheric Physics in Kühlungsborn, Germany. I am deeply grateful for being part of the AP group together with Jörg Gumbel, Jacek Stegman, Misha Khaplanov, Georg Witt, Bodil Karlsson, Stefan Lossow, Linda Megner, Farah Khosrawi, Heiner Körnich, Tomas Waldemarsson and Peggy Achtert. You have all certainly provided a wonderful and stimulating working atmosphere. I have very much enjoyed all our travels, campaigns and dinners (and of course work) together. I have learned so much from you all, especially Jörg, Misha, Jacek and Georg. Thank you for sharing your deep knowledge!

Special thanks also to Leif Bäcklin and Nils Walberg for the help in the machine shop, to Eva Tiberg for knowing and taking care of almost everything, to Birgitta Björling for help with finding old publications, to Solveig Larsson, Pat Johansson, Susanne Eriksson, and Janne Söderlund for help with a variety of things, and to Mikael Burger for the computer support. You are all really a very important part of the department.

I wish to thank everyone at Andøya Rocket Range, Norway, and Swedish Space Corporation at Esrange and in Solna for making my visits there

45 during meetings, experimental tests and campaigns very memorable. Many thanks also to Karl Heinrich Fricke for the instructive time at Esrange.

I thank the Swedish National Graduate School of Space Technology for the support in funding my Ph. D. studies.

And of course I thank Jonathan Friedman for introducing me to the middle atmosphere during my M. Sc. thesis work at the National Astronomy and Ionosphere Center, Arecibo Observatory, Puerto Rico. This certainly inspired me to continue in the field of atmospheric science.

Last but not least, I thank my family and friends for supporting and believing in me, and of course my Jenny, the worlds greatest, for everything.

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