measurements of stratospheric aerosol extinction coefficients by

Faculty of Physics and Astronomy

University of Heidelberg

Diploma thesis

in Physics

submitted by

Martin Hirsekornborn in Braunschweig

– 2001 –

Measurements of StratosphericAerosol Extinction Coefficients byBalloon-Borne DOAS Observations

This Diploma Thesis has been carried out by Martin Hirsekorn at the

Institute for Environmental Physics

of the

University of Heidelberg

under the supervision of

Priv.-Doz. Dr. Klaus Pfeilsticker

Abstract

A method has been developed to infer stratospheric aerosol extinction coefficients from balloon-borne solar occultation measurements. The data were recorded by a DOAS spectrograph, pri-marily constructed to measure stratospheric trace gas profiles. The aim is to simultaneouslyderive aerosol surface areas, which can be used as input parameters for modelling the hetero-geneous stratospheric chemistry. Furthermore, satellite-borne aerosol extinction measurementscan be validated. The aerosol extinction is examined over a continuous wavelength range from440 to 615 nm.

Between November 1996 and February 2000, eight balloon-flights were conducted in collab-oration with the French LPMA team (Laboratoire de Physique Moleculaire et Applications) athigh and mid-latitudes. The obtained vertical aerosol extinction profiles are compared with datarecorded by the satellite-borne instruments SAGE II and POAM III. The results show that theaerosol extinction measured by satellites is about a factor of two lower than the one measuredby the DOAS spectrograph for all spatially and temporally coinciding observations. The shapeof the vertical profiles, however, is in good agreement. A decrease of the particle size with in-creasing altitude is derived from the data, being in agreement with satellite and optical particlecounter observations.

The analyses of the observed wavelength dependence indicate a significant amount of verysmall particles not detected by these particle counters. If this is true a large increase of theaerosol surface area with respect to common models are suggested by the DOAS measurementsemphasizing more heterogenous chemistry in the stratosphere.

Zusammenfassung

Es wurde eine Methode entwickelt zur Bestimmung stratospharischer Aerosol-Extinktionskoeffizienten mit Hilfe ballongestutzter Sonnenokkultationsmessungen. Dafurwurde ein DOAS-Spektrograph zur Bestimmung stratospharischer Spurenstoffprofile ver-wendet. Das Ziel ist die Gewinnung simultan gemessener Aerosoloberflachendaten, die alsEingabeparameter fur heterogene Chemiemodelle verwendet werden konnen. Daruberhinausbietet sich die Moglichkeit der Validierung satellitengestutzter Aerosol-Extinktionsmessungen.Die Aerosol-Extinktion wurde uber einen zusammenhangenden Wellenlangenbereich zwischen440 und 615 nm untersucht.

Von November 1996 bis Februar 2000 wurden in Zusammenarbeit mit der franzosischenGruppe vom LPMA (Laboratoire de Physique Moleculaire et Applications) acht Ballonfluge inmittleren und hohen geographischen Breiten durchgefuhrt. Die erhaltenen Vertikalprofile wur-den mit Messungen der Satelliteninstrumente SAGE II und POAM III verglichen. Dabei zeigtesich, daß die von den Satelliten gemessene Aerosolextinktion fur alle Vergleichspunkte um etwaeinen Faktor Zwei niedriger ist als die aus den DOAS-Messungen erhaltene. Der relative Verlaufdes Hohenprofils zeigt jedoch gute Ubereinstimmungen. Auch konnte eine mit der Hohe ab-nehmende Teilchengroße festgestellt werden, was mit den Daten der Satelliteninstrumente undvon optischen Partikelzahlern ubereinstimmt.

Untersuchungen der gemessenen Wellenlangenabhangigkeiten weisen auf eine große Mengean sehr kleinen Aerosolen hin, die von den Partikelzahlern nicht detektiert werden. Eine durchdiese kleinen Teilchen verursachte starke Erhohung der Aerosoloberflache gegenuber bisherigenModellen hatte weitreichende Konsequenzen auf das Verstandnis der heterogenen Stratospharen-chemie.

i

Contents

1 Introduction 1

2 Formation of Aerosols 5

2.1 Various Types of Aerosols . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5

2.2 Sulfate Aerosols in the Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . 7

2.3 The Importance of Aerosols in the Stratosphere . . . . . . . . . . . . . . . . . . . 9

2.3.1 The Radiation Budget . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9

2.3.2 Heterogeneous Chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . 10

2.4 Elementary Stratospheric Dynamics . . . . . . . . . . . . . . . . . . . . . . . . . 13

2.5 Volcanic Aerosol . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15

2.6 Background Aerosol . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17

3 Optical Properties of Stratospheric Aerosols 19

3.1 Measurement Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19

3.1.1 In-situ Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19

3.1.2 Stray Light Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . 20

3.1.3 Extinction Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . 21

3.2 Rayleigh Scattering . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23

3.3 The Mie Theory . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24

3.4 Size Distributions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29

3.5 Aerosol Extinction Coefficients . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32

4 Balloon Borne DOAS of Direct Sunlight 37

4.1 Light Attenuation in the Atmosphere . . . . . . . . . . . . . . . . . . . . . . . . . 37

4.2 Differential Optical Absorption Spectroscopy (DOAS) . . . . . . . . . . . . . . . 39

4.3 Retrieval of Vertical Trace Gas Profiles . . . . . . . . . . . . . . . . . . . . . . . . 40

4.3.1 The Slant Column Density . . . . . . . . . . . . . . . . . . . . . . . . . . 42

4.3.2 Error Sources of the DOAS Retrieval . . . . . . . . . . . . . . . . . . . . . 44

4.3.3 Raytracing . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46

4.3.4 AMF Matrix Inversion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47

4.4 The DOAS Balloon Spectrograph . . . . . . . . . . . . . . . . . . . . . . . . . . . 49

4.4.1 The LPMA/DOAS Balloon Gondola . . . . . . . . . . . . . . . . . . . . . 49

4.4.2 Optical and Mechanical Properties of the Spectrograph . . . . . . . . . . 49

4.4.3 Measurements of Relative Intensities . . . . . . . . . . . . . . . . . . . . . 53

ii

5 Retrieval of Vertical Aerosol Extinction Profiles 575.1 Spectrograph Stray Light Correction . . . . . . . . . . . . . . . . . . . . . . . . . 605.2 Elimination of Trace Gas Absorptions . . . . . . . . . . . . . . . . . . . . . . . . 605.3 Removal of Rayleigh Extinction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 625.4 Zero Air Mass Correction (Langley Plot) . . . . . . . . . . . . . . . . . . . . . . . 655.5 Error Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66

6 Results and Comparison with Satellite Data 716.1 Vertical Aerosol Extinction Profiles . . . . . . . . . . . . . . . . . . . . . . . . . . 726.2 The SAGE and POAM Aerosol Retrieval Algorithm . . . . . . . . . . . . . . . . 776.3 Comparison of the DOAS with Satellite Aerosol Extinction Data . . . . . . . . . 80

7 Inferred Mie Scattering Wavelength Dependence 857.1 Observed Wavelength Dependencies . . . . . . . . . . . . . . . . . . . . . . . . . 857.2 Calculation of the Mie Scattering Wavelength Dependence . . . . . . . . . . . . . 867.3 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88

7.3.1 Mie Scattering Wavelength Dependence . . . . . . . . . . . . . . . . . . . 887.3.2 Apparent Size Distribution inferred for the DOAS Observations . . . . . . 907.3.3 Comparison with Auxiliary Data . . . . . . . . . . . . . . . . . . . . . . . 92

7.4 Suggestions and Recommendations for Future Investigations . . . . . . . . . . . . 94

8 Conclusion and Outlook 95

Bibliography 97

iii

Chapter 1

Introduction

Historical reports on widespread optical phenomena in the atmosphere following large volcaniceruptions illustrate the significant impact of airborne particles of different origins on the Earth’sradiation budget and climate. After the cataclysmic explosion of Krakatau on August 26 and 27,1883, a number of unusual observations were made all over the Earth, in particular at low lati-tudes. Immediately after the eruption, white glare in the sky, streaky haze, spectacular sunrises,and an unusually colored Sun and Moon were reported from Krakatau’s vicinity [Symons 1888].These phenomena were followed by prolonged purple and red twilight glows and by Bishop’srings surrounding the Sun [Bishop 1884]. The nightly duration of the twilight glows indicated,that the tropical glow layer dwelled at altitudes of about 30 km [Symons 1888; Pernter 1889].By January 1884, the main glow layer had sunk to about 17 km. The cloud responsible forthese peculiarities was spread around the Earth by stratospheric winds within a few weeks, butremained limited to low latitudes.

Associated with the Krakatau haze was a remarkable 3-year reduction in ground level solarradiation [Pernter 1889]. Published data on stellar visibilities indicate that the atmosphericextinction was increased by about 1-2 mag during the period between January 1884 and thesummer of 1885 [Gore 1884; Winlock 1884].

Similar observations were made after several volcanic eruptions during the past century[Stothers 1996], whereas other eruptions did not give rise to anything similar to the dramaticeffects reported after Krakatau’s eruption.

Major volcanic eruptions eventually inject significant amounts of sulfur gases (mostly OCS,CS2, SO2, and H2S), sulfur containing aerosols, and small silicate ash particles directly intothe stratosphere. The sulfur-bearing gases are then photochemically converted into sulfuricacid within several weeks. Together with water vapor, sulfuric acid condenses into the wellknown Junge aerosols [Toon and Pollack 1973; Castleman et al. 1973]. These aerosols remainsuspended in the stratosphere with a typical time scale of 1.5 years, while the large ash particlesgravitationally settle out within a few days or weeks. Accordingly, after the last major volcaniceruption of Mt. Pinatubo in June 1991 a background level in sulfate aerosol has been reachedin the stratosphere by late 1996 [Barnes and Hofmann 1997 and 2001].

Stratospheric sulfate aerosols are known for its effects on Earth’s radiative budget and viaheterogeneous reactions on stratospheric ozone. For example, the importance of volcanic activityrelated stratospheric aerosols on the Earth’s surface temperature was already recognized afterthe eruption of Tambora in 1814 by record low temperatures occurring in Europe in summer1815.

1

2 CHAPTER 1. INTRODUCTION

The discovery of the Antarctic ozone hole by Farman et al. [1985] drew attention on thechemical aspects of stratospheric aerosols. Today heterogenous chemical reactions on PSC1-particles and cold sulfate aerosols are accepted to precondition the dramatic ozone depletionin both polar spring stratospheres [Solomon 1988]. Thereby, the most important heterogeneousprocess is the conversion of reactive nitrogen species into nitric acid in combination with theactivation of chlorine and bromine bearing reservoir gases (HCl, ClONO2, BrONO2) into ozonedestroying halogen oxides [e. g. WMO 1999]. Recent research has demonstrated that hetero-geneous reactions on the moderately cold sulfate aerosols also exert influence on mid-latitudestratospheric ozone, however more widespread but less intense than over the cold polar re-gions [e. g. Fahey et al. 1993; WMO/UNEP 1994; Avallone et al. 1993; Hofmann et al. 1994;Erle et al. 1998].

Physical properties of the stratospheric aerosols are either measured by in-situ or remote-sensing techniques. While in-situ techniques are best suited to measure the aerosol size distribu-tion and the aerosol’s chemical composition, remote sensing techniques are better in monitoringthe aerosol’s optical properties. The link between both sets of measurements is then eventuallyprovided by applying Mie theory to calculate either for measured optical properties and as-sumed chemical composition the size distribution or vice versa. An established in-situ techniqueto measure the aerosol size distribution is selective particle counting by instruments deployed onstratospheric balloons [e. g. Hofmann et al. 1975; Hofmann and Rosen 1983; Deshler et al. 1993;Ovarlez and Ovarlez 1995; Deshler et al. 1997]. Remote-sensing techniques either rely on activeinstruments like LIDARs2 [e. g. Abo and Nagasawa 1994; Iwasaka et al. 1994; Yasui et al. 1994;Barnes and Hofmann 1997; McCormick et al. 1995] or passive instruments that use directSun, moon or star light or scattered skylight [e. g. Herman et al. 1986; McCormick 1987;Rao et al. 1989; Farmer 1989; Russell III et al. 1993; Renard et al. 1996; Bevilacqua 1997].

Past passive techniques primarily observed the optical properties, however, only in limitedand separated wavelength intervals located in the UV to mid-IR spectral region. These observa-tions allow to infer Mie extinction in discrete wavelength intervals and from that accordingly thewavelength dependence of the aerosol extinction. If properly chosen the observation of discretewavelength intervals may then also provide information on interfering trace gas absorptions[e. g. Chu et al. 1989; Lumpe et al. 1997; Lucke et al. 1999]. Conversely, interfering trace gasabsorptions and Rayleigh scattering may introduce errors into the Mie extinction retrieval if notcorrectly accounted for.

In this study (in German called Diplomarbeit), the aerosol extinction is inferred from datarecorded by a balloon-borne DOAS spectrograph that monitored the rising or setting sun fromaltitudes well above the top of the Junge layer during a set of observations at different geo-physical conditions [e. g. Ferlemann et al. 1998b; Fitzenberger et al. 2000a; Bosch 2001]. Theseobservations offer a particular advantage to infer the stratospheric aerosol extinction in that theyprovide continuous spectra monitored in wavelength ranges 320 – 422 nm and 417 – 670 nm, re-spectively, each covered by 1024 channels. The employed DOAS technique allows us to measurevery precisely absorptions of minor stratospheric species (O3, NO2, BrO, OClO, O4, H2O, IO,...) with optical densities as low as 10−4. Accordingly, by removing the precisely measured tracegas absorptions from the total extinction, the remaining extinction can be interpreted as due toRayleigh and Mie scattering. This approach is investigated in detail in the present study.

1Polar Stratospheric Cloud2LIDAR: Light Detecting And Ranging

3

Chapter 2 provides an overview on the different types of stratospheric aerosols, their origin,formation, evolution, and impact on the Earth’s climate and the stratospheric photochemistry.In chapter 3 the optical properties of stratospheric sulfate aerosols are discussed with a par-ticular emphasis put on various employed measurement methods. Major features of the DOAStechnique aiming at the precise measurement of trace gas concentrations is reported in chapter4, and a description of the actual spectrograph used for the present observations is included.Chapter 5 details the retrieval algorithm of aerosol extinction profiles, including an analysis ofits error sources. In chapter 6 the results are compared with satellite-borne aerosol extinctionmeasurements (SAGE II and POAM III). Finally, in chapter 7 the inferred wavelength depen-dence of the visible aerosol extinction is discussed and compared with visible/near-IR satelliteand optical particle counter data. Chapter 8 concludes and summarizes the findings.

4 CHAPTER 1. INTRODUCTION

Chapter 2

Formation of Aerosols

Aerosols are small airborne particles of diameters from several nanometers up to 100 microns.They occur in the troposphere and stratosphere in liquid or solid state, or as liquid dropletscontaining a solid nucleus.

New aerosol droplets are produced by two different mechanisms

- Heterogeneous nucleation on solid materials (e.g. mineral dust, sea salt spray, soot inthe troposphere or volcanic ash and soot, micrometeorites, aluminium oxide in the tro-posphere and stratosphere [Turco et al. 1982; Zolensky et al. 1989; Blake and Kato 1995;Pueschel 1996]).

- Homogeneous condensation of supersaturated gases [Deshler et al. 1992;Sheridan et al. 1992]. These vapors can also arise from atmospheric gases by chemi-cal reactions (e.g. sulfuric acid vapor from sulfur dioxide).

2.1 Various Types of Aerosols

Tropospheric Aerosols

Tropospheric aerosols form a much more complex system than aerosols in the stratosphere. Asubstantial part of them are anthropogenic. Once in the atmosphere they may be transportedup to several thousand kilometers away from their point of origin before being removed by bothdry (sedimentation) and wet processes (rainout).

Anthropogenic and natural sulfur is a principal source of homogeneous condensed aerosols[section 2.2]. 70% of the atmospheric sulfur is of anthropogenic origin, 30% is produced bynatural sources, in particular emissions from oceans and volcanoes. Currently anthropogenicsources emit about 70-90 Mt of sulfur, which is equivalent to 150-180 Mt of sulfur dioxide (SO2)[Roedel 1992]. Approximately a quarter of the disposable sulfureous gases is oxidized to sulfuricacid (H2SO4).

The other important source of aerosols are particles from the Earth’s surface. The composi-tion of the emerging droplets is heavily dependent on their place of origin. Sea salt spray fromthe oceans can act as condensation nuclei as does windblown mineral dust from arid and semi-arid regions (e.g. the Sahara), and soot accrued from the combustion of biogenic matter (coal,petroleum, wood, ...).

5

6 CHAPTER 2. FORMATION OF AEROSOLS

Figure 2.1: Possible generation mechanisms of PCSs: In this plot are shown five scenarios of NAT for-mation. Circular symbols denote liquid particles, while square symbols are denoting solid particles.Adopted from Carslaw et al. [1999]

Stratospheric Aerosols

The most prevalent aerosol in the stratosphere is composed of sulfuric acid H2SO4 and water. Theparticle concentration is continuously decreasing at higher altitudes with a maximum of the largeparticles (0.1 µm and more) a few kilometers above the tropopause. Since the large particlescontain most of the aerosol mass, there is a definite aerosol layer in the stratosphere, calledJunge Layer [Junge et al. 1961]. As such a layer cannot be generated by mixing procedures ordiffusion, the aerosols must form in the stratosphere itself. That means that not only aerosols areinjected in the stratosphere through the tropical tropopause, but also relatively stable particlesource gases, which photodissociate in the stratosphere and form new particles by homogeneouscondensation of the chemical products.

Also possible is the direct injection of the particle source gases, soot or volcanic ash in thestratosphere by large sulfur-rich volcanic eruptions [section 2.5].

2.2. SULFATE AEROSOLS IN THE STRATOSPHERE 7

The other major class of stratospheric aerosols are the Polar Stratospheric Clouds (PSCs),which occur in very cold regions within the polar vortex. Primarily they were assumed to becomposed of water ice crystals. These PSCs, nowadays known as Type II PSCs, exist at tem-peratures below 188 K. Satellite measurements evidenced that PSCs are found at temperaturesof up to 195 K as well [McCormick et al. 1982], which are significantly higher than the freezingpoint of water. Crutzen and Arnold [1986] as well as Toon et al. [1986] independently suggestedthat the PSC particles may be composed of solid Nitric Acid Trihydrate (NAT), HNO3·3H2O,labelled as Type I PSCs. Hanson and Mauersberger [1988] showed the freezing point of suchcompounds to be approximately 5 K above the one of water. At temperatures below 210 Ksulfuric acid aerosol, in the form of sulfuric acid tetrahydrate (SAT), can freeze out and act ascondensation nuclei for NAT (T < 188 K) or ice (T < 188 K).

Type I PSCs are divided into two subclasses, as shown by LIDAR measurements of aerosolproperties. Type Ia PSCs consist of solid aspherical and therefore depolarizing particles, TypeIb PSCs consist of liquid spherical particles.

2.2 Sulfate Aerosols in the Stratosphere

There are several natural sources of atmospheric sulfur. The decomposition of organic mattercontinuously produces sulfur containing compounds, such as hydrogen sulfide H2S. Dimethylsulfide (CH3)2S and CS2 are emitted by the oceans. An important amount of atmospheric sulfur,bonded in sulfur dioxide and H2S is due to volcanic eruptions. The eruptions of El Chichon(southern Mexico in spring 1982) and Mount Pinatubo (Philippines, June 1991) considerablyincreased the amount of sulfur and consequently the aerosol amount in the atmosphere. Suchpowerful events can also induct sulfureous compounds and soot directly into the stratosphere.

Anthropogenic sources play an important role since the industrial emissions increased dra-matically during the past century. Large amounts of SO2 are produced by combustion of coalespecially in electric power plants, as well as smaller quantities of H2S, CS2 and carbonyl sulfideOCS. Sulfuric acid vapors and sulfur oxides are produced in all kinds of combustion engines,e.g. in automobiles.

Thus, sulfur has always been present in the atmosphere, but its concentration strongly de-pends on anthropogenic and geophysical influences, especially volcanic eruptions. In the strato-sphere and upper troposphere the lifetime of the compounds is an important factor. H2S and(CH3)2S have a very short lifetime in the troposphere and therefore are rapidly destroyed. Thetropospheric lifetime of SO2 is several days, so that it can be observed near the tropopause[Jaeksche et al. 1976]. Since OCS is very stable in the troposphere with a lifetime around oneyear, the concentration is almost constant up to the tropopause [Brasseur and Solomon 1986].The oxidation of OCS is believed to be the major nonvolcanic source of stratospheric sulfur[Crutzen 1976].

In the stratosphere, OCS is destroyed by photolysis [Brasseur and Solomon 1986]:

OCS + hν → CO + S (2.1)

Atomic sulfur immediately reacts with molecular oxygen:

S + O2 → SO + O (2.2)


The product is sulfur monoxide which can also be produced by the reaction of OCS and atomicoxygen:

OCS + O → CO + SO (2.3)

Sulfur monoxide is in rapid equilibrium with sulfur dioxide. The dominant reaction is:

SO + O2 → SO2 + O (2.4)

Less important [Black et al. 1982aa and 1982bb] are

SO + O3 → SO2 + O2 (2.5)

andSO + NO2 → SO2 + NO. (2.6)

Sulfur dioxide is quite stable in the stratosphere as well. It can be photodissociated by radiationin the 200-230 nm region [Brasseur and Solomon 1986]

SO2 + hν → SO + O, (2.7)

oxidized to sulfur trioxide by the following processes,

SO2 + HO2 → SO3 + OH (2.8)

SO2 + OM−→ SO3 (2.9)

or converted to HSO3SO2 + OH

M−→ HSO3 (2.10)

which itself is destroyed by the following reactions

HSO3 + O2 → HO2 + SO3 (2.11)

HSO3 + OH → H2O + S. (2.12)

SO3 can combine with water to sulfuric acid H2SO4. The acid can be photodissociated or pre-cipitate with water (as can SO2 and HSO3) to form sulfate aerosols.

Sulfuric acid is the most important substance to form new condensation nuclei. Its saturationvapor pressure is extremely low (around 3.3 · 10−5hPa at 23C [Roedel 1979]) and decreasesup to a factor of 100 if it is solved in water. The resulting large supersaturation of sulfuricacid increases the probability of precipitation into small droplets. Once existing, the size ofthe droplets rapidly increases due to nucleation of water vapor until an equilibrium with thesurrounding water vapor is reached [Steele and Hamill 1981]. Since in the case of sulfuric acidthe content of the surrounding water vapor is scarcely affected, the particle absorbs water untilit reaches a composition in which the water vapor pressure is equal to the surrounding waterpartial pressure. Thus, the mixing ratio of the droplets depends primarily on the humidityand the temperature. For typical stratospherical conditions (H2O partial pressure 10−4 – 10−3

mbar, 220 – 250 K) the weight percentage of H2SO4 is within the narrow range of 68-86%[Steele and Hamill 1981].

Another particle source gas is nitrous oxide (N2O) which is the principal precursor of nitricacid (HNO3). The main anthropogenic emissions of N2O are agricultural emissions and are abouthalf as large as the natural ones. HNO3 is responsible for the forming of NAT, of which Type IPSCs consist.

2.3. THE IMPORTANCE OF AEROSOLS IN THE STRATOSPHERE 9

2.3 The Importance of Aerosols in the Stratosphere

Both tropospheric and stratospheric aerosols play an important role in global climate change.Especially the sulfate aerosols globally affect the climate due to their even distribution. Thoughthe absorption at wavelengths below 2.5 to 3 µm is negligible, the high scattering cross sectionincreases the planetary albedo significantly. The absorption and emission of terrestrial radianceis an important factor as well.

The presence of aerosols in the stratosphere allows a large number of heterogenous chemicalreactions [section 2.3.2]. At temperatures below 210 K these processes intensify the chlorine andbromine activation, which leads to a large acceleration of the arctic ozone loss during late polarwinter [Wagner 1999; Fitzenberger 2000b].

2.3.1 The Radiation Budget

Natural variations of aerosols, especially those due to eruptions of large volcanoes, and the an-thropogenic emission of aerosol precursors have a significant impact on the planetary radiationbalance, and thus can cause global temperature changes. During the past century aerosols havebeen one of the greatest sources of uncertainty in the interpretation and the projection of cli-mate change [Hansen et al. 1998]. Additionally the knowledge of the atmospheric content andthe composition of aerosols is essential to comprehend the heterogeneous chemistry in the at-mosphere and to be able to predict the influence of natural and anthropogenic variations in theaerosol budget.

Investigations of Hofmann [1990] with balloon-borne aerosol counters close to Laramie,Wyoming (USA) seem to indicate an increase of the aerosol mass density within the Jungelayer of some percent per year. Remarkable is a shift of the size distribution to larger radii,whereas the particle number only slightly increases. Possible reasons can be the combustion oforganic matter or the increasing global temperature causing higher biological activity.

Volcanic eruptions alter much more the aerosol content in the stratosphere, injecting mineralparticles and sulfureous gases. But as they are singular events their influence generally is tran-sient [see section 2.5] and particular for each eruption. The effect thus sensitively depends on thequantity of exhausted gas for which the magnitude of the eruption is not necessarily relevant.For example the spectacular blast of Mount St. Helens in early 1980 almost did not enhancethe stratospheric aerosol content since injected particles were small and the emitted amount ofgas was very low (Figure 2.5). However, the primarily neglected eruption of El Chichon injectedlarge quantities of aerosol forming matter into the stratosphere because the emission of sulfure-ous gases was very high. Furthermore, the cooling (absorption of solar radiation, increase ofthe planetary albedo) and the warming (absorption and emission of terrestrial radiance) effectsof mineral aerosols compensate each other whereas sulfureous aerosols lead to a heating of thestratosphere and a cooling of the surface [Pollack et al. 1976].

Hansen et al. [1978] and Newell and Weare [1976] directly measured the increase of thestratospheric temperature and the decline of the temperature in the lower troposphere afterthe eruption of Gunung Agung in march 1963. Until the late summer of 1964 the the strato-spheric temperature increased up to a maximum of 5 K and the surface temperature decreasedup to 0.4 K. Surprisingly this effect failed to appear after the El Chichon eruption. A few monthslater the surface temperature even increased. The unusually strong El Nino event of 1982/83probably covered completely the consequences of the eruption [Roedel 1992]. The eruption of


Pinatubo raised about 20 Megatons of SO2 to the stratosphere, almost three times as much asEl Chichon. During the first fourteen months after the eruption the global temperature at thesurface decreased by about 0.5 K [Dutton and Christy 1992].

Detailed investigations of volcanic emissions and their consequences on the stratosphericaerosol content, the aerosol size distribution and composition and the successional climatechanges may open possibilities to calibrate climate models and to properly anticipate the ef-fects of the anthropogenic sulfur emissions.

2.3.2 Heterogeneous Chemistry

Chlorine and Bromine are the most important substances which catalyse the ozone depletionin polar regions at wintertime. At ground level chlorine is produced by many natural processes(e.g. sea salt and volcanic emissions). Due to its high solubility, precipitation rapidly removesHCl from the atmosphere. Chlorofluorocarbons of anthropogenic origin are very stable in thetroposphere and therefore easily reach the stratosphere where they are photodissociated. Themain precursor of stratospheric bromine is CH3Br which emerges from biological processes inthe oceans and is used as disinfectant and in plumbiferous gasoline.

To describe the rapid ozone depletion like it has been observed in the Antarctica since 1985every year during polar spring [Farman et al. 1985] one has to consider the following heteroge-neous chemical processes:

Polar Stratospheric Clouds

The presence of PSCs enables the hydrolysis of N2O5 [Solomon et al. 1986] which was confirmedin numerous laboratory measurements [e.g. Leu 1988; Tolbert et al. 1988]:

N2O5 + H2Ohet−→ 2HNO3 (2.13)

N2O5 is relatively stable in the stratosphere and thus acts as a reservoir of the inorganic nitrogenfamily (NOy). HNO3 is even more stable, and freezes out as NAT at temperatures below 195K forming PSC particles [section 2.1]. These particles sink to lower atmospheric layers afterreaching a definite size, and consequently reduce the stratospheric content of nitrogen compounds[denitrification, Toon et al. 1986]. Solomon et al. [1986] pointed out that the thereby suppressedNO2 concentration impedes the rapid reformation of the ClONO2 and BrONO2 reservoirs:

ClO + NO2M−→ ClONO2 (2.14)

BrO + NO2M−→ BrNO2 (2.15)

Thus, the concentration of the ClO and BrO radicals in the polar stratosphere is increasedsignificantly.

The most important reactions for the activation of long-lived compounds like ClONO2 andHCl on PSC particles were suggested by Solomon et al. [1986] and McElroy et al. [1986] andverified in the laboratory by Molina et al. [1987] and Tolbert et al. [1987]:

ClONO2 + HClhet−→ Cl2 + HNO3 (2.16)

ClONO2 + H2Ohet−→ HOCl + HNO3 (2.17)

2.3. THE IMPORTANCE OF AEROSOLS IN THE STRATOSPHERE 11

Further important heterogeneous reactions occurring on the surface of PSCs are1:

HOCl + HClhet−→ Cl2 + H2O (2.18)

BrONO2 + HClhet−→ BrCl + HNO3 (2.19)

HOBr + HClhet−→ BrCl + H2O (2.20)

N2O5 + HClhet−→ ClONO + HNO3 (2.21)

At the end of the polar night when the stratosphere is exposed to solar radiation again, theweakly bound halogen species (e. g. Cl2 and HOCl) are rapidly photolyzed:

Cl2hν−→ 2Cl (2.22)

HOClhν−→ Cl + OH (2.23)

The released atomic chlorine rapidly depletes ozone. For high concentrations of ClO the catalyticClO-dimer-cycle is very effective [Molina and Molina 1987]:

2(Cl + O3 −→ ClO + O2) (2.24)

ClO + ClOM−→ Cl2O2 (2.25)

Cl2O2hν−→ Cl + ClO2 (2.26)

ClO2M−→ Cl + O2 (2.27)

Net: 2O3 −→ 3O2 (2.28)

The rate limiting step of this cycle is the self reaction of ClO, making the ozone depletionpotential (ODP2) proportional to the square of the ClO concentration. This process is theprimary catalytic process responsible for about 75% of the ozone removal in the ozone hole.McElroy et al. [1986] pointed out the role of the bromine chemistry in the ozone depletion (inparticular the reaction between ClO and BrO). This cycle contributes about 20% to the forma-tion of the Antarctic ozone hole [Anderson et al. 1989].

Sulfate Aerosols

The discovery of the Antarctic ozone hole raised fears that also at lower latitudes great ozonelosses could occur. Only few years later a statistically significant ozone loss could be ob-served at mid-latitudes in both hemispheres, not only in winter but also in other seasons [e.g. WMO/UNEP 1994; Hollandsworth et al. 1995; Staehelin et al. 1998]. In the beginning of the1990s the ozone deficit was of 5-10%, substantially larger than calculated with homogeneous gaschemistry models [WMO 1999; Bojkov and Fioletov 1995; Sze et al. 1989].

Fahey et al. [1993] explained these phenomena with a heterogeneous chemistry on liq-uid stratospheric sulfuric acid aerosols. The ozone losses were dramatically increased athigh aerosol concentrations after volcanic eruptions [Avallone et al. 1993; Hofmann et al. 1994;

1The heterogeneous reactions of HOBr, HOCl, ClONO2, BrONO2 and N2O5 with HBr are less importantbecause of the low concentration of HBr

2An index called the Ozone Depletion Potential (ODP) has been adopted for regulatory purposes under theMontreal Protocol. The ODP of a compound is an estimate of the total ozone depletion due to 1 kg of thecompound divided by the total ozone depletion due to 1 kg of Freon 11 (CFCl3).


Figure 2.2: Reaction coefficients for the most important heterogeneous reactions on sulfuric acidaerosol surfaces (for a H2O partial pressure of 2.5 · 10−4 mbar and a HCl mixing ratio of 500ppt). Adopted from Erle [1999]

Solomon et al. 1996; Portmann et al. 1996]. After the eruption of Mount Pinatubo in June 1991,it was possible to study the effect of enhanced aerosol loading of the PSC-free stratosphere withregard to chlorine activation on sulfuric acid aerosols.

The N2O5 hydrolysis was the first reaction found to be relevant not only on PSCsbut also on sulfuric acid aerosol surfaces [Mozurkewich and Calvert 1988; Tolbert et al. 1988].This reaction leads to a reduction of NOx and is therefore indirectly responsible for an in-crease in reactive chlorine. The BrONO2 [Erle et al. 1998] hydrolysis has a similar effect,since it changes BrONO2 to HOBr, which photolyzes faster. The significance of theses re-actions both for polar regions and for mid-latitudes, after the eruptions of El Chichon andPinatubo was evidenced by model studies [Rodriguez et al. 1991; Brasseur and Granier 1992;Lary et al. 1996; Tie and Brasseur 1996] and by measurements of conspicuously reduced NO2densities [Johnston et al. 1992; Solomon et al. 1994; Slusser et al. 1997].

In addition, the above mentioned heterogeneous reactions on PSC particles play an im-portant role on sulfate aerosols in both polar regions [Portmann et al. 1996] and mid-latitudes

2.4. ELEMENTARY STRATOSPHERIC DYNAMICS 13

[Solomon et al. 1996]. Figure 2.2 summarizes the reaction probabilities for the different reactions.

The loss of ClONO2, BrONO2, HOCl and HOBr on HCl can be neglected for temperaturesabove 200 K and background aerosol conditions, but become important for temperatures between200 and 210 K when aerosol levels are slightly increased. The bromine reaction coefficients arehigher by 1-3 orders of magnitude than the respective chlorine reactivities. Thus, the effect ofbromine is comparable to the one of chlorine even though bromine is 100 times less abundantthan chlorine. The reaction coefficients of the heterogeneous bromine reactions measured in thelaboratory [Hanson and Ravishankara 1995; Hanson et al. 1996; Waschevsky and Abbatt 1999]argue for higher atmospheric relevance than assumed previously.

2.4 Elementary Stratospheric Dynamics

The distribution of the chemical species in the atmosphere is determined by chemical and dy-namical processes. Vice versa the distribution of some photochemical species, especially ozone,affects the thermal structure of the atmosphere and thus the dynamical processes. Transport ofchemical species can be realized by the prevailing winds (advection) and by turbulent mixing(diffusion).

The absorption of ultraviolet radiation, particularly by ozone and less by molecular oxygen,is the main heating process in the middle atmosphere. The observed increase in temperaturewith altitude in the stratosphere is attributed to the ozone UV absorption (see Figure 2.3).Radiative cooling occurs through infrared emissions by CO2, H2O and O3.

Brewer [1949] suggested that the H2O concentration in the stratosphere (around 1 ppm) wasgoverned primarily by the rising air from the troposphere. Therefore, the content is determinedby the lowest temperature the ascending air parcel experiences in the tropopause. He notedthat the middle and high latitude tropopause was too warm to explain the observed dryness ofthe stratosphere and that only in the tropics, where the tropopause is highest (17-18 km), thetemperature is sufficiently low. Thus he suggested that air parcels rose from the troposphere tothe stratosphere only in the tropics and descended at extra-tropical latitudes, forming the socalled Brewer-Dobson circulation or Hadley cell. Dobson [1968] noted that the maximum ozoneproduction was found in the upper tropical stratosphere but that the largest concentrations wereobserved at high latitudes, requiring a downward-poleward transport.

Figure 2.4 shows the mean meridional circulation. The summer stratosphere is near to ra-diative equilibrium with generally higher temperatures than the equilibrium temperatures inwintertime. The reasons are stationary planetary waves, which originate from the tropospherewith typical wavelengths close to the Earth’s diameter (corresponding to wave numbers of 1 to3). They only propagate into the stratosphere, when there are westerly winds, occurring in thestratosphere only at wintertime. During winter the polar stratosphere is cooling off strongly.Colder air is descending like in a tropospheric low pressure system. Analogously a vortex isforming around the polar region, which causes the stratospheric westerly winds. Their velocityis proportional to the pressure gradient and reaches typically 50 m

s . These high velocities almostcompletely impede the exchange of airmasses between the polar region and the exterior of thevortex [Waugh et al. 1994].

Since the distribution of water and land surfaces is very different for the two hemispheres,planetary waves are much more frequent and intensive in the northern hemisphere. Thus, thetemperatures in the northern polar vortex are not as low as in the Antarctic stratosphere.


Figure 2.3: Vertical structure of the atmosphere [Brasseur and Solomon 1986]

Planetary waves may lead to sudden minor warmings of the vortex or even to a total breakdownalready at wintertime [major warming, Pawson and Naujokat 1999]. Normally the polar vortexbreaks down in spring, when the stratosphere is again sunlit. This so called final warming yieldsan extreme increase in planetary wave activity. The vortex is split into different parts whichare transported to mid-latitudes. The winter circulation changes to summer circulation witheasterly winds. In the northern hemisphere this breakdown takes place in March/April, in thesouthern hemisphere between October and December.

2.5. VOLCANIC AEROSOL 15

Figure 2.4: Meridional circulation deduced by the theoretical study of Cunnold et al. [1975]

2.5 Volcanic Aerosol

All the disposable data recorded during the last decades clearly demonstrate the strong perturba-tions of stratospheric sulfate aerosols caused by volcanic eruptions with large sulfureous gas emis-sions [Hofmann 1990; Jager 1991; Chazette et al. 1995; Osborn et al. 1995; Uchino et al. 1995;Thomason et al. 1997; Stothers 2001]. Historical records over the last 100 years or more indicatethe most recent 30-year period as a relatively active volcanic period. Primarily through the useof historical pyrheliometric data, Stothers [1996] showed that eight major eruptions happenedduring the past century. Four of these occurred between 1880 and 1910 (Krakatau, an uniden-tified eruption, Santa Maria, and Katmai) and four occurred since 1960 (Agung, Fernandina,El Chichon, and Pinatubo). Between 1910 and 1960, the stratosphere was almost undisturbedby volcanic activity. Junge et al. [1961] performed his measurements of the stratospheric sulfateaerosol layer at the very end of this period, thus he examined pure stratospheric backgroundaerosol. Krakatau and Pinatubo were the two largest eruptions in the past century relatingto their impact on the stratosphere [Stothers 1996]. The analysis of ice cores from Greenland,


Figure 2.5: Stratospheric aerosol loading observed from 1976 to 1999. The aerosol loading is taken asthe integrated column backscatter above the tropopause measured by the LIDAR at Garmisch-Partenkirchen, Germany (47N, 11E). Important volcanic eruptions are indicated by arrows.(Adopted from Pyle et al. 1999)

which cover an even longer period, indicate that global volcanism has an approximately 80-yearperiodicity [Hammer et al. 1980].

The two most recent major eruptions of El Chichon (17N, April 1982) and Mount Pinatubo(15N, June 1991), both in the tropics, occurred at different phases of the quasi-biennial os-cillation (QBO), so that the dispersal and decay of the aerosol from the eruptions were dif-ferent. After low-latitude eruptions aerosol accumulates in a tropical stratospheric reservoir[Trepte and Hitchman 1992]. This reservoir has an abrupt, narrow boundary on the winter hemi-sphere side and a broad boundary on the summer hemisphere side, as noted after the Pinatuboeruption [Grant et al. 1996; Lambert et al. 1997]. QBO easterly winds over the equator impedethe propagation of planetary waves into the tropics. Thus, the aerosol reservoir remains relativelyisolated from the mid-latitudes. QBO westerly winds allow planetary waves to penetrate deeperinto the tropics and to transport the aerosols poleward more easily. After both El Chichon andPinatubo eruptions QBO easterly winds were prevalent. In autumn 1982 the equatorial strato-sphere made the transition to the westerly phase of the QBO. Hence a large increase in aerosolloading was observed in the northern hemisphere [Pollack et al. 1983]. In the case of El Chichonthe QBO easterly winds where very strong, so that most of the aerosols remained in the northernhemisphere [Trepte et al. 1993]. In contrast, Mount Pinatubo erupted at a weaker phase of east-erly QBO. The absence of strong winds may have allowed significant amounts to be transportedacross the equator. After the reversal of the QBO aerosols were rapidly transported to higherlatitudes of both hemispheres [Rosen et al. 1994; Godin et al. 1996; Deshler et al. 1997].

Measurements of peak backscatter, peak mass, column backscatter, and column massof stratospheric aerosol during the first three years following Mount Pinatubo eruptionshowed an exponential (e−1) decay time of 1 ± 0.2 years [Rosen et al. 1994; Jager et al. 1996;Barnes and Hofmann 1997; Deshler et al. 1997]. The peak parameters primarily reflect sedimen-tation, and are not influenced by fluctuations at tropopause height. The time required until theaerosol surface area reached pre-Pinatubo values is shorter at higher altitudes. At 25 km it tooksome 3.5 years to reach pre-Pinatubo conditions, whereas at 15 km they were not reached until

2.6. BACKGROUND AEROSOL 17

mid- to late 1996, more than 5 years after the eruption. The decay of aerosol surface area is about20-30% slower than the decay of backscatter or mass [Rosen et al. 1994; Jonsson et al. 1996;Deshler et al. 1997]. The decay of integrated backscatter was around 25% longer after MountPinatubo eruption than after El Chichon [Chazette et al. 1995], due to the higher lofting of par-ticles following the eruption of Pinatubo. Russel et al. [1996] showed that the effective radius[chapter 3.4] increased from 0.15 µm prior to Pinatubo to 0.55 µm one year later, decreasing to0.45 µm in spring 1993. In fall 1994 the effective radius reached a value below 0.2 µm and hasremained between 0.15 µm and 0.2 µm since then [Deshler et al. 1997].

These decay rates of volcanic stratospheric aerosol are similar to the observed rates for theremoval of strontium from the stratosphere [Fabian et al. 1968]. The relatively constant decayrate of stratospheric sulfate aerosols suggest that the removal is controlled by relatively steadyand robust processes, such as gravitational settling and stratospheric-tropospheric exchange. LI-DAR measurements provide sufficient temporal and vertical resolution to observe these exchangeprocesses [Menzies and Tratt 1995; Sassen et al. 1995; Post et al. 1996].

2.6 Background Aerosol

The measurements of Junge et al. [1961] were made at the end of a long period free of volcaniceruptions but were not extensive enough to draw conclusions on the source of the sulfuric acidaerosols in the stratosphere and the stability of the Junge layer. Continuous measurements ofaerosol properties and content were not performed until the 1970s [chapter 3.1]. There have beenfour periods since then when the volcanic influence has been at a minimum: 1974, 1979, 1989 toearly 1991, and the present.

In situ measurements at 41N presented by Hofmann [1990] showed an increase in strato-spheric sulfate aerosols of 30 to 50% between 1979 and 1989. SAGE data from this latitudeconfirmed this result [Thomason et al. 1997]. Hofmann [1990] suggested that increased surfaceemissions of OCS or SO2 might be responsible, but more recent studies [Chin and David 1995]predicated the amount of stratospheric sulfur derived from the oxidation of OCS as too small toexplain the observed minimum aerosol loading. Another reason may be the sulfur from aircraftexhaust [Hofmann 1991], but model studies by Bekki and Pyle [1992] suggest that the rise inair traffic from 1979 to 1989 was too small to account for the measured increase. This led to thequestion whether any pure background aerosol conditions had been reached during the abovementioned periods.

Column stratospheric backscatter measurements at Mauna Loa in late 1996 showed the lowestvalue observed during the previous 17 years, but as the QBO changed phase in 1997 it increasedby a factor of 2 [Barnes and Hofmann 1997]. This modulation can be explained if one regardsthe upper tropical troposphere as a source of condensation nuclei [Brock et al. 1995] and SO2[Weisenstein et al. 1997]. The consequence is a strong seasonal cycle: An increase and relativestability during winter and spring followed by a decrease in summer and fall, depending on thetropopause height. At very low levels of stratospheric sulfate aerosols the column backscattercan vary by a factor of 2 to 4 between different seasons. Several more years of data unaffectedby volcanic activity are required to determine if a stable background has been reached and tocompare with earlier minimum periods.

Chapter 3

Optical Properties of StratosphericAerosols

3.1 Measurement Methods

As a result of the various measuring methods characteristics, spatial distribution, formation andevolution of stratospheric aerosols have become much better known. Direct in-situ measurementsof aerosols using a particle counter or collector have been performed as well as optical methodsobserving the extinction due to scattering processes and absorption or the properties of thescattered light itself. Depending on the purpose of the measurements a variety of instrumentshave been used on various platforms such as ground, balloons, aircrafts, and satellites.

3.1.1 In-situ Measurements

Most commonly used for in-situ measurements of aerosols in the stratosphere are balloon-borneinstruments. Aircrafts are predominantly employed in the troposphere. Only those reaching veryhigh operating altitudes can also be applied at heights around the tropopause. For example,Keil et al. [2001] performed in-situ measurements of tropospheric aerosols by using a particleimpactor to measure chemical and physical aerosol characteristics on a small two-engine aircraftequipped with sensors for meteorological, avionic, and radiation parameters. Stratospheric bal-loons easily reach altitudes over 40 km, and provide a high vertical resolution of the observedparameters over a large altitude range.

In-situ measurements allow precise determination of particle number, size distribution andmass of the aerosols. Impactors and other particle collectors enable a direct analysis of the chemi-cal composition after recovering the gondola. Various in-situ measurements using a balloon-borneparticle counter were performed by Hofmann et al. [1975] and Ovarlez and Ovarlez [1995], espe-cially after the large volcanic eruptions of El Chichon [Hofmann and Rosen 1983] and Pinatubo[Deshler et al. 1993]. They also monitor refractive indices, extinction and scattering propertiesby drawing the surrounding air into the focus of a light beam, thus combining in-situ and straylight measurements. The great disadvantages of the in-situ measurements is the impossibility ofglobal monitoring and that it is very laborious to perform continuous observations. Furthermore,particle counters are not able to detect aerosols of a radius smaller than about 0.1 µm.

19

20 CHAPTER 3. OPTICAL PROPERTIES OF STRATOSPHERIC AEROSOLS

3.1.2 Stray Light Measurements

Measuring the intensity, the angular distribution and the polarization of light scattered byaerosols gives a lot of information about the physical and the chemical properties of the particles[see section 3.3]. In addition to sun and moonlight, scattered light from artificial sources, suchas a LASER, is observed. Being remote sensing techniques, such observations can be performedfrom all the different kinds of measurement platforms.

Remote sensing of stratospheric aerosols from the Earth’s surface is only possible bythe application of a LIDAR system [Barnes and Hofmann 1997; Abo and Nagasawa 1994;Iwasaka et al. 1994; Yasui et al. 1994]. Short LASER pulses are transmitted vertically into theatmosphere. At each altitude some of the light is backscattered into the detector. Light backscat-tered from high altitudes needs more time to reach the detector than light backscattered fromlow altitudes. By measuring the temporal dependence of the detected light one gets a verticalbackscatter profile. A global LIDAR networks covers a great part of the northern hemisphericcontinents. Because of the large distance from the ground to the stratosphere, ground-basedLIDARs are better suited for tropospheric than for stratospheric observations. To measurebackscattering of stratospheric aerosols a clear sky is necessary and a correction of the tro-pospheric influence has to be implemented.

To avoid tropospheric influence, stray light measurements have to be performed directly inthe stratosphere. RADIBAL [Herman et al. 1986] is a balloon-borne instrument with a narrowfield of view in horizontal direction which measures the radiance and the polarization ratio ofthe scattered sunlight. It operates within the aerosol layer, so that the observed scattered lightmainly accrues from aerosols and molecules within a few tens of kilometers from the gondola.The vertical resolution is about 2 km. Another example for a balloon-borne instrument whichobserves scattered sunlight is BALLAD [Ramon 1995].

Since a global network of balloon flights and ground-based measurements is too expensiveand even impossible over the oceans, satellites bear instruments which measure solar radiationscattered back from the atmosphere [Kaufmann 1995], as well as LIDAR systems. The French in-strument Polarization and Directionality of the Earth’s Reflectance (POLDER), launched aboardthe ADEOS spececraft, measures the polarization, and directional and spectral characteristics ofthe solar light reflected by aerosols, clouds, oceans and land surfaces [Herman et al. 1997]. TheAdvanced Very High Resolution Radiometer (AVHRR) instrument flies on the NOAA series ofpolar-orbiting Sun-synchronous satellites. 2048 channels span a viewing angle of ±55 degreesfrom the nadir [Rao et al. 1989; Stowe et al. 1992], in which upwelling radiance is measured atfive narrow-band wavelength channels in the visible and infrared.

Even the Space Shuttle served as a measurement platform. In 1989, a Fourier transformspectrometer [Farmer 1989] performing limb observations and in September 1994, a three-wavelength backscatter LIDAR [McCormick et al. 1995] flew on the Space Shuttle for severaldays. The goal was the validation of space-borne LIDAR technologies for future satellite instru-ments. PICASSO-CENA (the Pathfinder Instruments for Clouds and Aerosols using SpaceborneObservations-Climatologie Etendue des Nuages et des Aerosols), to be launched in 2003, uses adual-wavelength polarization-sensitive LIDAR. The gathered data presented the first highly de-tailed global view of the nadir-viewed vertical structure of clouds and aerosols from the Earth’ssurface up to the middle stratosphere.

3.1. MEASUREMENT METHODS 21

3.1.3 Extinction Measurements

A simple way to obtain aerosol characteristics is the measurement of direct sunlight extinction.Due to the high intensity of the solar light the total extinction can be measured very exactlyand the portion of light in the field of view originating from multiple scattering is negligible.In the infrared the aerosol extinction can directly be measured at certain wavelength ranges,while throughout the visible and the UV spectral range other absorbing and scattering effectshave to be corrected [chapter 5.2 and 3.2]. To derive other physical and chemical characteristics,such as size distribution or refractive index, the aerosol extinction has to be measured at variouswavelengths [Ahlquist and Charlson 1969; Thomalla and Quenzel 1982]. By implementing Miecalculations, diverse assumptions of these parameters can be compared with the measurementresults [see sections 3.3 and 3.5]. Aerosol extinction measurements are the most commonly usedmethod to obtain aerosol information by remote sensing from airborne instruments.

Satellite-borne Instruments

The only way to obtain a global view of the aerosol distribution and characteristics is theobservation by a satellite borne instrument. The Stratospheric Aerosol Measurement (SAM)II and the Satellite Aerosol and Gas Experiment (SAGE) I and II have provided the mostcomplete global aerosol extinction data set. Ground-based LIDAR networks, although mainlyin the northern mid- to high latitudes, the Polar Ozone and Aerosol Measurement II (POAM)[Bevilacqua 1997], and the Halogen Occultation Experiment (HALOE) [Russell III et al. 1993]have also contributed a global picture of the aerosol content.

The SAGE II instrument aboard the Earth Radiation Budget Satellite (ERBS) in a circularorbit with a 57-degree inclination has been measuring vertical profiles of aerosol extinction atfour wavelength channels during spacecraft sunrise and sunset since October 1984. Additionallythe gaseous absorbers nitrogen dioxide, ozone, and water are measured [Mauldin et al. 1985].Its observations cover approximately 75S to 75N over a year, providing near-global coverage.The SAGE II instrument takes 15 sunset and 15 sunrise measurements each day with a verticalresolution of 1 km. The latitudinal spacing between two subsequent measurements is roughly0.5 degrees depending on the latitude, while the longitudinal spacing is around 24 degrees. Moreinformation on the SAGE II program can be found in McCormick [1987]. The data retrievalalgorithm is described in Chu et al. [1989].

POAM II was launched aboard the French satellite System Pour l’Observation de la Terre(SPOT) 3 on September 25, 1993, into a Sun-synchronous polar orbit with a 98.7 degree in-clination. The POAM II instrument is another solar occultation instrument measuring verticalprofiles of aerosols, ozone, nitrogen dioxide, and water in nine wavelength channels between0.35 and 1.06 µm, with 1 km vertical resolution [Glaccum et al. 1996; Randall et al. 1996]. Ittakes 14 sunrise and 14 sunset measurements each day, whereas the sunrise events occur entirelyat high northern latitudes (55 to 71N) and the sunset events at high southern latitudes (63to 88S). The measurements also allow the detection of PSCs, and Polar Mesospheric Clouds(PMCs) [Fromm et al. 1997; Debrestian et al. 1997]. In Lumpe et al. [1997] the retrieval algo-rithm is described. The instrument operated successfully until November 14, 1996, when theSPOT 3 satellite went out of service [Bevilacqua 1997]. An improved version of POAM II, thePOAM III instrument [Lucke et al. 1999], was launched in March 1998 aboard the SPOT 4satellite and since then has continuously provided aerosol extinction profiles at approximatelythe same wavelengths as POAM II [Randall et al. 2001].


The Halogen Occultation Experiment (HALOE) is a satellite-borne instrument also usingthe solar occultation technique to measure vertical profiles of O3, NO2, NO, H2O, HCl, HF,CH4, and aerosol extinction, but in the infrared spectral range between 2.45 and 10.04 µm[Russell III et al. 1993]. It was launched on September 12, 1991, aboard the Upper AtmosphereResearch Satellite (UARS) into a circular orbit similar to the one of SAGE II. When thereare high aerosol loadings, as occurring after volcanic eruptions, it can provide stratosphericmicrophysical aerosol information [Hervig et al. 1998].

Another solar occultation instrument, which measures vertical profiles of O3, NO2, N2O,N2O5, HNO3, H2O, CH4, CFC-11, CFC-12, and aerosol extinction in the infrared band between6.21 and 11.77 µm and at 0.78 µm is the Improved Limb Atmospheric Spectrometer (ILAS)[Sasano 1996; Sasano et al. 1999]. It was launched on August 17, 1996, aboard the JapaneseAdvanced Earth Observation Satellite (ADEOS) into a polar orbit similar to the one of POAM.The launch of a revised version, the ILAS II instrument, aboard ADEOS II is scheduled for2002.

One of the most important instruments providing stratospheric aerosol data in the nearfuture is SAGE III. The Meteor-3M spacecraft carrying the SAGE III instrument was launchedon December 10, 2001, from the Baikonur Cosmodrome in Kazakhstan on a Ukrainian builtZenit-2 Rocket. Whereas previous instruments in the SAGE series used single silicon diodes,the SAGE III instrument uses an 800-pixel charge-coupled device (CCD) linear array detector.The CCD is designed to measure aerosol extinction coefficients centered at wavelengths of 0.385,0.450, 0.521, 0.676, 0.756, 0.869, and 1.0195 µm, and absorption features of O3, NO2, and H2O[McCormick et al. 1999]. To improve the size discrimination of larger aerosol particles and toseparate cloud and aerosol signals, the instrument also has a channel centered at 1.54 µm.With SAGE III, lunar occultation capability has been added, which increases the measurementopportunities with moonrises and moonsets, and allows to measure atmospheric species that arenot observable in the presence of daylight.

An extensive European project is the Scanning and Imaging Absorption Spectrometerfor Atmospheric Cartography (SCIAMACHY) [Burrows et al. 1995; Bovensmann et al. 1999;Noel et al. 2001]. It is one of the payload instruments embarked aboard EnviSat-1 (EnvironmentSatellite) of the European Space Agency (ESA). It will provide the most complete observationset ever of the Earth’s meteorological and climatic data. The launch by an Ariane 5 rocketfrom Europe’s launch site in Kourou, French Guiana, is scheduled for March 1st, 2002. Directsunlight extinction measurements as well as limb and nadir observations will be performed overthe wavelength ranges 0.24 to 1.70 , 1.94 to 2.04, and 2.27 to 2.38 µm. The primary objectiveis to generate global maps and vertical profiles of the aerosol’s optical density, and of tropo-spheric and stratospheric trace gases, in particular NO2 and ozone, to identify the aerosol type,to estimate cloud coverage and cloud top height, and to measure the surface reflectance.

Balloon-borne instruments

Only a few aerosol extinction measurements have been performed by balloon-borne instruments,so far exclusively in connection with trace gas absorption measurements. Balloon-borne mea-surements of direct sunlight are particularly suitable for the observation of key species of thestratospheric chemistry [chapter 4]. Since the latter one is strongly affected by aerosols [see 2.3.2],it is very important to observe aerosol characteristics under the same measurement conditions.

3.2. RAYLEIGH SCATTERING 23

The DOAS (Differential Optical Absorption Spectroscopy) instrument designed at the IUP(Institut fur Umweltphysik) of the University of Heidelberg, Germany, measures vertical profilesof O3, NO2, NO3, BrO, OClO, O4, H2O, IO, OIO, and aerosol extinction. The measurementtechnique is described in detail in chapter 4. During either ascent or descent of the balloon andduring the solar occultation, solar extinction is measured in the near ultraviolet (0.320 - 0.422µm) and visible (0.417 - 0.670 µm) spectral range at 1024 continuously distributed channels,respectively. Nine flights have been performed successfully since November 23, 1996. The re-trieval algorithm of the aerosol characteristics from the measured extinction data is describedin chapter 5.

The AMON (Absorption par Minoritaires Ozone et NOx) uses a similar measurementgeometry and additionally implements direct starlight measurements [Renard et al. 1996] inthe UV-visible to obtain vertical profiles of O3, NO2, NO3, OClO and possibly OBrO[Renard et al. 1998]. The aerosol extinction coefficient is obtained after subtracting the con-tribution of all the species mentioned above [Pirre et al. 2000]. The work is still in progress, buthas already allowed to detect the presence of a Polar Stratospheric Cloud at high latitudes onFebruary 1997 [Riviere et al. 2000]. The SALOMON (Spectroscopie d’Absorption Lunaire pourl’Observation des Minoritaires Ozone et NOx) instrument performs similar measurements usingthe Moon as a light source [Renard et al. 2000].

3.2 Rayleigh Scattering

Rayleigh scattering is defined as scattering by particles, whose dimensions are small comparedto the wavelength of the incident radiation. The radiation at the position of the particle canbe treated as an oscillating homogeneous electric field. The charges of the particle tend toredistribute themselves in the direction of the field, thus forming an electric dipole, oscillatingwith the same frequency as the incoming light. Rayleigh assumed, that scattering by smallparticles can be described as the irradiation of this dipole. For the derivation of the Rayleighscattering cross section in the atmosphere only air molecules are considered. Possibly occurringsmall aerosols are excepted.

The separation of positive and negative electric charges of a particle in an electric field iscalled polarization, but should not be confused with the polarization of light. The dipole moment~p of a polarizable particle is proportional to the surrounding electric field ~E,

~p = α~E, (3.1)

where α is called the polarizability. For spherical symmetric molecules which can be polarizedin each direction, Chandrasekhar [1960] found the radiation flux φ into a steradian dΩ′ to be

φ =128π5

3λ4α2I dΩ

3

4

(

1 + cos2 θ) dΩ′

4π. (3.2)

I is the intensity of the incoming light from the steradian dΩ, of wavelength λ. The ratio ofthe intensities of scattered light polarized parallel (I‖) and perpendicular (I⊥) to the scatteringplane is calculated to be

I‖I⊥

= cos θ, (3.3)

where θ is the scattering angle. That means, that light scattered perpendicular to the incominglight is completely polarized. In reality, at scattering angels of 90 no complete polarization is


observed. Most of the atmospheric molecules have a linear structure, and thus the polarizabilityis not isotropic. This effect can be described by the anisotropy γ and the depolarization factorρ, which are correlated by

ρ =2γ

γ + 1. (3.4)

The depolarization factor is defined by

ρ =I‖I⊥

∣

∣

∣

∣

∣

θ=90

. (3.5)

Therefore, the phase function of Rayleigh scattering can be written as [Chandrasekhar 1960;Penndorf 1957]

p(θ) =3

4 (1 + 2γ)

(

(1 + 3γ) + (1− γ) cos2 θ)

. (3.6)

Penndorf [1957] specifies the depolarization factor of air to be 0.035, leading to the Rayleighscattering phase function

p (θ) = 0.7629(

1 + 0.932 cos2 θ)

. (3.7)

Assuming an analytical expression for the wavelength dependence of the air refractive index,Brasseur and Solomon [1986] quoted the following approximation of the total Rayleigh scatter-ing cross section per molecule,

σRayleigh (λ) =4 · 10−28

λ3.916+0.074λ+0.05λ

cm2 (3.8)

where λ has to be inserted in the unit of microns. The total extinction coefficient of Rayleighscattering in air of molecular density nAir is given by

βRayleigh = σRayleigh · nAir. (3.9)

3.3 The Mie Theory

The Mie Theory formally describes the scattering of light by spherical particles of arbitrarydimension and was first treated by Mie [1908]. The intensity and the polarization factor of thescattered light can be computed depending on its wavelength, the scattering angle, the radiusof the droplet, and the complex refractive index of the aerosol forming liquid. Integration overan aerosol size distribution [see 3.4] and the scattering angle allows to calculate the extinctiondue to scattering. Since a complex refractive index is considered, the Mie theory completelydescribes the extinction of aerosols with a given size distribution and chemical composition.

A useful parameter to characterize the size of a droplet is the size parameter x:

x = 2πr

λ(3.10)

For small particles with a radius r much smaller than the wavelength λ of the scattered light, theMie theory approximates to the Rayleigh scattering. The Rayleigh scattering can be treated asthe irradiation of a single oscillating dipole, whereas the Mie theory satisfies Maxwell’s equationsfor the electromagnetic field about a dielectric sphere. Mie calculations are necessary if the sizeparameter is around unity or larger.

3.3. THE MIE THEORY 25

The derivation of the solution is a straightforward application of classical electrodynam-ics. The dielectric sphere of radius r is expanded in spherical harmonics and Bessel functions[Liou 1980]. The incoming light is supposed to be linearly polarized and outside the sphere vac-uum prevails (refractive index 1). The results can be given in form of infinite series, but witha slow convergence for large values of x. Modern computers easily enable calculations for bothreal and complex refractive indices even for large size parameters, not possible a few decadesago.

The intensity I of the scattered light depending on the scattering angle θ is calculated to be[Van De Hulst 1957]

I (θ) =I0 (i1 (θ) + i2 (θ))

2k2d2, (3.11)

where I0 is the intensity of the incoming unpolarized light, k = 2πλ

is the wave number, d is thedistance from the center of the particle, and i1 and i2 are the squares of the absolute value of theamplitude functions S1 and S2. S1 relates to the light polarized parallel to the scattering plane,S2 to the light polarized perpendicular to the scattering plane. The two amplitude functionshave the symmetrical form [Goody and Yung 1989]

S1 (θ) =∞∑

n=1

2n+ 1

n (n+ 1)(anπn (cos θ) + bnτn (cos θ)) (3.12)

S2 (θ) =∞∑

n=1

2n+ 1

n (n+ 1)(bnπn (cos θ) + anτn (cos θ)) , (3.13)

where

πn (cos θ) =1

sin θP 1n (cos θ) (3.14)

τn (cos θ) =d

dθP 1n (cos θ) . (3.15)

P 1n is an associated Legendre polynomial, and the coefficients an and bn which characterize thescattered wave are given by

an =ψ′n (mx)ψn (x)− mψn (mx)ψ

′n (x)

ψ′n (mx) ζn (x)− mψn (mx) ζ ′n (x)

(3.16)

bn =mψ′

n (mx)ψn (x)− ψn (mx)ψ′n (x)

mψ′n (mx) ζn (x)− ψn (mx) ζ ′n (x)

, (3.17)

where m is the complex refractive index,

ψn (x) =

(

1

2πx

)12

Jn+ 12(x) (3.18)

ζn (x) =

(

1

2πx

)12

H(2)

n+ 12

(x) (3.19)

are Riccati-Bessel functions, and Jn+ 12and H

(2)

n+ 12

are spherical Bessel functions. The spherical

Bessel functions can be expressed in sine and cosine terms. This provokes rapid changes of S1and S2 for small variations of x. Some results for real refractive indices are shown in Figure 3.1.


Figure 3.1: Scattering diagrams from Mie theory. The solid curves are for i1 = |S1|2 and the broken

curves for i2 = |S2|2. The vertical scale is logarithmic. Values of i1 and i2 at 0

and 180 (whereboth curves are equal) are given beside the diagram. (Adopted from Goody and Yung 1989)

For x = 1 the curves show some of the features of Rayleigh scattering, particularly the large,positive polarization (i1 > i2) at scattering angels near 90. For x > 2 both positive andnegative polarizations occur, alternating more frequently as x increases. The forward scatteringis accentuated and often has a net negative polarization for large particles.

The total extinction and scattering cross sections can be obtained by integration over thescattering angle θ. Reasonable is the application of the dimensionless scattering and extinctionefficiencies Qs and Qe, being the cross sections normalized to the particle’s diameter, respec-tively,

Qs =2

x2

∞∑

n=1

(2n+ 1)(

|an|2 + |bn|

2)

(3.20)

Qe =2

x2

∞∑

n=1

(2n+ 1)Re (an + bn) (3.21)

The scattering efficiency for a sphere with a real component of the refractive indexmr = 1.33 andseveral values of the imaginary component mi are shown in Figure 3.2. For real refractive indices(mi = 0), Qs increases quartically for small values of x, corresponding to Rayleigh scattering,and attains a maximum near x ' 2π (λ ' r). The series of oscillations for increasing x results

3.3. THE MIE THEORY 27

Figure 3.2: Scattering efficiency from a dielectric sphere having a real component of the refractiveindex mr = 1.33, for several imaginary components mi, as functions of the dimensionless sizeparameter x (adopted from Salby 1996).

from the interference between light transmitted through and diffracted about the sphere. Forlarge particles, Qs takes approximately a constant value of 2. In this case, scattered radiationincludes radiation diffracted about the sphere as well as radiation that is redirected by reflectioninside the sphere. The weak dependence on the wavelength and the fact that the size of clouddroplets can span over several oscillations in Qs explain why clouds appear white.

If absorption is prevalent (mi < 0), the oscillations due to interference are damped outwith increasing mi, but the maximum near x ' 2π remains. The limiting value of Qs for largeparticles is reduced by attenuation inside the sphere. Energy that would otherwise contributeto the scattered wave field is absorbed.

Since in the atmosphere scattering particles only exist as a mixture of particles of variousradii, the macroscopic scattering and extinction coefficients are given by integration over thesize distribution n(r),

βs =

∫ ∞

0Qs (r)πr

2n (r) dr (3.22)


βe =

∫ ∞

0Qe (r)πr

2n (r) dr (3.23)

Some common size distributions are specified in section 3.4. The ratio of scattering and extinctioncoefficients of an aerosol population is called its single scattering albedo ω.

ω =βsβe

(3.24)

ω describes the influence of scattering on the total extinction of the aerosol. In the case of realrefractive indices the values of an and bn are real, and the single scattering albedo is 1. In generalω depends on the wavelength and the particle size. Irvine and Pollack [1968] studied the singlescattering albedo of water droplets and ice crystals.

The optical depth τ of an aerosol population of homogeneous density (e. g., a cloud) and athickness ∆z can be estimated as:

τ = βe∆z (3.25)

In contrast to Rayleigh scattering the phase function of Mie scattering has no analytical so-lution. Exact computation is only possible by numerical algorithms. For practical applications inradiative transfer calculations it is often convenient to approximate the Mie phase function by theHenyey-Greenstein phase function [for formal treatment see Van De Hulst 1980 and Liou 1980],

PHG (cos θ; g) =1− g2

(1 + g2 − 2g cos θ)32

(3.26)

=∞∑

l=0

(2l + 1) glPl (cos θ) , (3.27)

where θ is the scattering angle, Pl (x) is a Legendre polynomial, and g is the so called asymmetryfactor,

g =1

2

∫ +1

−1P (θ) cos θ d(cos θ). (3.28)

The asymmetry factor g of a phase function P characterizes the symmetry or the asymmetryof the scattering process. For isotropic scattering g = 0, for pure forward scattering g = 1, andfor pure backward scattering g = −1 are obtained. Irvine and Pollack [1968] listed asymmetryfactors of water droplets and ice crystals for several wavelengths and particle sizes. Asymmetryfactors of mineral dust aerosols are given in Tegen et al. [1996].

The Henyey-Greenstein phase functions succeed in reproducing the forward peak of Miescattering, but the backscattering behavior is improperly approximated. This shortcoming ispartially removed by using a double Henyey-Greenstein phase function,

P (cos θ) = bPHG (cos θ; g1) + (1− b)PHG (cos θ; g2) , (3.29)

where b is a positive fraction. g2 can be assigned a negative value to account for the backscatterpeak.

3.4. SIZE DISTRIBUTIONS 29

3.4 Size Distributions

As seen in the previous section the scattering characteristics, such as polarization and intensityof the scattered light depending on its wavelength, as well as the wavelength dependence ofthe aerosol extinction, are quite different for varying particle sizes. Since the size spectrumof a typical aerosol population is very wide and depends on formation, dispersion, conversionand decay processes model assumptions for continuous size distributions have to be done. Theextinction and scattering properties of a specific size distribution can be evaluated by using theMie Theory, and can then be compared with experimental results. Due to the large number ofparameters to be considered, it is impossible to directly derive the chemical composition andthe size distribution of an aerosol population from its measured optical characteristics.

Stratospheric aerosols predominantly consist of sulfate aerosols [see chapter 2.2]. In the fol-lowing it is assumed that the background aerosols comprise solely H2SO4 and H2O, and areliquid, spherical particles with a real refractive index of 1.43-1.44 [Palmer and Williams 1975;Russell and Hamill 1984]. The refractive index slightly depends on the wavelength of the incom-ing light, the temperature [Longhurst 1964], and the percentage of sulfuric acid, varying between68% and 86% [Steele and Hamill 1981]. For example, a typical stratospheric background aerosolcomposed of 75% of H2SO4 and 25% of H2O at a temperature of 218 K and a wavelength of532 nm has a refractive index of 1.43 [Pinnick et al. 1976]. For the consideration of volcanicaerosol with solid nuclei consisting of volcanic ash a second mode with a different refractiveindex (generally complex) has to be modelled analogously, and the extinction calculated fromthis size distribution must be added to the background contribution.

Russell et al. [1981] recommended nine different models of stratospheric aerosol size distri-butions. In the following the most widely used ones are presented.

Power Law Size Distribution

Stratospheric aerosols were first examined in detail by Junge et al. [1961]. He found that theiroptical properties can be adequately described by a size distribution that he assumed in earlierlaboratory studies for a large number of particle samples [Junge 1952, 1953]

n (r) = A

(

r

r0

)−ν

, (3.30)

where r0 is the unity radius. ν was found by Junge [1963] to be around 3.5±1. Since the numberof particles increased infinitely with decreasing radius, a truncated power law size distributionmust be applied. Russell et al. [1981] confined the radius range of stratospheric aerosols to 0.1 <r < 0.5µm. Within this range he estimated a value of ν of about 4. For r < 0.1µm or r > 0.5µmν is equal to 0.

To take into account a small but not vanishing number of large particles a segmented powerlaw size distribution can be introduced [e. g. Thomason 1991; Burton et al. 1999], with a largevalue of ν beyond a critical radius.


Exponential Size Distribution

To avoid the abrupt cut-off at a particular minimum radius Pinnick et al. [1976] suggested touse an exponential size distribution of the form:

n (r) = A exp

(

−r

r0

)

, (3.31)

where A is a constant related to the aerosol number concentration and r0 is a variable parameterrelated to the effective radius of the whole aerosol size distribution. For non-volcanic aerosolsa value of r0 = 0.075 µm is recommended. Rosen et al. [1978] predicated that a stratosphericaerosol generated by a growth process may be better described by an exponential than by apower law size distribution.

Modified Gamma Size Distribution

Deirmendjian [1964] was the first to use the modified gamma size distribution to compute scat-tering and polarization properties of water clouds and hazes in the visible and infrared. Laterhe noticed that the modified gamma function also agrees fairly well with measurements on mosttypes of aerosols [Deirmendjian 1969],

n (r) = Arα exp (−brγ) . (3.32)

The function is rapidly but continuously decreasing for small and for large particle sizes, andthus displays a concentration of the particle sizes around a certain radius. This phenomenonis typical for growth processes reaching an equilibrium with decay processes [see also chapter2.2]. The shape of the distribution is determined by the parameters α, b, and γ. The modifiedgamma, and even more the following lognormal size distributions are the most common modelsfor all types of tropospheric and stratospheric aerosols.

Lognormal Size Distribution

The mostly used size distribution model for all types of tropospheric and stratospheric aerosols isthe lognormal size distribution. It was introduced by Davies [1974] as an alternative solution tothe problem that the Junge power law does not account accurately for large and small particles,

n (r) =A

r lnσexp

−ln2(

rrm

)

2 ln2 σ

, (3.33)

where A again is related to the total number of particles and rm is called the mode radius. rmand σ are the variable parameters of the size distribution, and are related to the effective radiusreff and the effective variance veff [Hansen and Hovenier 1974] by

reff = rm exp

(

5

2ln2 σ

)

(3.34)

veff = exp(

ln2 σ)

− 1. (3.35)

Up to the present, numerous stray light and extinction measurements have been verified bya lognormal approach. Some of the first results from extinction measurements were published

3.4. SIZE DISTRIBUTIONS 31

Figure 3.3: Unimodal or bimodal differential (n (r)) and integrated (N (≥ r)) lognormal size distri-butions fitted to 1-km averaged particle counter measurements (solid circles) at 20 km aboveLaramie, Wyoming, for the spring of 1991 – 1996. The parameters of the size distributions andthe inferred surface area (A) and mass mixing ratio (m) assuming a composition of 65% H2SO4by weight, are shown in the figure. Uncertainties in the concentration measurements are shownby the vertical bars on the right-hand side. The goodness of the fit is shown at the bottom asthe sum of the logarithm of the ratios of measured and calculated number concentrations (Nm,Nc). Adopted from WMO [1999]

by Lenoble and Pruvost [1983]. From SAGE satellite data they derived values of rm ' 0.10 µmand σ ' 1.6 for background, and rm ' 0.045 µm and σ ' 1.9 for volcanic aerosols. This showsthat the size parameters of stratospheric aerosols significantly vary with volcanic activity. Ingeneral both, background and volcanic aerosols exist in the stratosphere. Thus, they cannot becharacterized by a unimodal lognormal size distribution.


The problem can be remedied by using a multimodal lognormal size distribution, which isthe sum of the several individual components in a particle mixture,

n (r) =I∑

i=1

Ai

r lnσiexp

−ln2(

rri

)

2 ln2 σi

. (3.36)

Each of the I components is characterized microphysically by its own particle number densityNi, and its own particular mode radii ri and mode width σ. The several modes emphasizethe different origins of the respective components. Usually, one or two modes are sufficient[Deshler et al. 1993; Deshler 1994; Grainger et al. 1995]. Thus, the aerosol size distribution canbe fully described by three or six parameters in the case of a unimodal or bimodal model,respectively. Only for high volcanism a third mode may be necessary [Bingen and Fussen 2000].

Furthermore, theoretical and field studies have shown that, the mean particle size of strato-spheric aerosol decreases with increasing altitude [e. g. Steele and Hamill 1981; Thomason 1991;Grainger et al. 1995; Hervig et al. 1998; Deshler 2001].

3.5 Aerosol Extinction Coefficients

As seen in the previous two sections, the computation of the extinction wavelength dependenceis possible by knowing the physical and chemical properties of the aerosols and their size dis-tribution. The implementable way pursues the opposite direction. Extinction coefficients areexperimentally measured, and the other characteristics must be derived. Since in general thereis no analytical solution for the wavelength dependence of the aerosol extinction, a model ap-proximation with only a few variable parameters is preferable.

The Angstrom Coefficient

More than two decades before Junge [1952, 1953, 1961] examined the particle size distributionof various aerosols, Angstrom [1929, 1930] investigated their extinction wavelength dependenceon the basis of both atmospheric transmission measurements and laboratory experiments. Heshowed that the extinction coefficient varies as a function of the wavelength within the visibleand near infrared spectral range, approximately following the power law

β (λ) = β [1µm]

(

λ

1µm

)−α

. (3.37)

β [1µm] is the extinction coefficient measured at the wavelength λ = 1µm, and depends on thespectral features of the aerosol particles as well as on the total particle number along the opticalpath. The Angstrom Coefficient α is closely related to the particle size distribution.

For the Junge power law size distribution (equation 3.30) a simple linear relationship be-tween the Junge parameter ν and the Angstrom coefficient α was found [Van De Hulst 1957;Bullrich 1964]:

α = ν − 2. (3.38)

This approximation is fairly well for large spherical particles with a real refractive index and2 < ν < 5, and for absorbing particles in the range 3 < ν < 4 [Volz 1956; Curcio 1961;McCartney 1977]. Empirical values for the aerosol extinction vary between 1 < α < 1.5 with a

3.5. AEROSOL EXTINCTION COEFFICIENTS 33

Figure 3.4: Left panel: Aerosol wavelength dependence exponent (−α) profiles observed by LIDARover Mauna Loa Observatory (MLO), Hawaii, for high and low Integrated Aerosol Backscat-ter (IABS) conditions. Right panel: Same for winter (DJF) and summer (JJA). Adopted fromBarnes and Hofmann [2001]

mean value of α ' 1.3 which corresponds to a Junge parameter of ν ' 3.3. In cases in which αis not close to 1 or the size distribution does not correspond to the Junge power law, the linearapproximation 3.38 fails [Quenzel 1970; Tomasi et al. 1983]. Figure 3.4 shows vertical profiles ofAngstrom coefficients inferred from aerosol backscatter measurements, using a LIDAR.

A Wavelength Dependence Model

The Angstrom formula 3.37 differs from experimental results as well, especially if large wave-length ranges are considered. Lenoble and Pruvost [1983] showed that the Angstrom coefficientα is not a constant, but a function of λ for a given aerosol size distribution. Therefore, [Yue 1986]proposed a two-parametrical wavelength dependence model, which better agrees with the resultsof Mie calculations for lognormal size distributions than the Angstrom model does,

β (λ) ∝ exp (C0 − C1λ) , (3.39)

where C0 and C1 are constants independent of λ.

A similar model was used by Brogniez and Lenoble [1988] to approximate the extinctionwavelength dependence measured by the satellite-borne SAGE II instrument. To be able to infertrace gas profiles from the observed extinction data, the aerosol extinction must be interpolatedfrom the aerosol channels to the channels responsible for the measurement of the trace gasabsorptions.


The Angstrom formula (3.37) is only valid in a small wavelength region around a measuredvalue λ0. Its logarithm has the form

ln (β (λ)) = ln (β (λ0))− α ln

(

λ

λ0

)

. (3.40)

3.5. AEROSOL EXTINCTION COEFFICIENTS 35

To cover a larger wavelength range a quadratic term is added which leads to the formulaused by Brogniez and Lenoble [1988],

ln (β (λ)) = ln (β (λ0))− α ln

(

λ

λ0

)

− b

(

ln

(

λ

λ0

))2

. (3.41)

This assumption is applied in the retrieval algorithms of the most important satellite-borne solaroccultation instruments, SAGE II [Brogniez et al. 1996, 1997] and POAM II [Lumpe et al. 1997]and III [Randall et al. 2001]. The inversion algorithm appendant to both instruments, and theneed to introduce a wavelength dependence model for the aerosol extinction is described in moredetail in chapter 6.2.

A third approximation of the same scheme was used by Fussen and Bingen [1999] to modelthe geographical change of the aerosol extinction and its dependence on volcanic activity basedon SAGE II data [also Bingen and Fussen 2000]:

β (λ) = exp(

a1 (λ− λ0) + a2 (λ− λ0)2)

(3.42)

If applied to a limited wavelength range all three models can match the same measured wave-length dependencies, since the aerosol extinction only slightly varies with the wavelength. Thetwo-parametric models either do not completely describe the extinction wavelength dependenceof an aerosol population, but the approximation is valid over a large wavelength range. Thus,the model serves for interpolating values between measurement channels but is not suitable forextrapolating to different wavelength ranges.

Chapter 4

Balloon Borne DOAS (DifferentialOptical Absorption Spectroscopy) ofDirect Sunlight

4.1 Light Attenuation in the Atmosphere

Solar radiation reaching the Earth is fundamental for all meteorological and climatic processes,the air composition and temperature at all altitudes, and many more phenomena which affectany form of life on Earth. The emission spectrum of the Sun can be approximated by a blackbody spectrum with a temperature of 5770 K, the mean temperature of the photosphere. Thephotosphere is a thin layer, only a few hundred kilometers deep, which emits almost all ofthe solar radiation. Superimposed on the smooth continuum are absorption lines produced bydifferent elements of the higher photosphere layers [Sonett et al. 1991]. The strongest lines inthe Sun’s spectrum are the so-called H and K lines due to ionized calcium (Ca2+) at 396.8 nmand 393.4 nm. Their confusing labels date from the original descriptions of the solar absorptionlines by Fraunhofer in 1817. Since the origin of the lines was not understood at that time, hesimply labelled the nine most remarkable lines in alphabetical order starting at the red end ofthe spectrum. Nowadays, the enormous number of absorption lines in the solar spectrum are allcalled Fraunhofer lines. Their characteristic series show that at least 63 elements occur in thesolar photosphere.

Passing through the atmosphere the light is attenuated by various absorption and scatteringprocesses. A continuous extinction is caused by Rayleigh [section 3.2], Raman [section 4.3.2], andMie [section 3.3] scattering, and by absorbing aerosols. Further, each molecule in the atmospherehas its own characteristic absorption cross section which can be more or less continuous withwavelength (e. g. O3), or composed of an ensemble of absorption lines (e. g. H2O) (see Figure4.6 and 4.7). Most notable is the absorption of UV light below 350 nm by ozone and oxygen1,and the strong infrared absorptions of H2O and CO2.

1Huggins band of O3: 310-350 nm, Hartley band of O3: 220-310 nm, Herzberg continuum of O2: 190-220 nm,Schumann-Runge band of O2: ≤ 190 nm, and the Schumann continuum of O2: ≤ 170 nm

37

38 CHAPTER 4. BALLOON BORNE DOAS OF DIRECT SUNLIGHT

Figure 4.1: Spectral distribution of the incident solar radiation outside the atmosphere and at sealevel. Major absorption bands of the most important atmospheric gases are indicated. Adoptedfrom [Boeker and van Grondelle 1995]

Lambert-Beer’s Law

The extinction of radiation by each substance is characterized by its particular wavelength-dependent extinction coefficient α (λ,~s). The light attenuation dI (λ) is proportional to thetotal light intensity I (λ), the extinction coefficient, and the width of the traversed layer ds.

dI (λ,~s) = −I (λ)α (λ,~s) ds, (4.1)

where the extinction coefficient of a substance is related to the extinction cross section σ (λ) fora particle, and the density of the substance n (~s) by

α (λ,~s) = σ (λ) · n (~s) . (4.2)

Neglecting the inelastic Raman scattering and the intensity of light scattered into the lightbeam from an arbitrary direction [section 5.1], the total extinction coefficient α (λ,~s) is the sum

4.2. DIFFERENTIAL OPTICAL ABSORPTION SPECTROSCOPY (DOAS) 39

of the partial extinctions due to Rayleigh and Mie scattering, and to the molecular absorbers i,

α (λ,~s) = σRayleigh (λ) · nAir (~s) +∑

i

σi (λ) · ni (~s) + αMie (λ,~s) . (4.3)

The total extinction along the light path L is obtained by integrating equation 4.1,

I (λ) = I0 (λ) exp

(

−

∫

L

(

σRayleigh (λ) · nAir (~s) +∑

i

σi (λ) · ni (~s) + αMie (λ,~s)

)

ds

)

, (4.4)

called Lambert-Beer’s law. I0 (λ) is the intensity of the unattenuated light. The logarithmic ratioof the light intensities is called the optical density (or depth) τ (λ)

τ (λ) = − lnI (λ)

I0 (λ). (4.5)

Thus, equation 4.4 can be written as

I (λ) = I0 (λ) e−τ(λ). (4.6)

4.2 Differential Optical Absorption Spectroscopy (DOAS)

Direct sunlight measurements applying Lambert-Beer’s law to infer trace gas absorptions re-quire the exact knowledge of the extraterrestrial solar intensity I0 (λ), and of the extinctiondue to Rayleigh and Mie scattering. It is a serious shortcoming that even for the most simplemeasurements with direct sunlight it is often not possible to do absolute measurements of I0.

If not a single wavelength but an entire spectrum is measured, theDifferential Optical Absorp-tion Spectroscopy (DOAS) technique can be applied to avoid above limitations [Platt et al. 1979;Platt 1994]. The principle of DOAS is to split the total extinction coefficient α into a low fre-quency component αb and a high frequency component α′ (differential extinction coefficient)with regard to the wavelength.

α (λ) = αb (λ) + α′ (λ) (4.7)

The extinction coefficients due to Rayleigh and Mie scattering vary only slightly with the wave-length, in particular continuously, whereas the molecular absorption cross sections show verystrong variations within a small wavelength range (see Figures 4.6 and 4.7). Thus, the influenceof both processes can be properly separated using the DOAS technique. Evaluating first the lowfrequency and then the high frequency component in Lambert-Beer’s law (4.4), it can be writtenas

I (λ) = I0 (λ) exp

(

−

∫

L

(

αb (λ) + α′ (λ))

ds

)

= I ′0 (λ) exp

(

−

∫

Lα′ (λ) ds

)

. (4.8)

The differential optical density τ ′ is given by

τ ′ (λ) = lnI ′0 (λ)

I (λ)=

∫

Lα′ (λ) ds. (4.9)

In a very simple approach, I ′0 (λ) can be determined by applying a high pass filter to I (λ) (seeFigure 4.2), to easily obtain τ ′ (λ). Actually, a polynomial of forth or fifth degree accuratelyapproximates the continuous wavelength dependence of Rayleigh and Mie extinction [chapters


Figure 4.2: Illustration of the DOAS principle. Narrow band structures of the absorption cross sectioncan be separated from broad band extinction.

3.3 and 3.2]. Therefore, the loss of information due to the high pass filtering can be avoidedapplying a polynomial fit to take into account the low frequency part.

Assuming that the temperature and pressure dependence of the absorption cross section ofthe trace gas i is negligible, the differential optical density τ ′i (λ) can be expressed by

τ ′i (λ) =

∫

Lα′i (λ,~s) ds = σ′i (λ)

∫

Lni (~s) ds = σ′i (λ)SCDi (4.10)

where the Slant Column Density SCDi of the trace gas i is defined as

SCDi =

∫

Lni (~s) ds (4.11)

Especially ozone and nitrogen dioxide show large temperature dependencies in some spectralranges [see section 4.3]. The use of two cross sections of the same substance at two differenttemperatures occurring along the light path, is sufficient to take into account the temperaturedependent changes of the absorption cross section.

The DOAS technique is especially suited to determine the influence of trace gas absorbersto the total extinction, since the broad-band extinction (including also the aerosol extinction)is omitted. By subtracting the trace gas absorptions of the total extinction, and correcting forthe contribution of Rayleigh scattering, the aerosol extinction can be directly measured over alarge wavelength range. The retrieval algorithm is explained in detail in chapter 5.

4.3 Retrieval of Vertical Trace Gas Profiles

During a typical DOAS balloon flight two different measurement geometries permit an altitudedependent observation of trace gases or aerosol extinctions. During the ascent or descent ofthe balloon, at lower altitudes more and lower atmospheric layers contribute to the measuredSCD than at higher altitudes (Figure 4.3). During the solar occultation (sunset or sunrise) theballoon remains at an almost constant altitude. For high solar zenith angles (SZA) the lightpath passes through lower atmospheric layers than for low SZAs (Figure 4.4). These differencesin the summation of the SCDs allow the inference of vertical profiles (see 4.3.4).

4.3. RETRIEVAL OF VERTICAL TRACE GAS PROFILES 41

Earth

Sun

Alti

tud

e

Balloon

Pos. 1

Balloon

Pos. 2

Trace gas profile

TroposphereStratosphere

Figure 4.3: Measurement geometry during the balloon’s ascent. Adopted from Osterkamp [1997]

During the ascent or descent for low solar zenith angles between 75 and 85 the solar light isonly marginally refracted on its way through the atmosphere. Thus, the computation of the lightpath [Raytracing, see 4.3.3] is quite straight forward. Since the ascent velocity of the balloondoes not exceed 10kmh the vertical resolution is very high. During the solar occultation, thelight passes a long way through the atmosphere. A very accurate measurement of the SCDs ispossible due to the large values of the traversed air masses, dominated by the contributions ofaltitude segments around the tangent point. Particularly gases with low concentrations like thehalogen oxides can be observed much better with the solar occultation technique. Though, thecalculation of the light path is much more complex. Small variations in the apparent SZA canlead to large differences in the tangent altitude. Thus, the vertical resolution is limited to around1 km.

Since both techniques measure the vertical trace gas profile for a different insolation, pho-tochemistry models can be directly verified in the stratosphere, and reaction rates measured inthe laboratory can be revised.


Earth

Alti

tud

e

Sun Pos.1

Sun Pos.2

Trace gas profile

Troposphere

Stratosphere

SZA 1

SZA 2

Figure 4.4: Measurement geometry during the solar occultation. Adopted from Osterkamp [1997]

4.3.1 The Slant Column Density

To derive the SCD of a trace gas from the measured total extinction, reference cross sections ofthe respective trace gas are needed. Since the spectra are recorded by a spectrograph of limitedresolution, in general having a different instrument function H (λ) than the spectrograph usedfor recording the reference spectra, either reference spectra recorded in the laboratory withthe own instrument at stratospheric temperature and pressure conditions are used, or a highlyresolved reference cross section σhigh (λ) is convoluted with the instrument function:

σref (λ) = H (λ) ∗ σhigh (λ) =

∫

H(

λ− λ′)

· σhigh(

λ′)

dλ′ (4.12)

The observed wavelength range is mapped to N discrete wavelength channels i of width 2∆λi

and center wavelength λci :

I (i) =

∫ λci+∆λi

λci−∆λi

I(

λ′)

dλ′ (4.13)

A detailed description of the effects of the instrument functions on the DOAS retrieval algorithmcan be found in Wenig [2001].


When the balloon reaches the maximum altitude (30 - 40 km, depending on the flight),the so called Fraunhofer reference spectrum is recorded for a low solar zenith angle. At thattime the solar light passes a minimum slant column to reach the detector. That way a solarspectrum I0 (λ) is measured with the instrument’s spectral resolution. For the evaluation ofballoon-borne measurements it is better to use this spectrum instead of a high-resolution solarspectrum [Kurucz et al. 1984] convoluted with the instrument function. So far, no highly resolvedextraterrestrial solar spectrum is available. Trace gas absorptions and extinction due to Rayleighand Mie scattering have to be eliminated from the Kurucz-spectrum as well.

The solar spectrum is highly structured, and the radiation intensity can vary from maximumvalues to zero within very small wavelength ranges. Further, the slit function of the spectrographis non-gaussian and asymmetric, and must be interpolated between the measurement channels.The convolution with this function is performed by numerical methods, and thus not being exact.Small errors in the slit function can change the convoluted low-resolution spectrum significantly.All these factors complicate the proper calculation of an extraterrestrial solar spectrum like itwould be measured by the applied instrument.

To take into account the possible differences of the wavelength-pixel mapping during theflight due to pressure and temperature changes of the spectrograph, each spectrum and eachreference cross section can be shifted and linearly stretched in relation to the Fraunhofer referencespectrum. In practice, the reference cross sections are reconciled with the Fraunhofer reference,and during the evaluation only the measured spectrum I is shifted and stretched.

The slant column densities of the trace gases j are now obtained by fitting the model functionF (i) to the shifted and stretched spectrum I? (i),

F (i) = ln I0 −m∑

j=1

SCDj · σj (i)− Pd (i) . (4.14)

The broad band structures of Rayleigh and Mie scattering are approximated by a polynomialPd (i) of degree d

Pd (i) =d∑

k=0

ck (i− ic)k (4.15)

with the center pixel ic of the considered spectral range.The spectral analysis consists of a non-linear Levenberg-Marquardt fit [Levenberg 1944 and

Marquardt 1963] to derive the shift and stretch parameters and a linear least square fit to derivethe parameters SCDj , in terms of minimizing the expression

χ2 =N∑

i=0

(

ln I? (i)− F (i)

εi

)2

, (4.16)

where εi is the measurement error of channel i. In general it is assumed that εi = ε = const.The influence of statistic noise and systematical residual structures on the fit is described inStutz and Platt [1996] and Hausmann et al. [1999].

For the numerical analysis of the measured spectra the software packages WinDOAS[Van Roozendael and Fayt 2000] and MFC [Gomer et al. 1996] were used, both providing theabove fitting procedure. WinDOAS additionally features the approximation of spectrographstray light by an polynomial intensity offset of a degree up to two [see chapter 5.1]. The coeffi-cients of this polynomial are additional parameters of the non-linear fit.


4.3.2 Error Sources of the DOAS Retrieval

Various errors due to the instrument’s properties and the measurement geometry contributeto the total error of the SCDs. These inaccuracies are not based on numerical errors in thefit process. A proper correction of the effects is not always possible, and can cause systemati-cal structures. In the following list only the most important error sources are mentioned. Thediversity is much larger and the effects are of varying importance for different instruments.

Instrumental Noise

The number of electrons generated by the photons hitting the photodiodes is subject to thePoisson statistic, causing the photoelectron noise. It is proportional to the square root of thetotal number of electrons. Also Poisson distributed is the noise of the photodiode dark current.Since the dark current is not changing with time, the dark-current noise is proportional to thesquare root of the integration time.

The preamplifier noise due to fluctuations in the current and voltage of the preamplifier, thereadout noise of a diode pixel due to switching capacities, and the analog-to-digital converternoise are constant contributions to the total noise.

For short exposure times, the dark current noise can be neglected. To minimize the photo-electron noise, a high light intensity is preferable.

A detailed description of the noise effects of the DOAS balloon spectrograph can be foundin Ferlemann [1998 and 2000].

Temperature Dependence of the Cross Sections

Most of the absorption cross sections show a strong dependence on the temperature in the UVand visible spectral range. Variations of the absolute value over the whole spectral range candirectly affect the retrieval of the SCDs, while changes in the shape of a cross section can result inlarge residual structures which adulterate the measurement of the aerosol extinction wavelengthdependence.

Burkholder and Talukdar [1994] investigated the temperature dependence of the ozone ab-sorption in the Chappuis band (410 – 760 nm). Near the maximum absorption around 550 –650 nm the cross section varies only slightly (< 1%). At larger and smaller wavelengths thetemperature dependence increases. At 420 nm the absorption cross section decreases by 40%when changing the temperature from 298 K to 220 K. In recent studies, the O3 absorption crosssection was measured between 203 and 293 K over a large wavelength range [230 – 850 nm byVoigt et al. 2001, 215 – 2400 nm by Bogumil et al. 2001] including the strong absorption bandsin the UV (see Figure 4.6).

Clearly, the ozone absorption features can be identified in the continuous wavelength rangeobserved by the DOAS spectrograph. The temperature dependence of the ozone absorptionis still a potential source of error. In our retrieval this error is minimized by including twonumerically orthogonal ozone reference spectra measured at different temperatures (−40C and−60C, see chapter 5.2).

The temperature dependence of the NO2 cross section in the visible spectral range hasbeen examined in a large number of studies [e. g. Davidson et al. 1988; Amoruso et al. 1993;Harwood and Jones 1994; Harder et al. 1997; Kirmse et al. 1997; Pfeilsticker et al. 1999]. The


absolute absorption cross section shows only small variations with temperature, since its broad-band shape remains almost constant. The differential cross section in contrast suffers from strongchanges, especially in the near UV and at the blue end of the visible spectral range. At 448 nmthe differential cross section increases almost linearly with temperature by almost 40 %, whencooling from 298 K to 200 K. At high spectral resolution also a large pressure dependence isobserved, but for the low resolution of the DOAS spectrograph this effect is negligible.

For the present study, O3 and NO2 reference absorption cross sections were recorded withthe present DOAS spectrograph in the laboratory for various temperatures [e. g. Harder 1999and Ferlemann 1998]. These spectra were used later on for the different DOAS retrievals.

The shape of the O4 collisional pair absorption cross section does not depend on pressure ortemperature, but the absolute value decreases by about 11 % when increasing the temperatureby 50 K [Osterkamp 1997 and Pfeilsticker et al. 2001]. For stratospheric balloon-borne trace gasand aerosol measurements the O4 cross section is included only to correct its absorption. Thus,its temperature dependence is irrelevant for the DOAS retrieval.

Spectrograph Stray Light

Usually various filters are integrated in a spectrograph to suppress the light outside the detectedwavelength range. Not always filter materials can be found whose transparency is exactly limitedto the desired range. A holographic grating is commonly used to spread the light of differentwavelength to different photodiodes. Light of larger or shorter wavelengths is diffracted intodifferent regions of the spectrograph, where it can be reflected or diffracted again. The samehappens to light from higher diffraction orders of the grating. The result is a continuous straylight offset, which can be corrected by including an additional intensity offset in the DOASfit procedure [see chapter 5.1]. This feature is integrated in the DOAS software WinDOAS[Van Roozendael and Fayt 2000].

Atmospheric Stray Light

Inelastic scattering processes on molecules are called Raman scattering. The wavelength of theincoming light is changed due to excitation of rotational or vibrational states of the molecule.Thereby the Fraunhofer lines in the solar spectrum are ”filled” by scattered light. This effectis called the Ring effect [Grainger and Ring 1962]. Strong atmospheric absorption lines can beaffected by the Ring effect as well [Fish and Jones 1995]. Calculations of the Raman cross sectionfor nitrogen and oxygen can be found in Bussemer [1993], Haug [1996], and Funk [2000].

The effect is usually corrected by including an additional spectrum in the fitting process[Solomon et al. 1987]. This Ring spectrum is obtained by polarization measurements2 or modelcalculations [Bussemer 1993; Funk 2000]. For direct sunlight measurements the Ring effect isgenerally supposed to be negligible [Bauer 1997; Pundt et al. 1998].

Rayleigh and Mie scattering are elastic processes, and do not change the wavelength ofthe incoming light. Thus, light scattered into the light beam by those elastic processes doesnot affect the differential optical density, since the scattering cross sections have a broad-bandstructure. However, the extinction due to Rayleigh and Mie scattering can be underestimated

2The Raman scattered light is much less polarized than the Rayleigh scattered light. Using polarization filtersit is possible to distinguish between both effects.


not considering the additional incoming light. For direct sunlight measurements these effects arenegligible, since the ratio of the stray light to the total intensity is very small.

Nonlinearity of the Spectrograph

The total extinction, which is the ratio of the Fraunhofer reference spectrum and the measuredspectrum (see above sections of this chapter), varies significantly over the whole observed wave-length region. Thus, a nonlinear response of the detector to the incoming light intensity canalter the shape of the extinction wavelength dependence, and therefore the inferred trace gasSCDs and the aerosol extinction wavelength dependence.

The nonlinearity of the applied balloon-borne DOAS spectrograph was examined byBauer [1997] and Ferlemann [1998]. It is negligible up to 90 % of maximum signal, but thenincreases drastically. Since a large signal is desirable to suppress noise effects, it is automaticallyregulated to values of some 80 % [see section 4.4.3].

4.3.3 Raytracing

For the numerical retrieval of the vertical distribution of a chemical species from the mea-sured slant column densities, the atmosphere has to be partitioned into layers of constant con-centrations of the species. The vector

−−−→V CDk of dimension n summarizes the vertical column

densities of the substance k in the n atmospheric layers. The slant column densities of thespecies k obtained by the DOAS fit are combined in the vector

−−−→SCDk of dimension m, where

m is the number of measured spectra. The relation between−−−→SCDk and

−−−→V CDk is given by

the Air Mass Factor matrix AMF,

−−−→SCDk =

AMF11 · · · AMF1n...

. . ....

AMFm1 · · · AMFmn

·−−−→V CDk. (4.17)

The component AMFij of the matrix can be seen as the ratio of the length of the light paththrough the atmospheric layer j of spectrum i and the width of the j.

In the easiest model atmosphere with plane and parallel layers and constant pressure andtemperature throughout the atmosphere, the airmass factors for all measurements and layersare given by 1/ cos (SZA). In a real atmosphere the temperature and especially the pressureare subject to changes of several orders of magnitude. Because of the temperature and pressuredependent refractive index of air and the spherical geometry of the atmosphere, a raytracingprogram has to be used to calculate the AMF matrix. The program damf [Schulte 1996] com-putes the airmass factors and retrieves the vertical profiles by inverting the AMF matrix [seesection 4.3.4].

The choice of the division of the atmosphere into segments of constant trace gas concen-trations should not affect the calculation of the optical path. Thus, the refraction index, andtherefore temperature and pressure, are treated as continuous functions of altitude. Horizontalvariations are neglected in the model.

A differential equation for the light path is derived from Fermat’s principle,

δ

∫

n (~x) ds = 0, (4.18)


where n (~x) is the refractive index of the air, which is only very little larger than 1[Penndorf 1957], but slightly varies with pressure and temperature, and thus depends on thelocation ~x. The light path is smooth and only little bent. Hence, a Runge-Kutta-algorithm (hereof fourth order [Bronstein and Semendjajew 1991]) is sufficient to solve equation 4.18, and thestepsize can be chosen constant.

4.3.4 AMF Matrix Inversion

For the inversion of the airmass factor matrix, a singular value decomposition (SVD) is used[Press et al. 1995]. The SVD algorithm represents a generalization of the principal axis transfor-mation to arbitrary real matrices. It is applied to solve most linear least-square problems, andwell suited to retrieve vertical profiles since in most cases there are more measured spectra thanaltitude segments.

The SVD method is the decomposition of an M ×N matrix A, whose number of rows M isgreater than or equal to its number of columns N , into a column-orthogonal M ×N matrix U,a diagonal N ×N matrix W, and the transpose of an orthogonal N ×N matrix V,

A = U ·W ·VT. (4.19)

The pseudo-inverse of the matrix A (an N ×M matrix) is defined as

A† =(

AT ·A)−1

·AT. (4.20)

Since the transpose of an orthogonal matrix V is equal to its inverse, the pseudo-inverse of Ausing the SVD can be expressed as

A† = V ·W−1 ·UT, (4.21)

where the inverse of W is given by

W−1 =[

diag (wj)]−1

=[

diag(

w−1j

)]

, (4.22)

and wj , 1 ≤ j ≤ N are the diagonal elements of W.If one of the wj is zero or so small that its value is dominated by numerical roundoff errors,

this construction poses problems. The condition number is defined as the ratio of the largestof wj to the smallest of wj . A matrix is singular (it has at least one Eigenvalue of zero) if itscondition number is infinite, and it is ill-conditioned if its condition number is too large, that is,if its reciprocal approaches the machine’s floating-point precision (e. g. 10−6 for single precisionor 10−12 for double precision).

To invert the airmass factor matrix in principle a linear equation of the form

A · ~x = ~b (4.23)

has to be solved. Using the SVD the vectors with the smallest length |~x|2 are singled out byreplacing w−1

j by 0 if wj = 0. If ~b is not in the range of A the best estimate of a solution is givenby a least square fit which minimizes

r =∣

∣

∣A · ~x−~b∣

∣

∣ . (4.24)


Pointing system

on boardcomputer

Hg Cd Tedetector

In Sbdetector

Damping ball

Fixedmirror

Movingmirror

Interferometer

Pointingsystem

electronics

Heliostat(sun tracker)

Acquisitionmirror

Sun sensor

CCD camera

Folding mirrorGe window

Cd Nibattery

Telemetry / TelecommandCNES

DOAS

fiber

Figure 4.5: Schematic drawing of the LPMA/DOAS balloon gondola. The acquisition mirror followsthe Sun with a precision of 160 degree and reflects the light onto the folding mirror which redirectsthe light directly into the FTIR and the telescopes of the DOAS spectrograph

The main source of error of the airmass matrix inversion is the SCD error. Mapping theSCD vector onto the VCD vector each of its components is amplified by the reciprocal of thecorresponding Eigenvalue. Thus, a component corresponding to a relatively large Eigenvaluemay be hidden among the noise from the contributions of small Eigenvalues. The relative errorof the SCD measurements is amplified by the condition of the AMF matrix. Therefore, it isreasonable to confine the condition to a certain maximum by setting the reciprocals of smallEigenvalues to zero. The VCD error can then be calculated by

∆V CDi =

√

√

√

√

M∑

j=1

(

(AMF )−1ij ·∆SCDj

)2for 1 ≤ i ≤ N (4.25)

4.4. THE DOAS BALLOON SPECTROGRAPH 49

4.4 The DOAS Balloon Spectrograph

The aerosol extinction profiles presented in this thesis have been inferred from data col-lected by a DOAS double spectrograph for balloon-borne measurements [Ferlemann 1998;Ferlemann et al. 2000; Harder 1999]. The instrument was primarily developed for direct sun-light measurements of stratospheric trace gases applying the DOAS technique. The coveredwavelength range in near ultraviolet (0.320 - 0.422 µm) and in visible (0.417 - 0.670 µm) facili-tates the observation of O3, O4, NO2, NO3, HNO2, H2O, HCHO, ClO, OClO, BrO, IO, and SO2(Figures 4.6 and 4.7) throughout the stratosphere, and in the upper troposphere. The spectro-graph is launched aboard the LPMA/DOAS3 balloon gondola together with an infrared FourierTransform Interferometer (FTIR). Up to the present, nine flights have been performed success-fully at three different locations [chapter 6]: Leon (Spain, 42.6N 5.7W), and Gap (France,44.0N 6.1E) at mid-latitudes, and Kiruna (Sweden, 67.9N 21.1E) at high latitudes.

4.4.1 The LPMA/DOAS Balloon Gondola

The balloon gondola used to accommodate the LPMA and the DOAS instruments goes back toan azimuth controlled gondola developed by the Observatoire de Geneve for astronomical obser-vations. Later it has been optimized for atmospheric applications by Camy-Peyret et al. [1995].The gondola is rotatably attached to a large stratospheric balloon (100000 – 400000 m3) floatingahead. A gyroscope stabilizes the gondola in azimuthal direction to within a range of ±1 degree.The fine-pointing of the Sun’s observation is performed by the active Sun tracker (in Frenchcalled Heliostat) [Hawat et al. 1995] which allows to point to the Sun to within a range of ± 1

60degree. The Sun tracker directs the sunlight into the evacuated Fourier transform interferometerand both telescopes of the DOAS spectrographs.

The FTIR is a DA2 Michelson type interferometer customized for balloon operationsby the French LPMA team. It measures the spectral signatures of O3, H2O, CO2, CH4,NO, N2O, HNO3, ClONO2, and CCl2F2 in mid-infrared, and those of O2, H2O, NO2, HCl,CH4, and HF in near-infrared [Camy-Peyret et al. 1995; Payan et al. 1998; Payan et al. 1999;Camy-Peyret et al. 1999].

A GPS antenna is installed aboard which records the balloon trajectory to ±2 meters inthe horizontal and ±10 meters in the vertical direction. On-board sensors measure the ambienttemperature and pressure, an information used to remove the small GPS-jitter in post-flighttrajectory calculations. Finally, since 1998 an ECC4 ozone sonde is assembled aboard whichallows to verify the DOAS ozone profile.

4.4.2 Optical and Mechanical Properties of the Spectrograph

The DOAS balloon spectrograph was optimized for airborne applications [Ferlemann 1998;Ferlemann et al. 2000; Harder 1999], with particular emphasis put on low mass (42 kg), lowpower consumption (30 W), low spectral drift of the optical imaging system due to ambienttemperature and pressure changes, low detector noise, and low spectrograph stray light. Theserequirements were met by enclosing the entire spectrograph into a pressurized and thermostatedstainless steel container. Prior to launch, it is evacuated to a pressure of around 10−6 mbar,

3LPMA = Laboratoire de Physique Moleculaire et Applications4Electro-Chemical Cell


2

4

6

8

10

12

NO2 (x10-19)

300 320 340 360 380 400 4202

4

6

8

10

12

SO2(x10-19)

Wavelength [nm]

300 320 340 360 380 400 420

0

5

10

15

0

5

10

15

OClO (x10-18)

300 320 340 360 380 400 420

0

2

4

6

8

0

2

4

6

8

Ab

so

lute

Cro

ss

Se

cti

on

(c

m

2 ) HCHO (x10-20)

300 320 340 360 380 400 420

0

5

10

15

20

BrO (x10-18)

0,00,51,01,52,02,53,03,5

IO(x10-17)

300 320 340 360 380 400 4200

2

4

6

8

0

2

4

6

8

HNO2(x10-19)

300 320 340 360 380 400 420-25-24-23-22-21-20-19-18-17-16

-25-24-23-22-21-20-19-18-17-16

logO3

300 320 340 360 380 400 420

0

2

4

6

0

2

4

6

ClO (x10-18)

Figure 4.6: Absolute cross sections of the chemical species in the UV balloon double spectrographwavelength range (320 - 422 nm). Adopted from [Bauer 1997].


450 500 550 600 650

0.0

0.5

1.0

1.5

2.0

0.0

0.5

1.0

1.5

2.0

450 500 550 600 650

0

1

2

3

4

5

0

1

2

3

4

5

450 500 550 600 650

0

2

4

6

8

0

2

4

6

8450 500 550 600 650

0.00.51.01.52.02.53.03.5

0.00.51.01.52.02.53.03.5

450 500 550 600 650

0

5

10

15

20

25

30

0

5

10

15

20

25

30

450 500 550 600 650

0.90

0.95

1.00

1.05

1.10

0.90

0.95

1.00

1.05

1.10

O4(x10-46cm5/molec2)

Wavelength [nm]

OClO (x10-18)

NO2 (x10-19)

IO(x10-17)

NO3 (x10-18)

450 500 550 600 650

-28

-26

-24

-28

-26

-24

Abs

olut

e C

ross

Sec

tion

[cm

2 ]

logH2O

O3(Differential Optical Density)

Figure 4.7: Absolute cross sections of the chemical species in the Vis balloon double spectrographwavelength range (417 - 670 nm). Adopted from [Bauer 1997].


Figure 4.8: Schematic drawing of the DOAS double spectrograph. The total weight of the instrumentincluding 10 liters of ice water mixture is approximately 42 kg, the total size is 45 x 45 x 55 cm3.

and even though small leakages exist it is maintained at less than 10−2 mbar during the wholeflight. The temperature is stabilized by an ice water mixture surrounding the container. Thetemperature of the photodiode arrays is fixed to −10 ± 0.03C by peltier elements. The warmside is cooled by a refrigerant cycled through the ice water.

The sunlight collected by the Sun tracker is fed into two small telescopes to which quartzfibre bundles are attached. The Telescopes contain (a) filters for bandwidth restriction, andto guarantee an appropriate signal also at short wavelengths which are strongly attenuated byRayleigh scattering, (b) diffusers to average the radiance over the solar disc, and (c) skimmersto match the f-numbers of the quartz fibre bundles and the spectrographs. Within the quartzfibre bundles, the f-number of the light remains approximately constant as they are only around5 meters long. The exits of the quartz fibre bundles form a rectangular slit (100 x 2000 µm) fromwhich the light is directed on two holographic gratings. The two holographic gratings analyzethe sunlight either in the near UV (0.320 - 0.422 µm) or visible (0.417 - 0.670 µm) spectral


range. The holographic gratings disperse and focus the light onto two photodiodes array with1024 channels each 25 µm wide. Accordingly, the FWHM5 resolutions of the spectrographsare 0.45 nm (4.5 detector pixels) for the UV and 1.48 nm (5.7 detector pixels) for the visiblespectrograph. It thus fulfills the sampling criteria given by Roscoe et al. [1996], stating thatfor a proper DOAS application the FWHM resolution should match at least 4.5 detector pixels.Particular care was put on reducing the spectrograph stray light by suppressing the higher-orderand zero-order grating reflections using light traps and a dispersive prism pre-analyzer in frontof the UV telescope [Vradelis 1998].

A MC68332 CPU module controls the whole measurement process, i. e. the read-out elec-tronics, the peltier elements, and the exposure time to obtain an optimal signal. A more de-tailed description of the electronics and the programming can be found in the Ph. D. thesis ofFerlemann [1998].

4.4.3 Measurements of Relative Intensities

Evidently the apparent solar intensity changes over the course of a flight. Therefore, the exposuretime to record a single spectrum has to be adapted to the available sunlight. At high solar zenithangles, i. e. solar occultation, exposure times quickly change from several seconds to up to twominutes.

The self-sufficient single-board computer MC68332 CPU module (shortly called 68k) con-tains a control program written on its EPROM6. In the automatic mode, the control programcalculates the new exposure time from the maximum signal of the preceding spectrum, in thatthe maximum signal of the present spectrum is 80 % saturated7. As shown by [Ferlemann 1998]unwanted nonlinear effects of the photodiode array occur for signals exceeding 90 % of themaximum signal.

Evidently, a reliable measurement of relative intensities requires similar good records of thesignal as function of the exposure time. Hence, nonlinear effects of the photodiode or a wronglyrecording of the exposure time are highly unwanted.

Past Deficits in the Control Program

Unfortunately, the read-out controlling program of the 68k showed some serious deficits in theversion used up to March 2001 (version 2.92 and previous). These deficits mainly affected thecorrect storage of the total exposure time by (a) incorrectly terminating the record of the lastspectrum in a sequence of spectra for a given total exposure time, (b) resetting the exposure timeto zero after a so called dummy-readout but keeping the recorded intensity, and (c) resetting theintensity of an oversaturated spectrum to zero while adding the exposure time of that particularspectrum to the total exposure time of a spectrum recording sequence.8

5FWHM = Full Width at Half Maximum6EPROM = Electrically Programmable Read Only Memory7The photodiodes are operated with a reverse-bias of 2 V. Since reverse-biasing inhibits power flow, the

photodiode acts as a capacitor, which is then charged by around 120 · 106 electrons. Photoelectrons released bythe insolation discharge the photodiode. When the photodiode is discharged completely, the maximum signal isreached. The read-out electronics measures the difference between the voltage after the insolation and the 2 Vreference voltage. The response is linear up to 90 % of the maximum signal

8Note that DOAS measurements are in any case not affected by incorrect storage of exposure times, sinceDOAS does not use time information.


0

1

2

3

4

5

6

7

8

9

10

Num

ber

of S

cans

4000 4100 4200 4300 4400 4500

65000

70000

75000

80000

85000

90000

Cou

nts

per

Tim

e [s

-1 ]

Spectrum Number - 220000

Figure 4.9: Demonstration of the effect of the dummy-read on the submitted exposure time. Theblack dots accord to the averages of the spectra divided by the submitted exposure time, theblue ones indicate the result after division by the corrected exposure time. The red dots representthe number of performed scans, each of 0.4 s exposure time. The variations at constant scannumbers are due to fluctuations in the lamp’s intensity. The lamp shows a decreasing intensitywith time, due to a voltage decrease of the power supply. The numbers of the spectra correspondto a IUP (Institut fur Umweltphysik, Heidelberg) internal numbering of the spectra recorded bythe DOAS balloon spectrograph.

Details to (a): The exposure time is calculated for each scan respectively from the signalof the former scan to reach the optimum of 80 % of the maximum signal. A series of scans isperformed until the remaining time of the user given total exposure time is shorter than thecalculated time for the next scan. The submitted spectrum consists of the sum of all scans, sosingle scans cannot be analyzed. If the remaining time is longer than 75 % of the required timefor the scan, the scan is performed and terminated after the remaining time has elapsed. InVersion 2.92 the scan is always performed completely, but only the given total exposure time issubmitted. Thus, the submitted exposure time can be up to 25 % of the scan time shorter thanthe actual exposure time, without a chance of correcting the discrepancy. Hence, every spectrumexposed with exactly the given total exposure time is expelled from the aerosol retrieval.


0 100 200 300 400 500 600 700 800 900 1000 0.0050

0.0052

0.0054

0.0056

0.0058

0.0060

0.0062

0.0064

0.0066

0.0068

0.0070

0.0072

ln(S

pect

rum

212

313/

Spe

ctru

m 2

1231

5)

Channel

0

50000

100000

150000

200000

250000

300000

Spe

ctru

m 2

1231

5

Wavelength [nm]

420 440 460 480 500 520 540 560 580 600 620 640 660

Figure 4.10: Illustration of the effect of nonlinearities in the photodiodes exceeding 90 % of saturation.The black curve shows the logarithmic ratio of the two spectra 212313 and 212315 In spectrum212315 the first scan is discarded, so nonlinear effects are prevented, in spectrum 212313 thefirst scan is included. The bias of 0.0051 is caused by differences in the real time exposure. Bothspectra are recorded within 11 s. During this time intensity changes can be neglected. Spectrum212315 is plotted in red to point out the dependence of the nonlinearities on the signal andthe Fraunhofer line structures in the black curve. The balloon spectra 212313 and 212315 wererecorded during the Kiruna balloon flight on February 18, 2000, at float altitude and a solarzenith angle of 86.6

Details to (b): Before each series of scans, a dummy-readout of 60 ms is performed. In version2.92 the photodiode array is not reset subsequently. Thus, in every spectrum the first scan isexposed 60 ms longer than the calculated exposure time for that scan. This effect can easilybe corrected by adding 60 ms to the submitted exposure time of each spectrum, respectively.In Figure 4.9 a series of spectra recorded with a halogen lamp in the laboratory for differentnumber of scans with constant exposure time is shown. The average light intensity over thevisible spectral range is calculated first dividing by the submitted and then by the correctedexposure time.


During a typical balloon flight, the light intensity at float altitude (30 - 40 km) is veryhigh. The photodiodes reach 80 % of saturation after around 0.3 s of exposure. At such a highinsolation an additional exposure time of 60 ms can lead to signals in the nonlinear range ofthe photodiodes. To avoid errors in the spectrum, a scan with a signal in the nonlinear rangeis automatically abolished. In version 2.92 the exposure time of a discarded scan is neverthelessadded to the total sum. Such spectra can easily be eliminated by the strong divergence from thecontinuous intensity trend throughout the flight.

Details to (c): In version 2.92 the rejection criteria for spectra in the nonlinear range, isa signal of 100 % saturation. In version 3.01 it is lowered to 90 %. Hence, at high insolation,there may be nonlinearities for the first scan of some spectra. In Figure 4.10, the ratio of twospectra recorded at maximum light intensity during the Kiruna balloon flight on February 18,2000 is shown. In the DOAS retrieval of trace gases, the broad-band structure is compensatedby the fitted polynomial, and the Fraunhofer line structures are compensated by the stray lightoffset. For the aerosol retrieval, only the broad-band structure is relevant. As can be seen inFigure 4.10, the variance of the optical density is around 0.002. As this effect only occurs atfloat altitude and for low solar zenith angles (high light intensity), it only affects the aerosolobservations above the float altitude, but it contributes to the measured extinction of the wholeflight, since the extinction of the Fraunhofer spectrum is determined by a Langley plot [chapter5]. Typical aerosol extinction optical densities along the light path are around 0.5 up to valueslarger than 1 for the relevant tangent altitudes within the Junge layer. So this effect can beconsidered to be negligible.

All the mentioned deficits are removed in version 3.01, now installed in the balloon spectro-graph electronics.

Chapter 5

Retrieval of Vertical AerosolExtinction Profiles

The method to infer aerosol extinction profiles from measured total atmospheric extinctions relieson a precise measurement of trace gas absorption and Rayleigh scattering of the atmosphere.The basic underlying formula goes back to Lambert-Beer (equation 4.4),

I (λ) = I (λ)−Istraylight (λ) = I0 (λ) exp

(

−σRayleigh (λ) · SCDair −∑

i

σi (λ) · SCDi − αMie (λ)

)

.

(5.1)with a correction due to spectrograph stray light (Istraylight (λ)), I (λ) being the measured spec-

trum, I (λ) the spectrograph stray light corrected spectrum, I0 (λ) the so called Fraunhoferreference spectrum recorded at low total air column (balloon float altitude), and attenuationterms due to Rayleigh scattering (σRayleigh (λ) · SCDair), the absorption of the different tracegases i (

∑

i σi (λ) · SCDi), and Mie scattering (αMie (λ)).Mie extinction is evaluated by firstly correcting the stray light from the measured spectrum.

Then, both the measured spectrum and the Fraunhofer reference spectrum are normalized withrespect to their total and if necessary corrected exposure times (as described in chapter 4.4.3).The negative logarithm of the ratio of (I (λ)− Istraylight (λ)) and I0 (λ) is taken as the totaloptical density αtotal (λ) of all processes contributing to the total extinction (equation 4.5). TheMie extinction is then evaluated by subtracting the trace gas absorptions due to O3, NO2, O4and the calculated Rayleigh scattering from the total extinction αtotal (λ).

Eventually, Rayleigh scattering dominates the whole atmospheric attenuation as it stronglyincreases with decreasing wavelength (∝ λ−4, see section 3.2). It then comes clear that theaerosol retrieval should be performed where the errors in the inferred Rayleigh scattering arestill much smaller than the expected Mie extinction. Therefore, in the present study the Mieanalysis is performed from spectra recorded with the visible spectrograph, i. e. from 440 to 615nm, corresponding to 700 channels.

The trace gas absorptions are obtained by a DOAS retrieval on the measured spectra [chap-ter 4]. Since the cross sections of all gases absorbing in the considered spectral range are wellknown, the retrieval yields the SCDi for each of the gases. More details are given in chapter 5.2.The extinction due to Rayleigh scattering is calculated from the known Rayleigh scattering crosssection (equation 3.8 3.8) and calculated air SCDs. The air SCD is calculated by simulating theline of sight in a refractive spherical atmosphere in the raytracing program damf [Schulte 1996

57

58 CHAPTER 5. RETRIEVAL OF VERTICAL AEROSOL EXTINCTION PROFILES

440 460 480 500 520 540 560 580 600 1.5

2.0

2.5

Wavelength [nm]

(c)

Opt

ical

Den

sity

0.1

0.2

1.0

1.5

(b)

(a) In

tens

ity [

10 5 C

ts/s

]

Figure 5.1: First three steps of a typical aerosol evaluation: (a) The Fraunhofer reference spectrumI0 (λ) recorded at a solar zenith angle of 89.07 and at an altitude of 30.15 km. (b) The spectrumI (λ) recorded at a solar zenith angle of 93.8, corresponding to a tangent altitude of 15.8 km.(c) The black curve shows the logarithm of the ratio of I0 (λ) over I (λ); the red curve showsthe same but with stray light corrected I (λ). The spectrum is taken from the balloon flight inKiruna on February 18, 2000.

and chapter 4.3.3] with atmospheric temperature and pressure data measured by on-board sen-sors, radiosondes, or analyzed meteorological fields. For altitudes above the atmospheric rangecovered by the instruments, the US Standard Atmosphere [1976] is used to approximate the at-mospheric conditions. Since at these altitudes, the atmospheric density is very low, departuresfrom the actual conditions propagate only weakly into the determination of the ”zero air mass”extinction (detailed in chapter 5.4).

Figures 5.1 and 5.2 shows the individual steps of the aerosol retrieval for a sample spectrumrecorded during the balloon flight in Kiruna (northern Sweden) on February 18, 2000, for a solarzenith angle of 93.8, corresponding to a tangent altitude of 15.8 km.

59

440 460 480 500 520 540 560 580 600

0.4

0.5

(h)

Wavelength [nm]

1

2

(g)

1

2

(f)

Opt

ical

Den

sity

1

2

(e)

1

2

(d)

Figure 5.2: Trace gas absorption and Rayleigh scattering removal. In figure (d) – (g) the red curvesshow the remaining optical density after subtracting the contribution of the respective extinctionprocess due to O3 (trace d), NO2 (trace e), and O4 (trace f) absorption, and Rayleigh scattering(trace g) from the previous result (black drawn line). Trace (h) shows the resulting Mie extinctionin black and the fitted approximation (3.41) in red.


5.1 Spectrograph Stray Light Correction

The spectrograph stray light intensity is obtained from the DOAS retrieval of the trace gasabsorptions [see chapter 4.3]. A linear intensity offset included in the trace gas evaluations isfound to be sufficient to eliminate the influence of the spectrograph stray light,

Istraylight (λ) = (c0 + c1 · (λ− λ0)) · Imean, (5.2)

where c0 and c1 are the fitting parameters. Imean is the mean value of the measured spectrumI (λ). Hence, the value of the parameter c0 corresponds to the ratio of the stray light intensityand the total measured intensity at the middle channel of the analysis range at a wavelength λ0(around 530 nm for this evaluation).

Figure 5.3 shows values of c0 as a function of SZA for all investigated eight DOAS balloonflights performed till 2000. The linear fitting parameter c1 is usually about 0.1% of c0 indicatinga small variation of the spectrograph stray light with wavelength of about 10% over the analyzedwavelength range. Figure 5.3 illustrates that the spectrograph stray light is about ±0.003 of thetotal intensity throughout most of the solar occultation, and only for very large SZAs (tangentaltitudes close to the tropopause) c0 is dropping to higher negative values. This behavioursuggests that the spectrograph stray light is originates mainly from reflections of UV lightentering the UV spectrograph, i. e. a cross-talk of both optically not totally sealed setups. Athigh sun the UV insolation and the spectrograph stray light cross-talk are expected to be largerthan at low sun, and accordingly the spectrograph stray light should then assume negativevalues. In any case, the spectrograph stray light only causes small variations in the retrievedtotal optical density (Figure 5.1c). Spectra containing unreasonably large spectrograph straylight contribution (of several percent) are in any case discarded because they tend to suffer frominsufficient light intensity.

5.2 Elimination of Trace Gas Absorptions

The contribution to the extinction by gas absorption is determined from the SCD of each gas byperforming a DOAS fit with the programWinDOAS [chapter 4; Van Roozendael and Fayt 2000].

The major stratospheric absorbers in the visible spectral range are summarized in Figure4.7. The most important species is ozone, which reaches optical densities of more than 1 in deepoccultation spectra. Smaller but significant absorptions in certain wavelength ranges are causedby NO2 and O4. NO3 absorption is not important as significant absorption bands are locatedbeyond the considered wavelength interval. In the spectrograph’s spectral interval O2 γ-band(628 nm) absorbs as well, but in order to avoid its troublesome contribution the analysis rangeis limited to 615 nm.

OClO is also absorbing but its optical densities along the line of sight are usually very small(below 2 · 10−3), and the strongest absorption bands are below 440 nm Fitzenberger [2000].Likewise, IO and OIO is found to absorb by as much as the detection limit of the DOAS balloonspectrograph for stratospheric observations [Bosch 2001].

Stratospheric water vapor occurs in very small concentrations (3 – 6 ppm), but its visibleabsorption were below the detection limit even for solar occultation observations. It was alsofound that omitting the H2O reference in the fit did not change the inferred SCD of the othertrace gases. Hence, only the absorptions of O3, NO2, and O4 have to be considered here.

5.2. ELIMINATION OF TRACE GAS ABSORPTIONS 61

-0.008

-0.006

-0.004

-0.002

0.000

0.002

0.004

-0.008

-0.006

-0.004

-0.002

0.000

0.002

0.004

-0.008

-0.006

-0.004

-0.002

0.000

0.002

0.004

-0.008

-0.006

-0.004

-0.002

0.000

0.002

0.004

-0.008

-0.006

-0.004

-0.002

0.000

0.002

0.004

-0.008

-0.006

-0.004

-0.002

0.000

0.002

0.004

86 88 90 92 94 96-0.008

-0.006

-0.004

-0.002

0.000

0.002

0.004

86 88 90 92 94 96-0.008

-0.006

-0.004

-0.002

0.000

0.002

0.004

León 1996

Kiruna 1997

Kiruna 1999

León 1998Gap 1997

c 0

c0

Kiruna 1998

Kiruna 2000Gap 1999

SZA [deg]

Figure 5.3: Relative stray light offset c0 at the central wavelength λ0 for the eight DOAS balloonflights performed till 2000. Only the solar occultation spectra after recording the Fraunhoferreference spectrum are shown. For the two flights in Gap (France) the Fraunhofer referenceswere recorded for SZAs lower than 86. Occultation measurements during sunset are plotted inred, during sunrise they are plotted in blue.


Figure 5.2 illustrates the subsequent removal of the contributing trace gas absorptions (O3trace (d), NO2 trace (e), and O4 trace (f)). In panel (d) the black curve clearly shows theChappuis ozone band, which disappears completely (red curve) after removing the retrievedozone absorption. The Chappuis band, however, has a weak temperature dependence increasingtowards both wings of the band. It has a broad-band envelope which can be approximated bya low-degree polynomial [Burkholder and Talukdar 1994] and residual small-band (differential)absorptions in the blue to green spectral range (440 – 540 nm). The latter are used to removethe broad-band temperature dependence of the ozone absorption which otherwise would havemimicked a fake broad-band Mie extinction. Here the temperature dependence of the residualsmall-band ozone absorptions are used by employing two ozone reference spectra at -40C and-60C (see chapter 4) to remove that fake broad-band Mie extinction.

The NO2 absorption cross section is increasingly temperature dependent when going from theblue to the UV spectral range. The most troublesome spectral region is excluded by the choiceof the analyzed wavelength interval (440 – 615 nm). The remaining temperature dependenceis accounted for by including in the DOAS fit two NO2 absorption spectra recorded in thelaboratory at -45 and -70. The small contribution of the NO2 absorption on the measuredresidual extinction is demonstrated in panel (e) of Figure 5.2, from which follows that theaccuracy of the aerosol extinction is not as much affected by the accuracy of the NO2 retrieval.

The collision pair absorptions of O4 has relatively broad-band absorption features. Theirabsorptions are proportional to the square of the oxygen concentration. Osterkamp [1997] andPfeilsticker et al. [2001] investigated the atmospheric O4 absorption using the present instru-ment. In the laboratory the O4 absorption was measured at 55 bar and room temperature e. g.by Greenblatt et al. [1990]. Due to a small pressure dependence, the UV-visible absorption bandof Greenblatt et al. are somewhat broader than those observed in the atmosphere as noted byHerman et al. [1996]. Here we use the cross section of Greenblatt et al. with a correction givenby Herman et al.. Within the visible spectral range there are three significant absorption bandsof O4 (see Figure 4.7). Panel (f) of Figure 5.2 shows the residual absorption before and afterremoving the three contributing O4 absorption bands.

After having accounted for all the molecular gas absorptions, the remaining extinction isonly due to Rayleigh and Mie scattering. Its smooth wavelength dependence can be clearly seenin the black curve of Figure 5.2 (g).

5.3 Removal of Rayleigh Extinction

Both Rayleigh and Mie extinction show broad-band features, which precludes to separate themin one step. Here, the Rayleigh scattering is accounted for from the known Rayleigh scatteringcross section 3.2, known pressure and temperature profiles, and calculated line of sights in aspherical refractive atmosphere.

For atmospheric pressures and temperatures in the present study three sources are used.First, prior and post to the LPMA/DOAS balloon launch, regularly meteorological radiosondesare launched (Vaisala Series RS80) which deliver high resolution temperature and pressure pro-files with a precision given Table 5.1. Ambient pressures and temperatures are also measuredaboard the balloon gondola. In post-flight data analysis, both sets are compared. In most casesit is found that they agree within the precision given in Table 5.1. In all other cases, they arecombined for a homogenized T and p data set (Figures 5.4 and 5.5). For altitudes around and

5.3. REMOVAL OF RAYLEIGH EXTINCTION 63

1E-4 1E-3 0.01 0.1 1 10 100 10000

20

40

60

80

100

120

León 1996 León 1998 Gap 1997 Gap 1999 Kiruna 1997 Kiruna 1998 Kiruna 1999 Kiruna 2000

A

ltit

ud

e [k

m]

Pressure [mbar]

Figure 5.4: Pressure profiles used for the calculation of the Rayleigh extinction. For details see text.

above the balloon float altitude T and p data from the US Standard Atmosphere [1976] for theappropriate latitude and season are taken with a connection between the measured and theprescribed data made at the largest altitude where the respective data agree. Note that eventualmisalignments of the measured and prescribed profiles are in any case not important as Rayleighscattering above the balloon float altitude is small.

Pressure measurement precision (1060 to 3 hPa) 0.5 hPaTemperature measurement precision 0.2 KRelative Humidity (RH) measurement precision 3% RHWind vector measurement precision using Loran-C 0.7 m

s

Table 5.1: Vaisala Series RS80 Loran-C Radiosonde

The line of sight taken from the balloon to the sun for observation (or spectrum) i is per-formed by the raytracing program damf [chapter 4.3.3 and Schulte 1996]. It considers a refrac-tive and spherical atmosphere which is sliced into atmospheric altitude segments (onion peels).


160 180 200 220 240 260 280 300 320 340 360 3800

20

40

60

80

100

120

León 1996 León 1998 Gap 1997 Gap 1999 Kiruna 1997 Kiruna 1998 Kiruna 1999 Kiruna 2000

Alt

itu

de

[km

]

Temperature [K]

Figure 5.5: Temperature profiles used for the calculation of the Rayleigh extinction. For details seetext.

The contributions of the various altitude segments j to the total slant air column density (de-noted as AMFij) are calculated straightforwardly by matrix multiplication of the air mass factormatrix AMF

SCDAiri =

N∑

j=1

AMFij · V CDAirj , (5.3)

whereV CDAir

j = nj ·∆hj =pjk Tj

·∆hj (5.4)

is the vertical air column density of the altitude segment j, nj is the molecular density, pj thepressure, Tj the temperature, ∆hj the width of the altitude segment j, and k is the Boltzmannconstant. For an individual altitude segment being about 50 meters thick, pressures and temper-atures are assumed to be constant. Since the sunlight traverses the longest distances in altitudesegments located around the tangent altitude, a very fine altitude segmentation is required there.

The accuracy of the air SCD is obtained from

∆SCDAiri =

N∑

j=1

AMFij ·∆V CDAirj , (5.5)

5.4. ZERO AIR MASS CORRECTION (LANGLEY PLOT) 65

82 84 86 88 90 92 94 96

1E25

1E26

SC

DA

ir [c

m-2]

SZA [deg]

10

15

20

25

30 (a)

Alt

itu

de

[km

]

82 84 86 88 90 92 94 96

(b)

SZA [deg]

1

10

SC

DA

ir/S

CD

Air [%

]

Figure 5.6: The total slant air column densities for the Kiruna balloon flight on February 18, 2000,are printed in black. Graph (a) additionally shows the altitude of the gondola. In Graph (b) therelative errors of the SCD are printed in red. Obviously, large SCDs are necessary for accurateaerosol extinction measurements.

where∆V CDAir

j

V CDAirj

=∆p

pj+

∆T

Tj(5.6)

∆T and ∆p are the temperature and pressure precisions given in Table 5.1.

The calculated air-SCDs of the balloon flight on February 18, 2000 at Kiruna are shown inFigure 5.6 (a) and the corresponding relative errors in Figure 5.6 (b). During the solar occultationmuch larger SCDs occurred than during the ascent. The relative error of a slant column with25 km tangent altitude (92.2 SZA) does not exceed 5%. At a tangent altitude of 20 km (93.2

SZA), the relative error decreases to about 2%. Thus, the most accurate aerosol extinction dataare expected for altitudes between the tropopause and 20 km. For an error discussion of theaerosol extinction see section 5.5.

5.4 Zero Air Mass Correction (Langley Plot)

The Fraunhofer reference spectrum to which all measurements are related to is recorded forminimal atmospheric extinction, i. e. around balloon float altitude. It thus contains a small ex-tinction due to atmospheric scattering and absorption overhead. A zero air mass extinction (i.e. the extinction contained in the Fraunhofer reference spectrum) is thus obtained by extrapo-lating the measured extinction as a function of air mass to zero air mass. To obtain the totalobserved aerosol extinction, this contribution must be added to the aerosol extinction retrievedfor each observation. Assuming for the moment that the aerosol extinction (αmie (λ)) per unitair number density nAir is constant afloat (indeed above the Junge layer it is very small andtherefore the assumption appears to be justified) the measured aerosol extinction along the lineof sight (βMie (λ)) is given by

βMie (λ) =αMie (λ)

nAir· SCDAir − β0 (λ) , (5.7)


0 2 4 6 8 10-0.015

-0.010

-0.005

0.000

0.005

0.010

0.015

0.020

(-0.00607 ± 0.00307) + 1.4234·10-27cm2 · SCD

AirA

ero

sol O

pti

cal D

epth

SCDAir

[1024cm-2]

Figure 5.7: Langley plot from which the aerosol extinction at λ = 603nm is determined. The dataare from balloon float observations (SZA < 90) taken during the Kiruna balloon flight onFebruary 18, 2000. The red line shows a linear fit to the data according to Equation 5.7 with anextrapolation to zero air mass (β0 (λ) = (6.1± 3.1) · 10−3).

where the zero air mass extinction β0 (λ) and SCDAir is the slant air column density.

Figure 5.7 (Langley plot) shows the data for the Kiruna balloon flight on February 18, 2000.The data are for observations at balloon float altitude and SZAs smaller than 90. A fit to thedata according to Equation 5.7 is indicated by the red line. The departure of the data from thered line demonstrates the validity of the above mentioned assumption. The extrapolation to zeroair mass (SCDAir = 0) then yields β0 (λ) (= (6.1± 3.1) · 10−3).

5.5 Error Analysis

First it is worthwhile to consider whether observational effects such as the pointing to the sunmay impact the inferred Mie extinction. This may come through effects arising from the largefield of view of the telescopes (16 opening angle) or from changes of the sun’s brightness centerduring solar occultation. Evidently, the telescopes also receive atmospheric scattered light thatadds to the direct sunlight. Accordingly, that effect would systematically decrease the inferredaerosol extinctions and trace gas absorptions, the latter however is not observed. Together withthe much larger (> 104) direct sun than received scattered light intensity, this effect is unlikelyto be important.

5.5. ERROR ANALYSIS 67

Apparent changes of the sun’s brightness center during solar occultation are known to beimportant in the retrieval of minor trace gases [e. g. IO, OIO, see Bosch 2001]. In the calculationsof the tangent points this effect, however, is considered in the raytracing model by adding inten-sity weighted contributions of sunlight emitted from four equally spaced slices of the solar disc.During the measurements, the suntracker automatically tracks towards the apparent brightnesscenter of the sun. This first order correction apparently keeps the calculated below the actualbrightness center. Consequently, the calculated Rayleigh scattering is somewhat overestimatedwhich may lead to an underestimation of the inferred aerosol extinction.

Errors due to spectrograph stray light, Rayleigh scattering, and trace gas absorptions (O3,NO2, and O4) are calculated according to their wavelength dependence given in Equation 5.1(for the results see Figure 5.9).

Assuming the spectrograph stray light is linear in wavelength ((c0 − c1 (λ− λ0)) ·Imean) thenwhen added to the radiance received by the telescope I (λ) leads to the equation

I (λ) + (c0 − c1 (λ− λ0)) · Imean = I (λ) . (5.8)

where I (λ) measured intensity.Assuming Gaussian error propagation, the radiance error ∆I (λ) is then obtained from

∆I (λ) =

√

(

∆I (λ))2

+(

∆c0 · Imean

)2+(

(λ− λ0)∆c1 · Imean

)2, (5.9)

where ∆c0 and ∆c1 are the errors in the fitting parameters c0 and c1 given by the DOAS fit-ting procedure, that requires to minimize residual Fraunhofer absorptions (see chapter 4.2). Theerror of the measured intensity ∆I (λ) results from the various contributions to the instrumen-tal noise (see chapter 4.3.2). For the very large intensities encountered during direct sunlightmeasurements, Ferlemann et al. [2000] has shown that the instrumental noise is of the order ofabout 10−4, i. e. around a factor of 10 smaller than the spectrograph stray light offset errors, inthus not important here.

The total optical density αtotal along the line of sight is obtained by taking the logarithm ofthe ratio of I and the Fraunhofer reference spectrum I0

1,

αtotal = − ln

(

I

I0

)

. (5.10)

and accordingly the error

∆αtotal =

√

√

√

√

(

∆I

I

)2

+

(

∆I0I0

)2

. (5.11)

In general, due to the larger intensity of the Fraunhofer reference spectrum than of any othermeasured spectrum the instrumental noise of the former (∆I0) is smaller than than of the latter(∆I), the first term of the equation’s 5.11 right hand side dominates over the second term.

Apparently, the main error source in the aerosol extinction retrieval algorithm is the uncer-tainty due to Rayleigh scattering (Figure 5.2), being in particular large at short wavelengths.Primarily the error is given by uncertainties in atmospheric temperatures and pressures ratherthan by uncertainties in the Rayleigh scattering cross section. The air SCD and its error arecalculated using equation 5.5 and 5.6.

1To simplify matters, in the following equations the argument (λ) is omitted but it applies for all wavelengths.


A sensitivity test exchanging in-situ measured temperatures and pressures with data fromNCEP (National Center for Environmental Prediction) analysis only shows a weak dependenceof the inferred Mie extinction with the used data (for details see Figure 5.8).

1E-5 1E-4 1E-35

10

15

20

25

30

35

Alt

itu

de

[km

]

Aerosol Extinction Coeffificient [km-1]

Figure 5.8: Change in the retrievedaerosol extinction coefficients usingtwo different sets or air data forthe Kiruna 2000 measurements. Thered/blue curves show the DOAS re-sults using the pressure and tempera-ture data as described in chapter 5.3and from NCEP, respectively. The blackcurve shows the aerosol extinction coeffi-cient retrieved by POAM III using NCEP

(for details see chapter 6).

The trace gas absorptions and the involved errors are obtained from the DOAS fitting routine[see chapter 4 and Ferlemann et al. 1998]. The aerosol extinction αMie is obtained by subtractingthe trace gas absorptions and the Rayleigh extinction from the total extinction αtotal,

αMie = αtotal −∑

i

σi · SCDi − σRayleigh · SCDAir. (5.12)

The error propagation is thus given by

∆αMie =

√

(∆αtotal)2 +

∑

i

(σi ·∆SCDi)2 + (σRayleigh ·∆SCDAir)

2. (5.13)

Figure 5.9 outlines the error contributions of each term to the error of the total extinction for thesame spectrum as shown in Figure 5.1 and 5.2. Evidently uncertainties in Rayleigh scattering (orcalculated slant air column density) are dominating the total error in the retrieval. Unfortunately,these uncertainties cannot be reduced significantly as due to given reasons both the measurementaccuracy of the radiosondes and expected horizontal gradients in the atmospheric pressure andtemperature fields are matter of facts.

The absolute errors in the spectrograph stray light parameters (∆c0 and ∆c1) eventuallyreach 1

6 of the total extinction error at the long wavelength end, but it hardly exceeds 1%.Even smaller are errors due to uncertainties the inferred trace gas absorptions with the largestindividual contribution coming from the Chappuis ozone absorption. At most it is of the sameorder as the error contribution of the spectrograph stray light. Errors due to uncertainties inthe inferred O4 and NO2 absorption are at least a factor of 20 smaller than the total extinctionerror and therefore in practice negligible.

Overall the most accurate Mie extinction measurements in slant air columns can be performedat the longest wavelengths as there Rayleigh scattering, and hence its errors, contribute least tothe total extinction. Typically, uncertainties in Rayleigh scattering introduce an relative errorof 1.5%/3.4% into the inferred Mie extinction at the long/short wavelength end, respectively.The errors in the vertical aerosol extinction profile are somewhat larger than for the slant air

5.5. ERROR ANALYSIS 69

450 500 550 600

1E-4

1E-3

0.01

NO2

O4

O3

Rayleigh

SSL

R

elat

ive

Err

or

(''DD

Mie/DD

Mie)

Wavelength [nm]

Figure 5.9: Relative errors of the total extinction originating from the various contributions. Thedominating error source for all wavelengths is the Rayleigh extinction.

columns, since in the inversion algorithm the errors are multiplied by the reciprocal of theAMF matrix’ respective eigenvalues for the considered altitude segment [see chapter 4.3.4 andRodgers 1976]. Particularly for altitudes above 25 km where the air mass factors are gettingsmall, the inferred aerosol extinction profile is becoming increasingly uncertain. For tangentheights located in the troposphere, due to dramatically increasing Mie and Rayleigh scatteringthe light intensity is small preventing an accurate measurement of the aerosol extinction. Hence,a reliable measurement of aerosol extinction profiles is limited to an altitude range between thetropopause and no more than about 25 km.

Chapter 6

Results and Comparison withSatellite Data

From 1996 to date, nine flights of the LPMA/DOAS balloon payload have been performed (Ta-ble 6.1). Funding came through the European Union and the German Ministry for Educationand Research (BMBF). The scientific objectives are to study in the lower stratosphere the sea-sonal cycle of ozone relevant trace gases. The projects CHORUS and CHELOSBA (Chemistryof the Lower Stratosphere with Balloons) aim to tackle the chemistry of the odd nitrogen (NOy)and halogen components. The projects ODIN and ADEOS/ILAS focussed on the validation ofthe Japanese ILAS [see chapter 3.1] instrument and of the Canadian/Swedish/Finnish ODINinstrument on OSIRIS. THESEO and THESEO2000 examined the process that lead to strato-spheric ozone loss in the Arctic and at mid-latitudes. The latter was closely coordinated withthe American NASA-led SOLVE campaign. Most recently, a balloon flight was conducted fromKiruna on August 21/22, 2001 of which the data retrieval is in progress. In near future, a setof LPMA/DOAS balloon payloads will be launched primarily addressing the validation of theSCIAMACHY instrument on ENVISAT.

No. Location Date Campaign

1 Leon 42.6N 5.7W Nov. 23, 1996 CHORUS

2 Kiruna 67.9N 21.1E Feb. 14, 1997 ADEOS

3 Gap 44.0N 6.1E Jun. 20, 1997 CHELOSBA

4 Leon 42.6N 5.7W Mar. 19, 1998 CHELOSBA

5 Kiruna 67.9N 21.1E Aug. 19/20, 1998 SABINE/CHELOSBA

6 Kiruna 67.9N 21.1E Feb. 10, 1999 THESEO

7 Gap 44.0N 6.1E Jun. 25, 1999 THESEO

8 Kiruna 67.9N 21.1E Feb. 18, 2000 THESEO2000/SOLVE

9 Kiruna 67.9N 21.1E Aug. 21/22, 2001 ODIN

Table 6.1: Summary of the LPMA/DOAS balloon flights. The used launch stations for the big balloonsoperated by the French balloon launching company CNES (Centre National d’Etudes Spatiales)are in Leon (Spain), Gap (France), and Kiruna (Sweden)

The geophysical conditions encountered during the various balloon flights evidently differedwhen going from high to mid-latitudes and to different seasons. The given set of balloon flights is

71

72 CHAPTER 6. RESULTS AND COMPARISON WITH SATELLITE DATA

well suited to investigate the variety of different geophysical conditions and to validate satelliteborne instruments. In contrast, it does not allow a straightforward comparison between theresults obtained during the different balloon flights and in particular the investigation of temporalevolutions of the measured parameters is not possible.

Two balloon flights (No. 1 and 4) were undertaken from Leon (northern Spain) in spring andautumn, when the stratospheric air is transported into easterly directions. Both flights were socalled balloon ascent measurements continuing into complete sunset.

Two balloon flights (No. 3 and 7) took place from Gap (southern France) in June 1997 andJune 1999. During summer stratospheric winds are easterlies at mid-latitudes. The local weatherconditions in Gap allow to launch big stratospheric balloons during the night only, thus sunriseand balloon descent measurements could be performed in the morning stratosphere.

The balloon flights 2, 6, and 8 were performed from the European Space Range (Esrange)near Kiruna in northern Sweden in winter. On February 14, 1997 unfortunately, the edge ofthe polar vortex was located closely over Kiruna. Therefore, the line of sights mostly crossedthe polar vortex’ edge region which separates chlorine activated air masses low in NOx from thevortex’ interior from mid-latitude air high in NOx and low in active chlorine [Fitzenberger 2000].The measurements in February 1999 and 2000 took place completely inside the polar vortex andthus provide an excellent data set to study the stratospheric chemistry leading to the ozonedepletion in polar spring.

The balloon flights 5 and 9 were also conducted from Esrange, however, at turnover of thestratospheric circulation in late summer. Therefore, the stratospheric winds were very weakallowing long balloon flights (8 hours) without leaving the range of telecommunication. Accord-ingly, the measurements extended from the balloon’s ascent in late afternoon to sunset, sunrise,and the balloon’s descent in the early morning.

More detailed information including maps of the balloon flights’ trajectories and the positionsof the tangent points for different solar zenith angles can be found in the Ph. D. thesis ofFitzenberger [2000] (all flights), Ferlemann [1998], and Harder [1999] (only the first three flights).

6.1 Vertical Aerosol Extinction Profiles

An aerosol extinction profile−−−→V CD is obtained by multiplication of the

−−−→SCD the aerosol extinc-

tion vector with the inverted air mass factor (or observation) matrix AMF−−−→V CD = AMF−1 ·

−−−→SCD. (6.1)

In our case, no vertical or slant column densities of trace gases are calculated, but Mie extinctionper unit length. The nomenclature

−−−→V CD and

−−−→SCD are retained, because the same inversion

algorithm is implemented as for the trace gas retrieval. Since the aerosol extinction coefficientis wavelength dependent, the extinction at a particular wavelength is taken for the inversion.

The AMF matrix is calculated by a raytracing program (see chapter 4.3.3), requiring asinput parameters a vertical pressure and temperature profile of the atmosphere, partitionedsegments with a constant concentration (or in this case, aerosol extinction). The larger thesegments, the more accurate is the resulting mean vertical aerosol extinction. The observationgeometry of the DOAS spectrograph permits a vertical resolution of around 1 km at large SZAs.Thus, the width of the segments is chosen to 1 km. At high altitudes, the air masses along theline of sight are becoming smaller, and the extinction measurement less accurate. Hence, forsome flights the resolution is reduced at high altitudes.

6.1. VERTICAL AEROSOL EXTINCTION PROFILES 73

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Aerosol Extinction Coefficient [km-1]

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Figure 6.1: Sunset aerosol extinction observa-tions at Kiruna on February 14, 1997.

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Figure 6.2: Same as 6.1 but for Kiruna onFebruary 10, 1999.

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Figure 6.3: Same as 6.1 but for Kiruna onFebruary 18, 2000.

603 nm 525 nm 452 nm 442 nm

The most complete global data sets of stratospheric aerosol extinction are provided by thesatellite borne instruments SAGE II, and POAM II and III. In the visible spectral range, POAMinfers vertical aerosol extinction profiles at 442 and 603 nm, SAGE at 452 and 525 nm1 (for detailssee section 6.2). In the following, the aerosol extinction coefficients measured by DOAS at thesefour wavelengths are considered, with a particular emphasis put on the 603 nm channel, whoseretrieval is most accurate due to the smallest Rayleigh scattering cross section there.

Figures 6.1 to 6.9 show vertical profiles of the aerosol extinction coefficient observed duringdifferent solar occultations at 442, 452, 525, and 603 nm, respectively. The aerosol extinctionprofiles of the sunset polar winter flights in 1997, 1999, and 2000 for Kiruna (Figures 6.1,6.2, and 6.3) corresponds to expectations for a volcanically quiet non PSC-loaded polar winterstratosphere. Likewise, the sunset observations at Leon 1998 (Figure 6.5) and Kiruna 1998

1In addition, SAGE II observes aerosol extinction at 386, and 1020 nm, POAM III at 353, 779, 922, and 1018nm (see section 6.2)


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Figure 6.4: Sunset aerosol extinction observa-tions at Leon on November 23, 1996.

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Figure 6.5: Sunset aerosol extinction observa-tions at Leon on March 19, 1998.

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Figure 6.6: Sunrise aerosol extinction obser-vations at Gap on June 20, 1997.

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Figure 6.7: Sunrise aerosol extinction obser-vations at Gap on June 25, 1999.

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Figure 6.8: Sunset aerosol extinction observa-tions at Kiruna on August 19, 1998.

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Figure 6.9: Sunrise aerosol extinction obser-vations at Kiruna on August 20, 1998.

6.1. VERTICAL AEROSOL EXTINCTION PROFILES 75

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Figure 6.10: Comparison of 603 nm aerosolextinction coefficient observed for sunsetflights at Kiruna in wintertime.

Figure 6.11: Comparison of 603 nm aerosolextinction coefficient observed for sunsetand sunrise during the Kiruna flight insummer 1998.

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León 1996 León 1998 Gap 1997 Gap 1999

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Balloon flight Tropopause

Leon 1996 12.6 kmKiruna 1997 9.1 kmGap 1997 11.8 kmLeon 1998 12.3 kmKiruna 1998 10.2 kmKiruna 1999 8.0 kmGap 1999 12.1 km

Kiruna 2000 9.0 km

Figure 6.12: Comparison of 603 nm aerosolextinction coefficient observed for themid-latitudinal flights at Leon and Gap.

Table 6.2: Local tropopause altitudes duringthe DOAS balloon flights.

(Figure 6.8) are as expected for a volcanically quiet mid-latitude and high-latitude stratospherein fall and summer. Note that the relative errors for the latter flights are almost twice as largeas for the Kiruna 1999 and 2000 flights, because there the float altitude (where the ambientpressure is small and thus less accurate) was much larger and less solar occultation spectra wererecorded than elsewhere.

Unfortunately, the sunrise solar occultation measurements (at Gap in 1997 [Figure 6.6] andin 1999 [Figure 6.7], and at Kiruna in 1998 [Figure 6.9]) lack in observations for low tangentheights. This is because the suntracker Heliostat [chapter 4.4 and Hawat et al. 1995] needs sometime to lock in to the sun as it rises above the horizon2. Accordingly, during the sunrise flights at

2In general, the gyroscope adjusts the gondola approximately into the calculated direction of the Sun before itis visible, but under adverse circumstances, Heliostat may have to scan the whole viewing angle. Once the Sun isfound, only small corrections are necessary, since in the stratosphere the wind conditions are very constant, and


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Tro

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[km

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Figure 6.13: Summary of the results of all DOAS balloon flights. The vertical profile of the aerosolextinction coefficient at 603 nm is plotted relatively to the tropopause (≡ 0 km)

Gap in 1997, and at Kiruna in 1998, tangent heights upwards from 11 km could be monitored.During Gap flight in 1999, the suntracker could only lock into the sun at a solar zenith angleof 93.2, corresponding to a tangent altitude of 28.5 km. Sunset observations are not affectedby this lock-in problem, because the suntracker already adjusts to the sun long before solaroccultation, i. e. during balloon ascent. Due to logistic reasons, the sunset flight from Leon in1996 (Figure 6.4) had to be terminated at a solar zenith angle of 92.9 (corresponding to atangent altitude of 22.5 km) because it was heading into dangerous territory.

Figure 6.10 summarizes the results of the Kiruna balloon flights during in polar winter. The1999 and 2000 measurements took place completely inside the polar vortex, when the tropopausealtitudes were lower than for flights conducted outside the vortex. During the 1997 flight, theballoon’s trajectory was close to the edge of the polar vortex whereas for the other two flightsthe trajectory was located well within the vortex. Accordingly, for the 1997 flight air massesfrom outside the vortex were probed showing sightly elevated aerosol extinctions than for the1999 and 2000 vortex flights.

the gondola is not exposed to sudden torques.

6.2. THE SAGE AND POAM AEROSOL RETRIEVAL ALGORITHM 77

For the 1996 to 1999 period, the aerosol extinction profiles show good agreement for mid-latitudes (Figure 6.12). This behavior has three reasons: (a) since 1996 the stratospheric aerosolburden has reached background values (see Figure 2.5), (b) Gap and Leon are located almostat the same latitude precluding possible latitudinal effects, and (c) the tropopause heights weresimilar for all flights preventing dynamic changes in the aerosol layer’s shape.

In Figure 6.13 the aerosol extinction profiles from all DOAS balloon flights are shown rel-atively to the actual tropopause location. This removes the vertical displacement due to at-mospheric dynamics. The good overlap demonstrates the predominant effect of atmosphericdynamics on the profile shape for altitudes some 5 km above the tropopause.

In the following the Mie extinction measurements of the present studies will be comparedin more detail with SAGE II and POAM III measurements. Before doing so, the SAGE andPOAM aerosol retrieval algorithm will be briefly discussed.

6.2 The SAGE and POAM Aerosol Retrieval Algorithm

The major difference between balloon-borne DOAS observations and satellite measurementsis that the SAGE and POAM instruments record slant columns of atmospheric extinctionsin several narrow-banded channels ranging from about 350 to 1060 nm, whereas the DOASspectrograph observes a continuous spectrum in the wavelength range 417 – 670 nm. As shownin the previous chapters, the present aerosol retrieval method separates the contribution of theatmospheric absorption to the total extinction from Rayleigh and Mie scattering by applying theDOAS method for the trace gas measurements first (see chapters 4.3 and 5.2). In a second step theresulting extinction (being due to Rayleigh and Mie scattering only) is then interpreted as beingdue to a calculated Rayleigh scattering and the wanted Mie scattering. The SAGE and POAMaerosol retrieval algorithms rely on a simultaneous solution of a linear equation system [e. g.Chu et al. 1989 (SAGE II), Lumpe et al. 1997 (POAM II), and Lucke et al. 1999 (POAM III)].The linear equation system is built on summing up for each wavelength channel all processes(gas absorption, Rayleigh and Mie extinction) contributing to the total extinction with theircorresponding relative wavelength dependencies being either given by the relative absorption

Channel Center Width Particular sensitive to

1 352.3 nm 4.4 nm O3, NO2, Aerosol2 441.6 nm 2.0 nm NO2 (off peak), O3, Aerosol3 448.1 nm 2.1 nm NO2 (on peak), O34 601.4 nm 14.3 nm O3, Aerosol5 761.2 nm 2.2 nm O2 (on peak), O36 781.0 nm 16.7 nm O2 (off peak), O3, Aerosol7 921.0 nm 2.1 nm H2O (off peak), O3, Aerosol8 936.4 nm 2.3 nm H2O (on peak), O39 1060.3 nm 11.1 nm Aerosol

Table 6.3: Compendium of the nine POAM II channels, the respective center wavelengths and widths,and the processes contributing to the total atmospheric extinction. The specific channels with”Aerosol” are devoted for aerosol extinction measurements. Instrumental data are adopted fromLumpe et al. [1997].


Channel Center Width Particular sensitive to

1 353.4 nm 9.7 nm O3, NO2, Aerosol2 439.6 nm 2.1 nm NO2 (on peak), O33 442.2 nm 2.1 nm NO2 (off peak), O3, Aerosol4 603.0 nm 17.7 nm O3, Aerosol5 761.3 nm 2.3 nm O2 (on peak), O36 779.0 nm 10.2 nm O2 (off peak), O3, Aerosol7 922.4 nm 2.6 nm H2O (off peak), O3, Aerosol8 935.9 nm 2.9 nm H2O (on peak), O39 1018.0 nm 11.6 nm Aerosol

Table 6.4: Same as 6.3 but for POAM III. Instrumental details are adopted from Lucke et al. [1999]

or scattering cross sections. Obviously, this does not pose a big problem for the gas absorptioncross sections (except for a weak temperature dependence) and for the Rayleigh scattering crosssection, but for the a priori assumed wavelength dependence of the Mie scattering cross section.Clearly, for the latter a priori assumptions have to be made about the size distribution and thereal and complex refractive index when the wavelength dependence of the Mie scattering crosssection is calculated using Mie theory. In that respect, SAGE II, POAM II and III retrievalsare similar except for some small differences in the center wavelengths and widths of the chosenchannels.

Tables 6.3 (POAM II), 6.4 (POAM III), and 6.5 (SAGE II) summarize the measurementchannels of the three satellite instruments. For each channel, the different trace gases absorbingin the respective spectral range are quoted. Additionally to the quotation, Rayleigh scatteringand aerosol extinction contribute to the observed total extinction. Like the DOAS retrieval, theSAGE and POAM algorithms use atmospheric pressure and temperature data to calculate theRayleigh extinction for the raytraced line of sight. For that purpose, a global atmospheric dataset is provided by the National Center for Environmental Prediction (NCEP), and formerlyby the National Meteorological Center (NMC). Tropospheric data (1000 – 100 hPa) are fromthe NCEP Global Data Acquisition System (GDAS), while stratospheric data (70 – 0.4 hPa)are obtained from the NCEP analysis (from Climate Analysis Center CAC). The NCEP CACdata is described in Newman et al. [1989]. The scheme uses combined and assimilated satelliteretrievals and radiosonde data for altitudes up to the 10 hPa, and satellite data only for altitudesbetween 10 and 0.4 hPa [Finger et al. 1993].

Channel Center Particular sensitive to

1 386 NO2, Aerosol2 448 NO2 (on peak), O33 452 NO2 (off peak), O3, Aerosol4 525 O3, Aerosol5 600 O36 935 H2O, O37 1020 Aerosol

Table 6.5: Same as 6.3 but for SAGE II.

6.2. THE SAGE AND POAM AEROSOL RETRIEVAL ALGORITHM 79

The pressure and temperature data used for the present aerosol extinction analysis is alsofrom radiosonde measurements, in addition to the data monitored by aboard deployed sensors.My analysis has shown that small discrepancies between the radiosonde and NCEP pressure andtemperature data only weakly propagate into the resulting Mie extinction (for a sensitivity testwith the LPMA/DOAS data see Figure 5.8).

For the satellite retrievals a calculated Rayleigh extinction is subtracted from the measuredwavelength dependent total extinction. The remaining wavelength dependent extinction is thendue to varying strong contributions from trace gas absorptions (listed in Tables 6.3, 6.4, and 6.5),and Mie scattering. For example, the two visible channels 2 and 3 are in particular sensitive toNO2 absorption. One channel is placed on a local maximum of the NO2 absorption cross section,the other one at a local minimum. For these narrowly located wavelength, the ozone absorptioncross section and the aerosol extinction can assumed to be constant, and thus they deliver adifferential signal of the atmospheric NO2 absorption. In a similar manner, the O2 (channels 5and 6) and H2O absorption (channels 7 and 8) are inferred from the POAM II and III data.

Once the contributions from the NO2, O2, and H2O absorption are known, they are sub-tracted from the remaining extinction. As a result, channel 9 (POAM II and III) or channel7 (SAGE II) is particularly sensitive to the aerosol extinction, whereas to the resulting ex-tinction of all other channels is due to ozone absorption and Mie scattering. If the Mie ex-tinction in each channel were assumed to be independent, together with the ozone absorptionone would obtain the number of channels plus one independent equations. This is an ill-posedproblem. Therefore, a Mie theory based wavelength dependence of the Mie extinction is as-sumed (usually containing three independent parameters, see Equation 6.2) and the system ofequations is solved. Frequently, the following aerosol extinction wavelength dependence is taken[Brogniez and Lenoble 1988]:

ln (β (λ)) = ln (β (λ0))− α ln

(

λ

λ0

)

− b

(

ln

(

λ

λ0

))2

. (6.2)

Including this formula into the system of equations, the four unknowns are β (λ0), α, b, and theozone SCD. This method, however, is an idealized approximation.

First, absorption cross sections are known to vary with temperature (primarily O3 andNO2 in the visible spectral region). Therefore, measured line of sight absorptions in the realatmosphere are always a mixture of contributions from absorptions at different temperatures.This problem introduces new parameters into the system of linear equations rendering themill-conditioned. Avoiding ill-conditioning the approximation described above leads to potentiallysignificant errors in the inferred parameters. Most important is the error propagation of anassumed ozone absorption cross section constant in temperature into all other parameters.

On purpose, the POAM O2 channels were designed to derive the total slant air column andthe Rayleigh extinction. It was found, however, that channel 5 is saturated for tangent pointslocated above about 30 km where the instrument is exposed to nearly unattenuated sunlight[Glaccum et al. 1996]. Consequently, this channel has not been used for the profile retrieval.Alternatively, the O2 absorption for POAM channel 6 is calculated from air data (NCEP).

In the SAGE II measurement, H2O absorption is monitored in channel 6 only, and theinformation from channels 2 and 3 is combined to get NO2. Accordingly, 5 independent equations(for the 386, 452, 525, 600, and 1020 nm extinction) remain only to infer the aerosol extinction(3 parameters) and ozone absorption (1 parameter). In the POAM retrievals, however, afterhaving retrieved NO2, O2, and H2O, there remain 6 independent equations (for the 352.3, 441.6,


601.4, 781.0, 921.0, and 1060.3 nm POAM II extinctions, and for the 353.4, 442.2, 603.0, 779.0,922.4, and 1018.0 nm POAM III extinctions) to infer the 4 parameters.

A particular strength of the satellite instruments is that unlike in our measurements 5.5spectrograph stray light hardly plays a role as they use narrow band-pass filters known tosuppress efficiently the spectrograph stray light.

In the SAGE and POAM retrieval algorithm the ozone SCD and the aerosol extinction,however, are strongly correlated. Since the Chappuis ozone absorption is assumed to be in-dependent of temperature and because of the cross correlation errors are potentially intro-duced into the retrieval of the aerosol extinction. This is according to our opinion a par-ticular concern. On the other hand, ozone profiles measured by SAGE and the POAMshave been extensively validated and reasonable good agreement was found frequently [e.g. Attmannspacher et al. 1989; Cunnold et al. 1989; Yue et al. 1995; Lu et al. 1997 (SAGE) orRusch et al. 1997; Hansen et al. 1997; Sugita et al. 2000; Lumpe et al. 2001 (POAM)].

In concluding, all three instruments provide a very powerful tool to monitor globally in thestratosphere the near IR aerosol extinction (POAM channel 9, SAGE channel 7), NO2, O3, andto a less degree H2O and O2 profiles. A weakness is, however, that the wavelength dependence ofthe aerosol extinction has to be assumed while it possibly cross-interferes with the temperaturedependence of the ozone absorption.

6.3 Comparison of the DOAS with Satellite Aerosol ExtinctionData

As described in chapter 3.1, only few validated extinction observations of stratospheric aerosolsare available. Therefore, comparison of the DOAS with satellite aerosol extinction data is desired.For that purpose, collocated satellite and DOAS observations are further investigated.

SAGE II observations cover low and mid-latitudes throughout the year with a longitudi-nal spacing of 24. Since solar occultation measurements are performed, extinction profiles arerecorded up to latitudes of 75 in summer. In midwinter the spatial coverage is limited to about50 (Figure 6.14). The POAMs take 14 sunrise observations in the northern hemisphere between55 and 71N per day (Figure 6.15) with a longitudinal spacing of about 25.7. For PSC-freeand volcanically quiet periods, the longitudinal but not the latitudinal variation of stratosphericaerosol is expected to be small and therefore a good validation requires a good match of thetangent points’ latitudes but to a less degree of their longitudes.

Table 6.6 provides an overview on the satellite measurements suitable for a comparison withthe aerosol extinction data recorded by the DOAS balloon spectrograph. For the winter 1997Kiruna flight, no satellite overpass is available as SAGE II was not measuring that far northand POAM II stopped to deliver data by end of 1996. For the Gap 1997 flight, SAGE II’smeasurements did not cover northern mid-latitudes between June 5 and June 30, 1997. Postflight data analysis showed for the summer 1998 Kiruna flight that the probed air masses didnot sufficiently overlap as the line of sights of the balloon instrument were directed into northerndirections while the POAM III’s northernmost measurements in August 1998 were still morethan five degrees south of the balloon’s trajectory.

The aerosol extinction coefficients from SAGE II are Version 6.0 taken from a CD-ROM Setdistributed by the NASA Langley Research Center. The recent update, Version 6.1 (includingaerosol surface areas, densities and effective radii, and an improved NO2 retrieval algorithm)

6.3. COMPARISON OF THE DOAS WITH SATELLITE AEROSOL EXTINCTION DATA81

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec-80

-60

-40

-20

0

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80

Sunset Sunrise

L

atit

ud

e [d

eg]

Month of the year 1999

Figure 6.14: Location of SAGE II’s tangentpoint latitudes for measurements throughoutthe year 1999 (data adopted from CD-ROMSet SAGE II Version 6.0 provided by theNASA Langley Atmospheric Sciences Data

Center, ASDC).

Figure 6.15: Calculated tangent point lati-tudes of the POAM III measurement loca-tions. Adopted from the POAM web page

(wvms.nrl.navy.mil/POAM).

was not available during the time of the present study. The quoted aerosol extinction coefficienterrors, and the applied temperature and pressure data have been kindly provided by LarryThomason from the SAGE II Team.

The POAM III aerosol data used in this study have been downloaded from the Internet pagesof the NASA Langley Atmospheric Sciences Data Center, ASDC (http://eosweb.larc.nasa.gov/).The data have been recently revised (but a recent automated communication received fromthe NASA ASDC indicates a suspect quality). Nevertheless, due to the lack of other aerosolextinction data from high latitudes the available POAM III data were used.

In Figures 6.16 to 6.20 a comparison of satellite and balloon-borne extinction observationsis plotted. Note that each satellite instrument provides aerosol extinction at two wavelengthchannels for the analysis range of the DOAS aerosol retrieval. Due to the increasing Rayleighscattering cross section, the error bars are larger at shorter wavelengths. Due to the smallest

Balloon Flight Location Satellite Event Date

Leon 11/23/96 42.6N 5.7W SAGE 42.6N 0.1E 11/26/96

Kiruna 02/14/97 67.9N 21.1E no overpasses available

Gap 06/20/97 44.0N 6.1E no overpasses within 10 days

Leon 03/19/98 42.6N 5.7W SAGE 40.6N 15.7W 03/21/98

Kiruna 08/19/98 67.9N 21.1E no overpasses within ±5

Kiruna 02/10/99 67.9N 21.1E POAM 67.1N 39.0E 02/10/99

Gap 06/25/99 44.0N 6.1E SAGE 45.2N 6.5E 06/24/99

Kiruna 02/10/00 67.9N 21.1E POAM 67.5N 14.3E 02/18/00

Table 6.6: Compendium of closest DOAS and satellite measurement coincidences. For the flights inGap 1997, and in Kiruna 1997 and 1998 no satellite observations are available for intercomparison.


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Figure 6.16: Comparison of Leon balloon flight on November 23, 1996 (in red) with SAGE II data(in black, see Table 6.6). Left panel: 452 nm, right panel: 525 nm

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Figure 6.17: Comparison of Leon balloon flight on March 19, 1998 (in red) with SAGE II data (inblack, see Table 6.6). Left panel: 452 nm, right panel: 525 nm

errors, a most robust comparison can be made for measurements ranging between 15 and 25km of altitude. For this altitude range, the aerosol extinction coefficients inferred in the presentstudy are significantly larger for all flights than those measured by the satellite instruments.However, for each flight the vertical profiles show a very similar shape throughout the wholealtitude range, but the absolute value of the aerosol extinction coefficient seems to be underes-timated by the satellite instruments by a factor of 2 with respect to the DOAS results. Even formeasurements in the mid-stratosphere where the errors are large enough to engulf the satelliteresults, the retrieved extinction coefficients are systematically larger by the same factor thanthe satellite measurements. The discrepancy is somewhat larger at high latitudes, where POAMIII data is used for the comparison only. In comparison with the SAGE II results the inferredDOAS extinction coefficients are also larger. In particular this is also true for the mid-latitude

6.3. COMPARISON OF THE DOAS WITH SATELLITE AEROSOL EXTINCTION DATA83

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Figure 6.18: Comparison of Kiruna balloon flight on February 10, 1999 (in red) with POAM III data(in black, see Table 6.6). Left panel: 442.2 nm, right panel: 603.0 nm

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Figure 6.19: Comparison of Gap balloon flight on June 25, 1999 (in red) with SAGE II data (in black,see Table 6.6). Left panel: 452 nm, right panel: 525 nm

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Figure 6.20: Comparison of Kiruna balloon flight on February 18, 2000 (in red) with POAM III data(in black, see Table 6.6). Left panel: 442.2 nm, right panel: 603.0 nm


observation in Leon 1998, even though the DOAS errors for that flight are larger than for thehigh latitude observations, as for the former flight less spectra were recorded than for the latterset of flights.

Due to the good reasons given above there is no naive explanation for the stateddiscrepancy, since the satellite and the balloon-borne instruments were tested extensivelyprior to launch, and all of them show reasonable good agreement of inferred trace gasprofiles with other observations [e. g. Cunnold et al. 1989 and 1991; Chiou et al. 1993;Larsen et al. 1993; Lucke et al. 1999; Lumpe et al. 2001; Randall ]. The satellite aerosol extinc-tion measurements were, however, compared with each others, with in-situ measured aerosolsize distribution based Mie calculations or by other optical methods (c. f. LIDAR) [e. g.Ackerman et al. 1989; Russell and McCormick 1989; Brogniez et al. 1996; Burton et al. 1999;Randall et al. 2000; Randall et al. 2001]. In that respect, our DOAS aerosol measurements areunvalidated and one may suspect weather there are systematic drawbacks in our retrieval.

From the discussion given above it is evident that, systematic effects due to instrumentalcharacteristics, the measurement geometry, or difficulties with the DOAS retrieval algorithm areunlikely to explain the factor of 2 discrepancies in either of the inferred aerosol extinctions. Thisis also supported by the similar profile shape for all observations.

At this point one may speculate, however, that the discrepancies may arise from problemswith the assumed wavelength dependence of the aerosol extinction in the satellite retrievals(Equations 3.41 and 6.2). Accordingly, in the next chapter we will inspect the wavelength de-pendence of the aerosol extinction in our retrieval.

Chapter 7

Inferred Mie Scattering WavelengthDependence

7.1 Observed Wavelength Dependencies

The primary information in the DOAS retrieval on the wavelength dependence of the aerosolextinction is given after subtracting from the measured total optical densities the trace gasabsorption and the extinction due to Rayleigh scattering (see Figure 5.2 trace h, black line). Itis noteworthy that this signal, however, is an apparent Mie scattering wavelength dependencetaken as an integral along the line of sight. A similar signal is taken by the satellites. In a firstapproximation this can be taken as a local Mie scattering wavelength dependence for the aerosolslocated around the tangent point as most of the extinction is occurring there.

In order to avoid an error amplification in the inversion algorithm (see chapters 4.3.4 and5.5) we prefer here to analyze the along the line of sight integrated aerosol extinction as it is

Flight Altitude SZA Air-SCD α b χ2

Leon 1996 SS 22.5 km 92.9 5.0 · 1025 cm−2 -0.01 -1.27 9.0 · 10−6

Kiruna 1997 SS 10.0 km 94.5 2.7 · 1026 cm−2 -0.86 -1.55 5.6 · 10−6

16.0 km 93.8 1.2 · 1026 cm−2 0.15 -0.53 8.7 · 10−6

Gap 1997 SR 10.5 km 95.5 2.9 · 1026 cm−2 -0.60 -0.80 1.4 · 10−5

15.5 km 95.0 1.5 · 1026 cm−2 -1.36 -1.84 9.2 · 10−6

Leon 1998 SS 10.5 km 95.3 2.8 · 1026 cm−2 0.74 -0.19 1.6 · 10−5

15.5 km 94.8 1.5 · 1026 cm−2 0.68 -0.34 8.9 · 10−6

Kiruna 1998 SS 10.0 km 95.4 2.7 · 1026 cm−2 -0.42 -1.05 6.9 · 10−6

15.5 km 94.8 1.5 · 1026 cm−2 -0.41 -1.39 4.7 · 10−6

Kiruna 1998 SR 16.0 km 94.3 1.4 · 1026 cm−2 -1.32 -1.34 2.1 · 10−5

Kiruna 1999 SS 10.0 km 94.4 2.9 · 1026 cm−2 0.54 -0.39 1.1 · 10−5

16.0 km 93.6 1.2 · 1026 cm−2 -0.66 -1.55 9.9 · 10−6

Kiruna 2000 SS 10.0 km 94.5 2.7 · 1026 cm−2 0.42 -0.57 1.1 · 10−5

16.0 km 93.8 1.2 · 1026 cm−2 -0.77 -1.75 7.0 · 10−6

Table 7.1: Model parameters α and b obtained for different spectra at different tangent altitudes.The accuracy of the fit is characterized by χ2.

85

86 CHAPTER 7. INFERRED MIE SCATTERING WAVELENGTH DEPENDENCE

much more accurate than inferring the wavelength dependence from vertical extinction profilesat different wavelengths. For each solar occultation measurement that extends into the lower-most stratosphere, some sample spectra are evaluated in the whole spectral range for their Miescattering wavelength dependence. The inferred Mie scattering wavelength dependencies can bereasonably well approximated by the model (Equations 3.41 and 6.2). Table 7.1 shows the modelparameters (α and b) retrieved from a least square fit of the model to the data. In fact, the sec-ond degree polynomial ln (βfit) (Equations 3.41 and 6.2) is fitted to the logarithm of the aerosol

extinction ln (β) as a function of ln(

λλ0

)

, where λ0 is set to 1 µm. Small but sizeable departures

of the measurements from the model are found near the red end of the analyzed spectral range.

For a further interpretation of the results, it is necessary to inspect briefly Mie scatteringcalculations as a function of the aerosol composition and size distribution.

7.2 Calculation of the Mie Scattering Wavelength Dependence

According to the discussion in chapter 3.3, the Mie scattering wavelength dependence of aspherical particle population is calculated from Mie theory. For that purpose the aerosol sizedistribution and the complex refractive index of the aerosols have to be known.

Stratospheric aerosol is primarily composed of liquid droplets containing sulfuricacid and water mixtures for typical stratospheric temperatures [Toon and Pollack 1973;Steele and Hamill 1981]. For lower temperatures water vapor condenses onto the aerosol ac-cording to the law of thermodynamics, changing the aerosol composition from about 75% : 25%(weight percentage) at 220 K to about 50% : 50% at 200 K for the H2SO4/H2O mixture, re-spectively. For temperatures lower than 200 K substantial amounts of HNO3 condense onto theaerosol [Luo et al. 1996], transforming it into ternary H2SO4/HNO3/H2O-solutions. The latteris most important in polar stratospheric winter [Molina et al. 1993].

The equilibrium radius of the stratospheric sulfate aerosol depends on the partial pressures ofwater vapor and sulfuric acid for a given mixture that is for a given radius (due to the Kelvin ef-fect) in equilibrium with the gas phase [[Hamill et al. 1977]; Turco et al. 1979; Toon et al. 1979].Overall for temperatures larger than about 200 K and reasonable stratospheric water contents(vapor pressure around 10−4 – 10−3 mbar) the equilibrium radius is relatively insensitive totemperature and humidity changes.

Model calculations of equilibrium radii and chemical compositions of stratospheric aerosolwere performed by Steele and Hamill [1981] and Carslaw et. al. [1995a; 1995b; 1997]. The modelswere confirmed recently using an Aerosol Composition Mass Spectrometer (ACMS) that allowsto investigate simulated stratospheric aerosols in a laboratory setup [Zink et al. 2001]. Refractiveindices of sulfuric acid solutions have been measured at room temperature for wavelengths downto 360 nm by Palmer and Williams [1975] and from 365 to 214 nm by Beyer et al. [1996]. Therefractive index of a mixture changes proportionally to its density [Dale and Gladstone 1858] andcomposition. In particular, the refractive index of the ternary H2SO4/HNO3/H2O-solution wascalculated by Luo et al. [1996], and measured by Beyer et al. [1996]. Yue et al. [1994] as well asBaumgardner et al. [1996] inferred refractive indices of aerosols from stratospheric observations.

Observations of the stratospheric aerosol size distribution are regularly performed by launch-ing balloon-borne particle counters into the stratosphere at Laramie, Wyoming. For volcanicallyquiet conditions, they can be well represented by a lognormal size distribution with mode radiiranging from 0.06 to 0.09 µm, mode widths from 1.5 to 1.9, and number concentrations from 5

7.2. CALCULATION OF THE MIE SCATTERING WAVELENGTH DEPENDENCE 87

to 15 cm−3 in an altitude range of roughly 15 – 25 km [Deshler 2001]. It is found that the moderadius decreases with increasing altitude.

In this thesis, a computer program written by Levoni [1996] is used for Mie calculations.The algorithm was developed for applications to data gained by the satellite-borne instrumentGOME (Global Ozone Monitoring Experiment). It allows to retrieve aerosol and cloud products[Guzzi et al. 1998].

The program computes the extinction coefficient normalized to 550 nm, the single scatteringalbedo, the asymmetry factor, and the phase function of a lognormal aerosol size distribution.The wavelength dependent refractive index of the liquid aerosol, the mode radius and modewidth of the aerosol size distribution are required as input parameters. In the present versionthe program works in a spectral range from 0.3 to 0.86 µm, so that the infrared channels ofSAGE and POAM are unfortunately not covered. Table 7.2 summarizes the refractive indicesused as input parameters for present the Mie calculations. The refractive indices in the regionsbetween the given wavelengths are interpolated. The values are typical for stratospheric aerosolhaving a weight percentage of H2SO4 in the range of 60 – 70%. Therefore the calculations arereasonably good for stratospheric temperatures between 210 and 230 K.

Table 7.2: Refractive indices ofstratospheric sulfate aerosolsused as input parametersfor the Mie calculations[WCP-112 1986].

Wavelength Refractive Index

300.0 nm 1.469337.1 nm 1.459400.0 nm 1.440488.0 nm 1.432514.5 nm 1.431550.0 nm 1.430632.8 nm 1.429694.3 nm 1.428860.0 nm 1.425

Figure 7.1 shows calculated aerosol extinctions as a function of the wavelengths used forthe DOAS aerosol retrieval. Unimodal lognormal size distribution of mode radius rm = 0.07µmand different mode widths between 1.26 and 2.00 are assumed. Figure 7.2 shows results bychanging the mode radius but keeping the mode width at 1.78. The calculated aerosol extinctioncoefficients are normalized to the extinction at 550 nm.

It can be seen, that a decrease in the mode radius has the same effect on the wavelengthdependence as decreasing the mode width. This demonstrates that, an unambiguous derivationof the aerosol size distribution from extinction measurements is impossible without confining theaerosol size parameters to reasonable values given by in-situ measurements and model calcula-tions. The slope of the aerosol extinction with wavelengths in a double-logarithmic plot (relatedto a wavelength dependent Angstrom coefficient, see chapter 3.5 and Formula 7.1) increases ifthe mode radius or mode width decrease. The wavelength dependent aerosol extinction coeffi-cients all show small negative curvatures that can be approximated by a linear function for smallmode widths and small mode radii (c. f. the Rayleigh extinction would have an almost constantwavelength exponent of around -4).

When using a bimodal lognormal size distribution, the Mie calculation is performed sepa-rately for both modes. The obtained extinction coefficients are weighted by a fraction factor,proportional to the relative particle concentration of the two modes, which are then accordingly


0.45 0.5 0.55 0.60.6

0.7

0.80.9

1

2

V = 2.00 V = 1.78 V = 1.58 V = 1.41 V = 1.26A

ero

sol E

xtin

ctio

n [a

. u.]

Wavelength [µm]0.45 0.5 0.55 0.6

0.6

0.7

0.80.9

1

2

rm = 0.1 µm

rm = 0.07 µm

rm = 0.045 µm

rm = 0.02 µm

rm = 0.01 µmA

ero

sol E

xtin

ctio

n [a

. u.]

Wavelength [µm]

Figure 7.1: Calculated aerosol extinctions asa function of wavelength (and normalized at550 nm) for unimodal lognormal size distri-butions of mode radius rm = 0.07µm and

different mode widths σ.

Figure 7.2: Same as Figure 7.1 but changingthe mode radius but keeping the mode width

at σ = 1.78.

added to obtain the total extinction of bimodal distribution. The short wavelength extinctionis dominated by small particles, whereas the large wavelength extinction is dominated by largeparticles. Hence, with respect to a unimodal size distribution the curvature of the wavelengthdependence in a double-logarithmic plot increases for a bimodal size distribution. For certaincombinations of bimodal size distributions, the curvature can become positive depending on theconsidered wavelength interval.

In the SAGE and POAM aerosol retrievals, the wavelength dependence is approximated bythe model given in Equations 3.41 and 6.2. Using this model, the slope of the Mie scatteringwavelength dependence is described by a parameter α, and the curvature is described by aparameter b. The model is, however, not able to approximate the measured aerosol extinctionfor any wavelength simultaneously. In particular for bimodal lognormal size distributions, thesecond derivative may change sign (e. g. a positive curvature in the visible spectral range becomesnegative in the infrared, if the extinction due to the small particle mode becomes negligible atlarge wavelengths).

7.3 Discussion

7.3.1 Mie Scattering Wavelength Dependence

In the DOAS retrieval, for every spectrum b < 0 is obtained, indicating that the observedwavelength dependence has positive curvature in the analyzed wavelength range (see Table 7.1).The parameter α, however, is strongly dependent on the choice of λ0, and b.

A wavelength dependent Angstrom coefficient a (λ) is obained by inserting into the derivativeof Equation 3.41 with respect to lnλ Angstrom’s formula (Equation 3.37),

a (λ) = −d ln (βfit (λ))

d (lnλ)= α+ 2b · ln

(

λ

λ0

)

. (7.1)

7.3. DISCUSSION 89

0.4 0.5 0.6 0.7 0.8 0.9 10.2

0.3

0.4

0.50.60.70.80.9

1

2

rm = 0.09 µm, V = 2.00

rm = 0.07 µm, V = 1.86

rm = 0.06 µm, V = 1.58

A

ero

sol E

xtin

ctio

n [a

. u.]

Wavelength [µm]0.4 0.5 0.6 0.7 0.8 0.9 1

0.2

0.3

0.4

0.50.60.70.80.9

1

2

Aer

oso

l Ext

inct

ion

[a. u

.]

Wavelength [µm]

Figure 7.3: Mie calculated extinctions used forcomparison with satellite and balloon data in Fig.

7.4 to 7.7

Figure 7.4: Leon 1996: 22.5 km. In Fig. 7.4 to7.7 show red lines: balloon data, black crosses:satellite data, blue lines: Models shown in Fig.

7.3

0.4 0.5 0.6 0.7 0.8 0.9 10.2

0.3

0.4

0.50.60.70.80.9

1

2

Aer

oso

l Ext

inct

ion

[a. u

.]

Wavelength [µm]0.4 0.5 0.6 0.7 0.8 0.9 1

0.2

0.3

0.4

0.50.60.70.80.9

1

2

Aer

oso

l Ext

inct

ion

[a. u

.]

Wavelength [µm]

Figure 7.5: Leon 1998. Left panel: 10.5 km, Right panel: 15.5 km

0.4 0.5 0.6 0.7 0.8 0.9 10.2

0.3

0.4

0.50.60.70.80.9

1

2

Aer

oso

l Ext

inct

ion

[a. u

.]

Wavelength [µm]0.4 0.5 0.6 0.7 0.8 0.9 1

0.2

0.3

0.4

0.50.60.70.80.9

1

2

Aer

oso

l Ext

inct

ion

[a. u

.]

Wavelength [µm]

Figure 7.6: Kiruna 1999. Left panel: 10 km, Right panel: 16 km

0.4 0.5 0.6 0.7 0.8 0.9 10.3

0.4

0.5

0.60.70.80.9

1

2

Aer

oso

l Ext

inct

ion

[a. u

.]

Wavelength [µm]0.4 0.5 0.6 0.7 0.8 0.9 1

0.2

0.3

0.4

0.50.60.70.80.9

1

2

Aer

oso

l Ext

inct

ion

[a. u

.]

Wavelength [µm]

Figure 7.7: Kiruna 2000. Left panel: 10 km, Right panel: 16 km


At 500 nm, Angstrom coefficients ranging between 0.9 and 1.2, and 1.2 and 1.6 for about 10km and 16 km, respectively, are obtained. This result is in good agreement with the wavelengthexponents inferred by Barnes and Hofmann [2001] (Figure 3.4). For large tangent altitudes, theAngstrom coefficient increases, thus indicating a decreasing particle radius (see Figure 7.2). Forthe Leon flight in 1996, a = 1.75 is obtained at a tangent altitude of 22.5 km, somewhat largerthan observed by Barnes and Hofmann [2001].

According to the discussion given in the preceding section, the detected positive curvatureof the Mie scattering wavelength dependence can only be described by a bimodal lognormal sizedistribution. This is contrary to the satellite observations and Mie calculations based on in-situmeasured aerosol size distributions.

Figures 7.4 to 7.7 compare the wavelength dependent aerosol extinction of the present studywith collocated SAGE II or POAM III measurements (for a more detailed description of thesatellite measurements see section 7.3.3). Figure 7.3 shows calculations of aerosol extinctionsfor typical aerosol radius and mode distributions measured by in-situ optical particle countersfrom 15 to 25 km [Deshler 2001]. The red curve in the Figure is calculated using standard inputparameters given by [Levoni 1996].

For each DOAS measurement (red curves), the slope of the curve (Angstrom coefficient) fallswithin the range of the calculated wavelength dependent extinctions (Figure 7.3). The wave-length dependence observed by the satellite instruments, only falls between the limits given byparticle counter observations in some cases. The curvature of the aerosol extinction wavelengthdependence, however, is stronger and more negative than the corresponding value assuming aunimodal lognormal distribution. This suggests a larger mode radius seen by the satellite in-struments than a unimodal lognormal distribution of rm = 0.09µm and σ = 2.00 can accountfor. Assuming larger aerosols were actually there, would conversely result in a flatter slope forlarge wavelength than the satellites are actually detecting. Hence, the wavelength dependenciesof stratospheric aerosol extinction found from particle counter measurements, satellite borne,and balloon borne observations are inconsistent.

7.3.2 Apparent Size Distribution inferred for the DOAS Observations

It is clear that the aerosol extinction inferred from balloon and satellite observations is alwaysan integral along the line of sight, i. e. Mie scattering from aerosols located at all altitudes fromballoon float down to the tangent height contribute to the measured aerosol extinction. There-fore, the inferred apparent wavelength dependence is a complicated function of the wavelengthdependencies in all probed atmospheric layers. Accordingly, since small particles tend to be morepresent at large altitudes their influence to the apparent aerosol extinction potentially contam-inates the Mie scattering wavelength dependence attributed for tangent height observations.

As shown in the preceding section the aerosol extinction inferred from the DOAS measure-ments cannot be properly described by a unimodal lognormal size distribution. Since both adecrease in the mode radius or width of the assumed lognormal distribution have a similar effecton the extinction wavelength dependence, our measurements require at least a narrow small,and a broader large particle mode. In order to demonstrate possible solutions only some samplespectra are considered in detail. Since several sets of the four parameters of a bimodal lognormalsize distribution may lead to extinction wavelength dependencies compatible to our observations,the result is never unambiguous.

7.3. DISCUSSION 91

For example, Figures 7.8 and 7.9 show assumed size distributions and calculated Mie scat-tering wavelength dependencies which fit best to two spectra (tangent height 10 and 16 km, re-spectively) taken during the Kiruna balloon flight on February 18, 2000. For both observations,a large particle mode with a mode radius rm = 0.13µm and a mode width σ = 2.24 is requiredto properly fit the small slope in the wavelength dependencies at the long wavelength end. Forthe small particle mode, best agreements result for rm = 0.045µm and σ = 1.58/rm = 0.02µmand σ = 1.26 for the 10/16 km observation, respectively.

0.01 0.1 110-2

10-1

100

101

102

103

Par

ticl

e N

um

ber

[a. u

.]

Radius [µm]0.01 0.1 1

10-4

10-3

10-2

10-1

100

101

102

103

104

Par

ticl

e N

um

ber

[a. u

.]

Radius [µm]

Figure 7.8: Relative particle radius distribution of an aerosol population causing the extinction wave-length dependencies shown in Figure 7.9.

0.5 0.6 0.7 0.80.3

0.4

0.50.60.70.80.9

1

2

3

Aer

oso

l Ext

inct

ion

[a. u

.]

Wavelength [µm]0.5 0.6 0.7 0.8

0.3

0.4

0.50.60.70.80.9

1

2

3

Aer

oso

l Ext

inct

ion

[a. u

.]

Wavelength [µm]

Figure 7.9: Two sample aerosol extinction wavelength dependencies observed during Kiruna balloonflight in 2000 at tangent altitudes of 10 km (left panel) and 16 km (right panel). The extinctioncomputed for a bimodal lognormal size distribution (black curve) is approximated to the observedextinction (red curve). The blue lines show the resulting Mie scattering wavelength dependenceof each of the two unimodal size distributions individually.Left panel: rm,1 = 0.045µm, σ1 = 1.58, rm,2 = 0.13µm, σ2 = 2.24,

βtotal (550 nm) = 0.342 · β1 (550 nm) + 0.658 · β2 (550 nm).Right panel: rm,1 = 0.02µm, σ1 = 1.26, rm,2 = 0.13µm, σ2 = 2.24,

βtotal (550 nm) = 0.302 · β1 (550 nm) + 0.698 · β2 (550 nm).


Note that these solutions assuming bimodal lognormal size distributions are only examplesfrom a larger set of possible solutions. Likewise, since the mean particle radius is continuouslydecreasing with increasing altitude, other asymmetric size distributions lead to possible solutionsas well. Also the required small particle mode cannot be explained by observational effects, i.e. the fact that the contribution from small particles at large altitudes always contaminate thesignal from large particles at low altitudes. This is since the aerosol size distribution continuouslychange with height.

For both observations (tangent heights 10 and 16 km), it is found that the small particlemode contributes 34.2% and 30.2% to the total extinctions at 550 nm, respectively (see Figure7.9). Since Mie scattering in the visible spectral range due to small droplets is much less efficientthan due to large droplets, in order to result in a significant contribution the number of smallparticles has to be much larger than of large particles (see Figure 7.8).

7.3.3 Comparison with Auxiliary Data

Comparison with Satellite Data

Figures 7.4 to 7.7 also show a comparison of the inferred Mie scattering wavelength dependenceof the present study with collocated SAGE II or POAM III measurements.

It is clear that both satellites infer a negative curvature of the Mie scattering wavelengthdependence which contrasts our observations. There is no obvious reason for this discrepancy.While the aerosol extinction errors of the present studies are discussed in chapter 5.5, theerrors of SAGE II and the POAMs are described in detail in Chu et al. [1989] (SAGE II),Lumpe et al. [1997] (POAM II), and Lucke et al. [1999] (POAM III). It is also clear that bothobservation methods have their strengths and weaknesses. With DOAS a rather limited wave-length interval located in the visible region is observed, however, in 1024 independent andnarrowly spaced wavelength bands. Conversely, the satellites cover a much larger wavelengthinterval, however, in much less channels. Therefore, the strength of DOAS is to measure dif-ferential absorption structures of atmospheric absorbers (O3, NO2, O4, OClO, IO, OIO) veryprecisely, for the sake not covering wavelength regions that are in particular sensitive to aerosolextinction. Conversely, the satellites’ strength is to measure the aerosol extinction in a purelyaerosol sensitive band (near IR). Their potential weakness lies in the missing independenceof the trace gas measurements and inferred aerosol extinction at shorter wavelengths. For ex-ample, if both satellites were to attribute less measured extinction in the visible channels at525 nm (SAGE II) and 603 nm (POAM III) to ozone absorption, they would measure moreaerosol extinction. Quantitatively, agreement of our and the satellites’ observations, however,would require to reduce the attributed ozone extinction of the satellites by at least a factor of2. This requirement would cause the satellites’ ozone measurements to deviate largely from thevalidation instruments’ measurements. There is no primitive solution to that dilemma.

Conversely, it is questionable whether the satellite instruments are sensitive at all to thesuggested small particle mode. Since their visible NO2 is a differential measurement in twoadjacent wavelength bands, the inferred NO2 is hardly affected by a suggested small particlemode. Similarly, a small particle mode would dramatically affect the attributed ozone extinctionin the visible channels but much less in the red channels. Therefore, at this point the onlyconsistent solution is to require more efficient Rayleigh scattering in the visible wavelengthregion than attributed to it by the satellites.

7.3. DISCUSSION 93

Moreover, if one were to assume the satellite IR aerosol extinction and the small particlemode hypothesis were correct within the given error bars, then the satellites’ aerosol extinctioninferred for the visible channels would become larger but not as large as required to close the”factor of 2 gap” (see chapter 6.3).

Also, our Mie calculation (assuming a typical stratospheric aerosol size distribution likemeasured by [Deshler 2001], see Figure 7.3) cannot reproduce the Mie scattering wavelengthdependence inferred from the satellites’ measurements. At this point, we see no reason for thatdiscrepancy.

Comparison with Optical Particle Counters (OPC)

The inferred aerosol size distributions can also be compared with OPC measurements[Rosen 1964]. The OPC instruments of the University of Wyoming measure the stratosphericaerosols by illuminating particles in a sample chamber using a white light source. The scatteredlight is observed at a fixed angle range in forward direction. Using this technique particles withradii larger than about 0.15 µm can be detected directly. Smaller particles of radii down to0.01 µm are detected by using a growth chamber ahead of an OPC [Rosen and Hofmann 1977].There, the particles grow to optically detectable sizes. The minimum detectable concentrationsare approximately 0.007 cm−3 for particles > 0.01µm, and 0.0006 cm−3 for particles > 0.15µm[Deshler et al. 1993]. The instruments count the integrated number of particles larger than athreshold radius, subdivided into different size classes. An integral of particle size distributions(unimodal or multimodal lognormal functions) taken from a threshold radius to infinity are fit-ted to the observed particle concentrations. Some of the OPC measurements and the inferredapproximated size distributions are shown in Figure 3.3.

The OPC measurements are very robust for large particles. Although large particles areeasily detectable and contribute significantly to the measured aerosol extinction, their detectionis limited by the small numbers causing large statistical errors. In practice, this limits the aerosoldetection to particles smaller than 1 µm (see error bars in Figure 3.3). In the stratosphere, most ofthe sulfate aerosols are smaller than 0.15 µm. Due to their large number, they can be accuratelymeasured. Much smaller particles (some 0.01 µm in size), however, are difficult to be detected,since prior to detection they have to grow by more than a factor of 10.

Accordingly, bimodal size distributions as emphasized by our study (see Figure 7.8) canhardly be detected by OPCs. This is since unimodal size distribution with a medial mode radiuslook very similar to bimodal particle distributions, if the number of large particles is slightlyunderestimated and the radius of small particles is slightly overestimated. Thus, small errorsin the OPC’s detection of abundant small and sparse large particles would prevent a decisionwhether the size distribution is unimodal or bimodal.

Likewise, the aerosol size distribution parameters inferred from satellite observations showsome discrepancies to OPC measurements as well. Hervig and Deshler [2001] compared resultsfrom SAGE II (0.386, 0.452, 0.525, and 1.02 µm) and HALOE (2.45, 3.40, 3.46, and 5.26µm) multi-wavelength aerosol extinction observations with balloon-borne OPC measurements[Deshler et al. 1993]. They found the inferred SAGE II aerosol surface areas to be smaller thanthose inferred from HALOE and OPC measurements, in particular when the aerosol numbersare small. According to Hervig and Deshler [2001], one reason could be that SAGE II is lesssensitive to small aerosols.


Summarizing, the OPC measurements can hardly provide robust arguments which may serveto decide whether the stratospheric aerosol size distribution is unimodal or bimodal.

7.4 Suggestions and Recommendations for Future Investiga-tions

The retrieval of aerosol size distributions from DOAS extinction data is still in an early stage. Wediscussed some preliminary results that indicate some striking discrepancies to the measurementsof other instruments. In addition, it is shown that the aerosol extinction wavelength dependenceobserved over a limited wavelength range is not sufficient to unambiguously infer aerosol sizedistributions. Systematic studies of various size distributions may help to reconcile the differentfindings of the various instruments. Thereby, the effects of varying mode radius, mode width,and refractive index should be analyzed and compared with the DOAS results.

Vertical extinction profiles for each DOAS wavelength channel should also be retrieved (asdescribed in section 4.3.4 and chapter 5), which would allow to analyze variations of the observedaerosol extinction wavelength dependence with altitude. Then, for different altitude segments,the size distribution parameters (mode radius, mode width, and refractive index) could be in-ferred and compared with satellite, LIDAR, and OPC measurements.

The LPMA IR measurements should be investigated for the aerosol extinction, in order toextend the considered wavelength interval into the infrared. This would allow a more directintercomparison with the satellite IR aerosol channels.

An improved software for Mie calculations is essential to extrapolate the aerosol extinctionto wavelengths larger than 0.86 µm. This improved software should also allow to calculateabsolute extinction cross sections per particle for given radii, which would allow us to inferabsolute particle number densities. Furthermore, different size distribution models should beexamined. Possibly, other size distributions than lognormal distributions would better agreewith our observations, in particular for the emphasized small particle mode.

Chapter 8

Conclusion and Outlook

The present thesis aimed at the development of a technique to infer aerosol data fromballoon-borne optical DOAS extinction measurements. The observations by the DOAS spec-trograph described in chapter 4.4 was used in recent studies to measure trace gas concentra-tions in the stratosphere [e. g. Ferlemann et al. 1998b and 2000; Harder et al. 1998 and 2000;Fitzenberger et al. 2000a; Camy-Peyret et al. 1999; Pfeilsticker et al. 2000 and 2001] with a spe-cial regard to the radicals relevant for the ozone depletion in the stratosphere (ClO, BrO, NO2,see chapter 2.3.2). In future, aerosol size distributions and surface areas inferred from the samedata can be used for modelling the heterogenous chemistry of the observed trace gases. Identicalmeasurement geometry, temporal and spatial coincidence of aerosol and trace gas observationsprovide best conditions for a proper verification of the chemical models.

Aerosol extinction measurements are performed by most of the satellite-borne instrumentsglobally observing the stratospheric aerosol content, but only by few other instruments. Withinthe scope of this thesis, aerosol extinctions could be retrieved over a continuous wavelengthrange from 440 to 615 nm, which permits an extensive validation of satellite aerosol extinctionmeasurements in the visible spectral range. So far, the latter were only validated by other satelliteobservations, or inferred parameters such as size distributions were compared.

The results presented in chapter 6.3 reveal large discrepancies between the aerosol extinctiondata by SAGE II and POAM III and the DOAS data. The inferred extinction wavelengthdependence differs significantly from the satellite multiwavelength observations. Preliminary Miecalculations indicate the presence of a small particle mode with radii of a few tens of nanometers.Since in a first approximation, the aerosol extinction coefficient and the aerosol surface area areproportional to the square of the particle radius, the small particle mode significantly affects theheterogenous stratospheric chemistry and the observed aerosol extinction at short wavelengths,whereas the total mass is negligible compared to the large aerosol mode. Further studies asdescribed in chapter 7.4 have to be performed to properly infer the size distribution and theparticle density from the DOAS observations. Then, the data can be directly used for a revisionof the chemical models, and for comparison with instruments observing different spectral rangesor scattered light.

The main source of error of the aerosol extinction retrieval are uncertainties in the air data,resulting in errors in the calculated Rayleigh extinction (see chapter 5.5). In particular, thepressure data at high altitudes is not very accurate. Using better pressure sensors aboard thegondola or aboard simultaneously launched radiosondes will enhance the accuracy of the inferredaerosol extinction significantly.

95

96 CHAPTER 8. CONCLUSION AND OUTLOOK

Implementing the aerosol extinction retrieval algorithm developed in this thesis, future bal-loon flight campaigns using the DOAS spectrograph will provide extensive data for the valida-tion of the near future satellite-borne instruments SAGE III and SCIAMACHY. SAGE III waslaunched on December 10, 2001 aboard Meteor-3M satellite, and will start to provide aerosolextinction data from the beginning of January 2002. The launch of SCIAMACHY is scheduledfor March 1st, 2002 aboard EnviSat-1. Atmospheric extinction will be observed over a wide con-tinuous spectral range from the UV to the infrared with a spectral resolution of better than 1nm. Hence, the DOAS technique can be used to infer trace gas and aerosol extinction profilesfrom the SCIAMACHY data. Similar measurement geometry and retrieval algorithms make theballoon-borne DOAS spectrograph to be of major importance for the SCIAMACHY validation.

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Danksagung

Hiermit mochte ich all jenen danken, die mich fachlich wie auch privat bei der Durchfuhrungdieser Arbeit unterstutzt haben.

Zunachst mochte ich meinem Betreuer Dr. Klaus Pfeilsticker danken, der mir mit fachlichenund sprachlichen Tips beim Zusammenschreiben der Arbeit sehr geholfen hat, und wahrend desganzen Jahres mit fachlicher Beratung, Beschaffung interessanter Literatur und nicht zuletztdurch haufige Motivierung in scheinbar weniger produktiven Phasen einen erheblichen Teil zumGelingen dieser Arbeit beigetragen hat. Besonders freue ich mich auch uber die Gelegenheit,meine erste wissenschaftliche Publikation schon uber das Thema der Diplomarbeit schreiben zukonnen.

Herrn Prof. Dr. Ulrich Platt danke ich fur die Schaffung des wissenschaftlichen und organ-isatorischen Rahmens fur eine gute Kooperation zwischen den einzelnen Arbeitgruppen mit un-terschiedlichen aber nicht weniger interessanten Themenbereichen, fur die Zweitkorrektur dieserArbeit und fur Feuerzangenbowle und Wichteln auf den arbeitsgruppeninternen Weihnachts-feiern, die eine gelungene Einstimmung und einen kronenden Abschluß meines Diplomarbeits-jahres darstellten.

Ich danke meinen Kollegen Hartmut Bosch fur die ausfuhrliche und verstandliche Einar-beitung in die theoretischen Hintergrunde und fur die Betreuung beim Zusammenschreiben derArbeit trotz eigenen Stresses wegen der Doktorarbeit, Dr. Richard Fitzenberger fur die Ver-anschaulichung experimenteller und elektronischer Details des DOAS-Spektrographen und furdie Rettung der Kiruna-2001-Kapagne in letzter Sekunde, Gerd Honninger fur LaTeX-Tips,die tatsachlich funktionierten, Marcel Dorf dafur, daß er innerhalb weniger Tage drei Vier-tel meiner Arbeit korrekturgelesen hat, so daß sie nun hoffentlich auch fur ”native speaker”verstandlich ist, Frank Weidner fur Korrektur des verbleibenden Viertels und fur Aufstockungmeiner CD-Sammlung mit guter Musik, und letzteren beiden auch noch fur eine unterhaltsameKiruna-Kampagne und die nachfolgende Norwegen-Tour.

Besonders hervorheben mochte ich das zwischenmenschliche Arbeitsklima am Institut furUmweltphysik, das neben der interessanten Wissenschaft erheblich dazu beigetragen hat, jedenTag erneut mit großer Motivation an die Arbeit zu gehen, und den regen Austausch zwischenden Arbeitsgruppen, der meinen fachlichen Horizont deutlich erweitert hat.

Ein ganz besonderer Dank geht auch an diejenigen, die abseits des Instituts dafur gesorgthaben, daß es mir gut geht und daß ich die kleinen Probleme des Alltags bewaltige, dennnur unter diesen Voraussetzungen war es mir uberhaupt moglich tagsuber meine volle Leis-tungsfahigkeit zu entfalten. Hierbei denke ich in erster Linie an meine Mutter, meinen Vater undmeine Schwester, die mich seelisch, moralisch und zum Teil sogar fachlich unterstutzt haben.Genauso wichtig fur mich waren auch meine Freunde, von denen ich hier insbesondere Volker,Mareike und Sabine erwahnen mochte, die immer Zeit und ein offenes Ohr fur mich hatten, ins-besondere wenn es mal nicht um Physik ging. Ihr wißt selbst am besten, was Ihr mir bedeutet.

113

Erklarung

Ich versichere, daß ich diese Arbeit selbststandig verfaßt und keine anderen als die angegebenenQuellen und Hilfsmittel benutzt habe.

Heidelberg, den

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