Carbonaceous Aerosol

CARBONACEOUS AEROSOL

ATMOSPHERIC AND OCEANOGRAPHIC SCIENCES LIBRARY VOLUME 30

Editors Lawrence A. Mysak, Department of Atmospheric and Oceanographic Sciences, McGill University, Montreal, Canada Kevin Hamilton, International Pacific Research Center, University of Hawaii, Honolulu, HI, U.S.A.

Editorial Advisory Board L. Bengtsson A. Berger P.J. Crutzen J.R. Garratt G. Geernaert M. Hantel A. Hollingsworth H. Kelder T.N. Krishnamurti P. Lemke P. Malanotte-Rizzoli S.G.H. Philander D. Randall J.-L. Redelsperger R.D. Rosen S.H. Schneider F. Schott G.E. Swaters J.C. Wyngaard

Max-Planck-Institut für Meteorologie, Hamburg, Germany Université Catholique, Louvain, Belgium Max-Planck-Institut für Chemie, Mainz, Germany CSIRO, Aspendale, Victoria, Australia DMU-FOLU, Roskilde, Denmark Universität Wien, Austria European Centre for Medium Range Weather Forecasts, Reading, UK KNMI (Royal Netherlands Meteorological Institute), De Bilt, The Netherlands The Florida State University, Tallahassee, FL, U.S.A. Alfred-Wegener-Institute for Polar and Marine Research, Bremerhaven, Germany MIT, Cambridge, MA, U.S.A. Princeton University, NJ, U.S.A. Colorado State University, Fort Collins, CO, U.S.A. METEO-FRANCE, Centre National de Recherches Météorologiques, Toulouse, France AER, Inc., Lexington, MA, U.S.A. Stanford University, CA, U.S.A. Universität Kiel, Kiel, Germany University of Alberta, Edmonton, Canada Pennsylvania State University, University Park, PA, U.S.A.

The titles published in this series are listed at the end of this volume.

CARBONACEOUS AEROSOL

by

ANDRÁS GELENCSÉR Air Chemistry Group of the Hungarian Academy of Sciences, University of Veszprém, Veszprém, Hungary

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CONTENTS Preface……….. ........................................................................................................vii Acknowledgements ...................................................................................................ix 1. Introduction ..........................................................................................................1 2. Methods of Observation of Carbonaceous Aerosol ...........................................7 2.1 Sampling of Carbonaceous Aerosol ...............................................................7 2.1.1 Filter Sampling ....................................................................................8 2.1.2 Impactor Sampling ............................................................................21 2.2 Methods for Determination of Main Carbonaceous Aerosol Components ....................................................................................22 2.2.1 Methods for BC Determination ......................................................... 23 2.2.2 Methods for EC Determination ......................................................... 31 2.2.3 Thermo-optical Methods for OC/EC analysis ...................................35 2.2.4 Intercomparison of Various Methods ................................................42 3. Major Carbonaceous Particle Types and their Sources .................................45 3.1 Soot—Black Carbon (BC), Elemental Carbon (EC), Graphitic Carbon .......45 3.1.1 Definitions ......................................................................................... 45 3.1.2 History of Soot Pollution ...................................................................48 3.1.3 Properties and Chemical Composition of EC/BC/Soot ..................... 49 3.1.4 Emission Sources of EC/BC/Soot ..................................................... 55 3.1.5 Atmospheric Observations of BC/Soot ............................................. 58 3.1.6 Global Models of BC ........................................................................63 3.2 Primary Organic Aerosol ..............................................................................69 3.2.1 Natural Primary Aerosol Sources....................................................... 70 3.2.2 Anthropogenic Primary Aerosol Sources ..........................................74 3.3 Secondary Organic Aerosol (SOA) ..............................................................83 3.3.1 Emissions of Volatile Organic Compounds (VOC) relevant to Secondary Organic Aerosol Formation ...........................84 3.3.2 Gas-phase Chemical Reactions Leading to Condensable Species .....89 3.3.3 Gas-to-particle Partitioning of Reaction Products ........................... 110 3.3.4 Atmospheric Observations of Secondary Organic Aerosol ............. 129 3.3.5 Modeling of SOA Formation .......................................................... 131 3.3.6 Secondary Aerosol Formation in Heterogeneous and Multiphase Processes ...................................................................... 141 4. Organic Chemistry of Aerosol ........................................................................ 149 4.1 Chemical Properties of Bulk Organic Aerosol ........................................... 151 4.1.1 OC Mass Concentrations ................................................................. 151 4.1.2 Global Modeling of OC Concentration Distributions ..................... 155

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vi 4.1.3 Size Distribution of Organic Aerosol .............................................. 159 4.1.4 Organic Aerosol in Mass Closure Studies ....................................... 161 4.1.5 State of Mixing of Organic Aerosol as Inferred from Individual Particle Analyses ............................................................ 166 4.1.6 Bulk Chemical Characterization of OC ........................................... 168 4.1.7 Bulk OC/BC Source Apportionment ............................................... 180 4.2 Major Compound Classes and Atmospheric Tracer Compounds ............... 184 4.2.1 Major Classes of Organic Tracers ................................................... 184 4.2.2 Source-specific Organic Tracer Compounds ................................... 195 4.3 Chemistry of Water-Soluble Organic Compounds (WSOC) ...................... 211 4.3.1 Water-soluble Organic Compounds ................................................ 211 4.3.2 Scavenging Efficiency for Organic Species .................................... 217 4.3.3 Organic Species in Cloud/fog Water and Precipitation ................... 218 5. Properties of Carbonaceous Aerosol and their Role in the Global Atmosphere ....................................................................................................... 221 5.1 Optical Properties of Carbonaceous Particles and their Role in Radiative Transfer in the Atmosphere …..……………………………………………221 5.1.1 Optical Properties of Individual Carbonaceous Aerosol Components ....................................................................... 222 5.1.2 Observations Regarding the Optical Effects of Carbonaceous Aerosol ..................................................................... 236 5.1.3 Role of Carbonaceous Aerosol in Radiative Transfer in the Atmosphere ...................................................................................... 242 5.2 Hygroscopic Properties of Carbonaceous Particles and their Atmospheric Implications .................................................................. 256 5.2.1 Hygroscopic Growth Attributed to Carbonaceous Aerosol ............. 258 5.2.2 Cloud Condensation Nucleus (CCN) Activity of Carbonaceous Aerosol ..................................................................... 268 5.2.3 Ice-nucleating (IN) Activity of Carbonaceous Aerosol ................... 288 5.2.4 Contribution of Carbonaceous Aerosol to Aerosol Indirect Forcing 289 5.3 Heterogeneous Reactions on Carbonaceous Aerosol in the Atmosphere ... 295 5.3.1 Properties of Carbonaceous Particles Relevant to Heterogeneous Chemistry ............................................................... 295 5.3.2 Heterogeneous Reactions in the Stratosphere on Carbonaceous Aerosol ..................................................................... 296 5.3.3 Heterogeneous Reactions in the Troposphere on Carbonaceous Aerosol ..................................................................... 299 References .......................................................................................................... 301 Copyright Acknowledgements ............................................................................... 333 Index .......................................................................................................... 343 List of Abbreviations .............................................................................................. 349

PREFACE Recently there has been heightening scientific interest in carbonaceous aerosol which has been shown to be important in inadvertent climate modification yet it is poorly understood. In particular, carbonaceous aerosol was started to be viewed from a global perspective just barely more than a decade ago, and there are still many conflicting views in the field even regarding basic nomenclature and methodology. Despite of its atmospheric significance, only a handful of summary works have been published on some parts of the subject so far, in the form of book chapters or review papers. This monograph is the first attempt to condense all major aspects of carbonaceous aerosol chemistry and its atmospheric effects into a single volume. This endeavor inevitably involves many compromises. The most important is that the reader is assumed to be largely familiar with atmospheric chemistry in general, and more specifically with atmospheric aerosol chemistry. Nevertheless, the individual chapters of the book are intended to be written in such a style that they remain readable and understandable for a wide audience interested in any aspect of atmospheric science. Therefore, the average level of the book is set to about that of an advanced textbook. Since the research on carbonaceous aerosol is relatively recent and lacks the time and perspective to crystallize, many of its aspects are still being intensively studied, and quite often debated within the scientific community. This makes writing such a book a great challenge on the one hand, and an extremely risky business on the other. To reduce the risks, the book tries to be “democratic” in the sense that it presents contrasting ideas without prejudice, and acknowledges “majority votes” wherever they are applicable. Currently, the integration of available information on carbonaceous aerosol is being done in global atmospheric models. This book may provide a background for this work by presenting the latest achievements in each field, and pinpointing the sources of uncertainties and caveats of simplification that is usually required for model parameterization. Another objective of this work is to attempt to create internal consistency between the separate fields with a view to drawing a coherent picture of carbonaceous aerosol in the global atmosphere. This may help identify further research needs in the field and inspire atmospheric scientists and other researchers from various disciplines to venture into the still unexplored reigns of carbonaceous aerosol. András Gelencsér Veszprém, June 2004

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ACKNOWLEDGEMENTS Despite being the sole author of this book, I would have never succeeded without the support of many people and organizations whom I wish to acknowledge herein. I was encouraged to take the challenge to write this book by Ernő Mészáros, my mentor and one of the pioneers in atmospheric aerosol chemistry, whose exemplary scientific career and attitude inspired me. I am very much indebted to my immediate colleagues Szilvia Janitsek and Kornélia Imre for their tremendous help in bringing the manuscript to its final edited form. I have received useful comments on the drafts from my friends and fellow researchers Ágnes Molnár, Gyula Kiss, and Mihály Pósfai to whom I am thankful. I feel privileged to be able to work in the hospitable atmosphere of the Department of Earth- and Environmental Sciences at the University of Veszprém, having benefited from the help of its chair, József Hlavay, and many other colleagues, Mrs. Irén Bakos-Szalai, Mrs. Katalin Galló-Békefi, Ms. Eszter Tornyos, to name just a few. I am also very grateful to my editor at Kluwer, Marie Johnson, for her assistance in the preparation of this book. The Hungarian Academy of Sciences and the University of Veszprém have provided me with a sound background by supporting and hosting the Air Chemistry Research Group which I have been a member of. The Ministry of Education, the National Science Fund of Hungary as well as the European Commission have granted research funds which relieved me from the pressure of seeking others, thus indirectly helping me concentrate on writing this book. Last but not least, I would like to thank my family for their help, understanding and patience, my wife, Ágnes, and my little son, András, who was born at the same time as the idea of writing this book and has grown along the way to its completion.

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Chapter 1 INTRODUCTION The research of carbonaceous aerosol has recently become one of the most favorite topics in the field of atmospheric sciences. The main reason for its popularity is that carbonaceous aerosol represents an extreme diversity within atmospheric aerosol. No matter from which standpoint researchers look at atmospheric aerosol they likely run into some type of carbonaceous aerosol. For those who are concerned with the health effects of air pollution, there are thousands of potentially harmful organic compounds associated with the respirable fraction of atmospheric aerosol. Those, on the other hand, who are primarily interested in inadvertent climate modification, will find by far the largest uncertainties accompanying the effects of carbonaceous aerosol. It is now understood that important atmospheric reactions take place on or are affected by the surface of carbonaceous aerosol particles. Even in cloud formation, which had long been thought not to be affected by aerosol components other than sulfate or sea salt, carbonaceous aerosol now seems to play an important role. The widespread atmospheric importance of carbonaceous aerosol, together with its extreme complexity, and the experimental as well as theoretical difficulties associated with the characterization of its properties and atmospheric effects have all made it an ideal subject of contemporary aerosol research. Although some of the aspects of carbonaceous aerosol research have been explored for several decades, the general recognition of its global importance dates back only to the early 1990s. It was when carbonaceous aerosol was promoted from being seen as a pollutant of local or regional importance to an atmospheric component of global significance that the real breakthrough occurred. If one looks back now, one can readily see that by that time vast amounts of evidence to support the abovementioned view had already been in place in every major field of modern aerosol research. However, a “change of paradigm” was needed to bring these into perspective. This “revelation” was probably inspired by a paper by Novakov and Penner that appeared in Nature in 1993. Specifically, it was concerned with the recognition that organic aerosol may play a role in cloud nucleation, but it is more likely that by that time the aerosol community had already been ripe for the change and had been awaiting some great idea to break through. This simplified concept of 1

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the “renaissance” of modern carbonaceous research is clearly subjective and there may be several arguments against it. However, it is beyond any doubt that the number of such articles skyrocketed just after that particular paper had been published in 1993. But apart from the issue of exactly when and why carbonaceous aerosol came into the focus of atmospheric research, it is clear that this field—as we see it now—does not have a long history. One of the most important consequences is that its “information field” is very inhomogeneous: some aspects have been thoroughly explored and have accumulated an enormous suite of data; at the other end of the scale others which are relatively new and resource-intensive are characterized by an extreme paucity of data. Above all, there is generally a lack of solidified terminology and methodology even in the more traditional areas of carbonaceous aerosol research. These all make plenty of room for current investigations but, for the same reason, writing a “consolidated overview” of the field becomes excessively difficult. Let us first elaborate on the concept of “carbonaceous aerosol” itself. The term is now placed in quotation marks to indicate that it is in fact a fiction: over most of the troposphere carbon compounds occur in association with inorganic aerosol constituents; pure carbonaceous particles can barely be found. When we refer to “carbonaceous aerosol” in general, it should correctly refer to the “carbonaceous component of atmospheric aerosol”. Only with this in mind are we allowed to use the common term without quotation marks. Another simplifying axiom of carbonaceous aerosol terminology is the distinction between elemental and organic carbon. Whereas the distinct atmospheric effects of these generic classes may justify such a division, there is no clear borderline between the two. Conceptually, there is a smooth transition between organic and elemental carbon, so any division can only be operationally defined and arbitrary. Taking into account the large number of methods that are available for the purpose, it is not surprising that they introduce a large uncertainty into the determination of these basic forms of carbonaceous aerosol. Organic aerosol can also be approached from the molecular level, which brings thousands of organic compounds into our focus. This is a traditional concept which has developed in parallel with modern organic analytical techniques, and yielded enormous sets of data, especially for urban and source aerosol. However, this approach has failed to account for most of the organic carbon in aerosol, so our understanding of organic aerosol from the molecular level still remains rather incomplete. It follows from the above that in order to keep the balance between the various approaches, we need to restrict the scope of this book to the modern and holistic concept of carbonaceous aerosol. It means that we will primarily

Introduction

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look at carbonaceous aerosol from a global perspective, thus sacrificing many aspects that are more restricted to a local scale. Information on certain classes of compounds will be weighted on the basis of their significance in global atmospheric effects, which is largely—but not always— proportional to their contribution to the aerosol mass. Trace organic compounds will only be considered to the extent they can convey valuable information on carbonaceous aerosol sources. This restriction inevitably means that important toxic or carcinogenic compound classes of atmospheric aerosol will be ignored. It should be noted, however, that such pollutants are generally treated in depth in the field of environmental (analytical) chemistry, where aerosol particles are merely seen as a transport medium for these compounds. Even with the restrictions above, the field of carbonaceous aerosol is so wide and it is developing so rapidly that it cannot be covered comprehensively. Its coverage can only be structured along the most popular trends in its research, which have been—unfortunately—rather independent from one another. On the other hand, atmospheric modeling would require consistent and reliable data on carbonaceous aerosol with a sufficient spatial and temporal coverage, which clearly does not exist. The demand set by global atmospheric models, however, seems to be partly satisfied by recent large-scale atmospheric chemistry projects, which apply carefully designed protocols for the sampling, analysis, and evaluation of carbonaceous aerosol, with internal quality control. The methods used are based on scientifically sound principles that have been thoroughly tested and are understood to be the best available. Although limited in space and time, such projects represent the first step towards the better characterization of the global significance of carbonaceous aerosol. The diversity of the field as well as the abstract nature of carbonaceous aerosol both make it very difficult to set the level of any book on this topic. One approach may be to integrate this broad field into the more general topic of atmospheric aerosol chemistry, which would in fact be the most consistent with the atmospheric occurrence of carbonaceous aerosol components. However, keeping the balance would leave very little room for the discussion of the chemistry and effects of carbonaceous aerosol, much less than it would actually deserve. Until now, only a handful of book chapters and review papers have been devoted specifically to carbonaceous aerosol in spite of the fact that their global atmospheric significance has been recognized for about a decade. Among them, there is a compilation of the presentations from a conference on carbonaceous aerosol in 1982, entitled as Particulate Carbon—Atmospheric Life Cycle [Wolff, G. T. and Klimisch R. L. (eds.), Plenum Press, New York, 1982], and a volume only dealing with one major component of carbonaceous

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aerosol: black carbon [Goldberg, E. R. Black Carbon in the Environment, John Wiley, New York, 1985]. More recently, a book chapter was published under the title Carbonaceous Aerosols Influencing Atmospheric Radiation: Black and Organic Carbon by Joyce E. Penner [In: Charlson, R. J. and Heintzenberg, J. (Eds.) Aerosol Forcing of Climate, John Wiley and Sons, Chicester, pp. 91– 108, 1995]; another one entitled Carbonaceous Combustion Aerosols by Helène Cachier [In: Harrison, R. M., van Grieken, R. (Eds.) Atmospheric Particles, John Wiley, New York, pp.295–348, 1998] on an important subset of atmospheric carbon particulates; the most detailed and comprehensive coverage of carbonaceous aerosol available to date is found in the chapter entitled Organic Atmospheric Aerosol by Spyros Pandis and John H. Seinfeld [In: Seinfeld, J. H. and Pandis, S. N. Atmospheric Chemistry and Physics, from Air Pollution to Climate Change, John Wiley, New York, pp. 700–765, 1998]; and a separate chapter entitled Carbonaceous Particles by Ernő Mészáros [In: Mészáros, E. Atmospheric Aerosol Chemistry, Akadémiai Kiadó, Budapest, pp. 167–192, 1999]. In addition to these separate chapters, there are only a handful of review papers on carbonaceous aerosol [Duce, 1983; Jaenicke, 1978; Jacobson et al., 2000; Turpin et al., 2000; Seinfeld and Pankow, 2003]. Apart from these sporadic occurrences, however, several fundamental aspects of carbonaceous aerosol are widely scattered throughout the literature and therefore remain difficult to identify. This book is the first attempt to condense all major aspects of carbonaceous aerosol chemistry and its atmospheric effects into a single volume. However, the organization of all major aspects of carbonaceous aerosol into individual chapters and their integration into a coherent structure caused me a lot of headache, especially in the early stages of the work. The current structure is the result of several painful compromises. Apart from the Introduction and References, all information is contained in only four main chapters. Of these, it is only Chapter 4 (Organic Chemistry of Aerosol) that can be regarded as a conventional book chapter. The other three are “virtual” or “organizing” chapters, which means that their first-order sub-chapters should have been main chapters if I had found any logical structure to support them. Unfortunately, I was unable to find any structure which would have done so. Chapter 2 (Methods of Observations of Carbonaceous Aerosol) brings sampling and measurements pertinent to carbonaceous aerosol under a single umbrella. This integration may be justifiable since sampling and measurements are usually closely related, and are both experimentallyoriented. Chapter 3 (Major Carbonaceous Particle Types and their Sources) is sectioned along the traditional source-based approach which separates soot (EC/BC), as well as primary and secondary organic aerosol by their distinct

Introduction

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characteristics and/or mechanisms of formation. While soot particles have a very specific chemical composition that is addressed separately in the same sub-chapter, no chemical distinction can be made between primary and secondary organic aerosol either in bulk or on the molecular level. This necessitates a separate chapter to be devoted to the organic chemistry of aerosol (Chapter 4). Similarly, it would be neither possible nor logical to assign various physical properties and the resulting global atmospheric effects to the principal carbonaceous particle types. The most important properties (optical, hygroscopic and surface reactivity) of carbonaceous aerosol and their climateand atmospheric chemistry-oriented effects are addressed separately in three sub-chapters, and are brought together in Chapter 5 (Properties of Carbonaceous Aerosol and their Role in the Global Atmosphere). Within each chapter, the emphasis is on building up a structure which is capable of accommodating most of the information published. The chapters are neither intended to reach the depth of a regular review paper, nor to include all publications supporting each statement. This also means that the reference list is far from being comprehensive, though all of the statements are supported by high-standard references. Since the chapters and subchapters are rather loosely connected, a more specific introduction is provided at the beginning of each chapter and first-order sub-chapter.

Chapter 2 METHODS OF OBSERVATION OF CARBONACEOUS AEROSOL 2.1 Sampling of Carbonaceous Aerosol Sampling is an indispensable and quite often critical step for the chemical characterization of atmospheric aerosol particles. Except for some on-line techniques which are capable of measuring the chemical composition of individual particles, most analytical methods require prior collection of the particles on a substrate. The main objective of sampling is to collect sufficient amount of particulate matter from relatively large volumes of air which can satisfy the demands of the analytical techniques. Reliable sampling methods have long been established for atmospheric aerosol. They can be basically classified into two broad classes, filter-based and impactor sampling. There is actually no clearcut delineation between the two classes, for example in dichotomous samplers (virtual impactors) actually filters are used for particle collection. There are some inherent problems with either method which are not specific to carbonaceous aerosol. In particular, size classification is prone to errors and uncertainties. These problems have been amply discussed in the literature, and will not be considered here. This chapter will be devoted entirely to the sampling artifacts and problems which are specific only to organic aerosol, though in some instances making reference to the analogy with sampling of certain inorganic species is difficult to avoid. There are basically two sources of artifacts in organic aerosol sampling which are not fully independent. The first is that a considerable fraction of organic aerosol consists of semi-volatile species which are in dynamic equilibrium between the gas and particulate phase. During sampling this equilibrium is inevitably disturbed, and consequently, semi-volatile compounds can either adsorb to or desorb from the particulates depending on the changes in environmental conditions, such as temperature, mixing ratio or relative humidity. As a matter of course, the same compounds are also in dynamic equilibrium in the atmosphere, and their distribution between the gas and the aerosol phase responds instantly to changes in ambient conditions. Sample collection, on the other hand, takes place for a prolonged period of time, 7

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during which substantial changes in environmental parameters can be experienced. This means that the particles collected on a substrate are subjected to changing conditions, to which the semi-volatile species will continuously adjust. As a result the equilibrium will be strongly skewed towards ambient conditions prevalent near the end of the sampling. Although this is a major concern regarding sampling of organic aerosol, and in particular semi-volatile organic species, no correction is available to date for this effect in any of the sampling procedures. It should be noted that certain inorganic species, such as ammonium-nitrate is also semi-volatile, and the problem discussed here also pertains to them. The other source—which is more specific to organic compounds—is related to the filter substrate itself which is needed for subsequent chemical analyses for total carbon. Usually high-purity quartz is used for this purpose which has a large specific surface area and high adsorptivity toward many gas-phase organic species. This is highly critical since the mass concentration of gaseous organic carbon can be eight times as much as that of particulate organic carbon [Mader et al., 2001]. Unless removed prior to particle collection, the adsorption of even a small amount of gaseous species can lead to measurable positive artifact with respect to particulate phase organic carbon. The purpose of this detailed discussion on the sampling artifacts and more reliable sampling methods for organic aerosol is twofold. First, it is intended to highlight the pitfalls and caveats of sampling of organic aerosol with established methods which work quite well for other non-volatile aerosol species. Secondly, for existing measurements, it may help estimate the uncertainty (or perhaps the systematic error) in many data on organic aerosol that have been obtained without even the simplest method of artifact correction. It is especially important because such data are relied on in the validation of atmospheric models, which are in turn used for the assessment of anthropogenic influence on the atmosphere and climate. 2.1.1 Filter Sampling 2.1.1.1 Filter Medium Typically high-purity quartz is the preferred filtration medium for organic aerosol, especially for subsequent determination of total carbon by thermal analysis. This is because quartz fiber filters do not require any organic binder to provide the rigidity needed to withstand the pressure drop during sampling. The density of a typical quartz fiber filter is around 6.5 mg cm−2, its surface is neutral, and its particle collection efficiency (above particle

Sampling

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diameter of 30 nm) is close to 100 %. However, the BET surface area of a quartz fiber filter is 126 ± 9 m2 m−2, and the surface is not chemically inert, which makes such filter a relatively good adsorbent [Turpin et al., 1994]. For this reason the determination of atmospheric particulate carbon concentrations from the total carbon load measured on quartz fiber filters is not as straightforward as it may seem [Storey et al., 1995; Mader et al., 2001]. It is a matter of debate whether quartz filters have to be pre-baked at high temperatures (typically at 500–700 °C) in order to reduce organic blanks or not. Since commercially available quartz filters are in adsorption equilibrium with ambient air in terms of semi-volatile organic species and water, it follows that they yield a sizable carbon blank upon thermal analysis. If they are pre-baked prior to sampling at high temperatures and precautions are taken during storage and transport to avoid re-adsorption of gas-phase phase species, the blanks can be reduced considerably, often below detectable levels. However, during sampling the equilibrium will be eventually reestablished by the adsorption of volatile or semi-volatile organic species from the gas-phase. This leads to severe positive artifacts without blank correction or if only the pre-baked quartz filter is subtracted as blank. 2.1.1.2 Sampling Artifacts The sampling artifacts pertinent to organic aerosol sampling are generally classified by the sign of the error they cause relative to the particulate phase concentrations. Thus, positive artifact (also known as adsorption artifact), and negative artifact (also known as volatilization or evaporation artifact) can be distinguished, causing over- and underestimation of particulate phase concentration of organic carbon, and also of semi-volatile organic species, respectively. There is another type of organic sampling artifact, namely the reaction artifact, which results from the reaction of organic species on the filter substrate with reactive trace gases and radicals passing through the filter. It can be either positive or negative, depending on whether it produces less volatile or more volatile species, respectively. It is generally not known for certainty which artifacts predominate under certain sampling conditions, as to date no single method can quantify all artifacts at the same time. The magnitude of these artifacts is dependent on many variables, including sampling time, flow rate, ambient temperature and relative humidity, gas- and particulate phase concentration of semi-volatile organic compounds as well as of reactive trace gases. What is known for certainty is that if unaccounted for, either artifact can result in severe systematic errors in organic concentration measurements.

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The various sampling artifacts which are relevant for total particulate organic carbon concentrations as discussed above do not apply equally to all individual organic species. The magnitude and significance of these effects depend on the volatility of organic species and their interaction with the quartz filter medium. There exists a relatively narrow intermediate volatility range within each class of organic compounds, for which the finite gasparticle distribution gives rise to measurable sampling artifacts. As a matter of course, it depends on the ambient temperatures. For example, within the homologue series of n-alkanes, only homologues between C21 and C24 belong to this range in mid-latitude winter, whereas it is shifted to C22–C26 in summer. For monocarboxylic acids, the range is from lauric to stearic acid in summer, and to palmitic acid in winter. For polycyclic aromatic hydrocarbons (PAHs), this translates into molecular weight ranges of 178–228 and 202–252 in winter and summer, respectively [van Vaeck et al., 1984]. Within these ranges the distribution of the homologues, which are frequently used as fingerprints of various aerosol sources, become rather skewed due to the sampling artifacts. In addition, recent studies suggested that volatility is not the only parameter governing sampling artifacts. A considerable fraction (13–30 %) of the highly polar dicarboxylic acids can also be found on quartz backup filters, with the exception of malonic acid [Limbeck et al., 2001]. Apart from the possibility that these compounds may form by the photooxidation of unsaturated monocarboxylic acids adsorbed on the filter substrate [Kawamura and Gagosian, 1987], it seems that gas-particle partitioning is not solely dependent on vapor pressure, but other factors, such as absorption into the liquid phase may also play a role, as suggested by Pankow [1994]. a) Positive (adsorption) artifact As it has been pointed out in sub-chapter 2.1.1.1, the total carbon concentration measured on quartz fiber filter also comprises adsorbed gaseous-phase species due to the high adsorptive potential of pure quartz. This is a positive sampling artifact, which leads to the overestimation of organic aerosol mass concentration. Since the adsorptive capacity of the quartz fiber filter is limited, the amount of organic species adsorbed approaches saturation. In other terms, the surface of quartz is equilibrated with gaseous organic species during sampling, but it must be recalled that equilibrium is not attained instantly. As a quartz fiber filter cannot adsorb more gaseous organic species beyond saturation, the relative magnitude of the positive artifact gradually decreases with increasing particulate load on the filter.

Sampling

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Therefore positive artifacts are important for fresh (pre-baked) filter medium, for short sampling times and/or low particulate organic mass concentrations [Appel et al., 1989]. For example, aerosol sampling aboard an aircraft can be loaded with serious positive artifacts. Under unfavorable conditions, its magnitude can reach 100 % [Kirchstetter et al., 2001]. This is unacceptably high so the concentration of carbon measured on a single quartz filter without correction cannot be a measure of particulate carbon under these circumstances. On top of this, the apparent concentration of particulate organic carbon collected on quartz fiber filters seems to depend on face velocity (the volumetric flow rate divided by the exposed area of the filter) in the range of 10–120 cm s−1 [McDow and Huntzicker, 1990]. If the face velocity was increased at constant volumetric flow rate, significant decrease was observed in the apparent concentration of particulate carbon on both the front and backup quartz filters. On the contrary, for longer sampling times at urban locations, positive artifacts can be negligible [Turpin et al., 1994]. Prolonged sampling, however, could introduce other artifacts such as negative artifacts or reaction artifacts. Correction methods for positive sampling artifact—blank correction During short airborne sampling, the magnitude of the positive artifact could be as high as a factor of 3–18, and in most cases the amount of total carbon measured on single quartz fiber filters even exceeds the total aerosol mass determined gravimetrically on Teflon filters [Novakov et al., 1997b]. If, however, blank filters were allowed to be in contact with ambient air for sufficient time to establish equilibrium, preferably at the sampling site during sampling, this positive artifact can be greatly reduced. Conventional field blanks may not be fit for this purpose. In filter sampling, field blanks are prepared by placing a pre-baked quartz filter into the sample holder, then sampling air for a very short period of time (typically for a few minutes). Except for highly polluted environments, this procedure is clearly insufficient to bring the filter medium in equilibrium with ambient air. Pre-baked quartz filters may not become saturated during sampling even for several hours [Kirchstetter et al., 2001]. Therefore total carbon concentrations that have been obtained using a pre-baked single quartz filter with field blank correction, will certainly overestimate total particulate carbon concentration to a variable extent, depending on sampling time and conditions. If no field blank correction is applied, this overestimation would be even more severe. Allowing for equilibration of pre-baked quartz filters at the sampling site over the entire period of sampling with precautions against particle deposition on the filters, and using these filters as blanks to be subtracted

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Carbonaceous Aerosol

from identical filters that have been exposed is possibly the simplest method to approximate particulate carbon concentrations as they exist in the atmosphere. Correction methods for positive sampling artifact—Dual filter strategy Another relatively simple and well-established method to correct for positive artifact can be to use two quartz fiber filters for sampling placed directly on top of each other (tandem or dual filter strategy). This method, first suggested by Fitz [1990], does not require instrumental modification of commercially available high-volume or low-volume samplers and impactors, so it has been a frequently used method for artifact correction. However, it increases pressure drop—which might lead to negative artifact— and more importantly, doubles the number of samples and subsequent thermal analysis runs. Perhaps this is the main reason why most studies involving organic aerosol did not report the use of even this simple method of artifact correction. The dual filter strategy relies on the presumption that the front filter collects all particulates plus becomes saturated with adsorbed gaseous organic compounds, whereas the backup filter collects only gas-phase species up to saturation. In other terms, both the front and backup filters are assumed to have attained perfect equilibrium with the gas-phase, and—being identical— they have adsorbed the same type and amount of semi-volatile organic species. If these assumptions are valid, the concentration of particulate carbon can be calculated by simply subtracting the total carbon concentration measured on the backup filter from that of the front filter. Now let us consider critically each of the presumptions behind the dual filter strategy. Although not stated explicitly, it is generally assumed that the collected particulates do not adsorb significant amounts of gas-phase organic compounds. As a first measure, the area of the filter can be related to the calculated surface area of the particles, assuming a typical size-distribution of spherical particles. It turns out that while the BET surface area of quartz fiber filters is 126 ± 9 m2 m−2, the total surface area of particulates is in the order of 100 cm−2 which adds little to the surface area of the filter [Turpin et al., 1994]. In addition, since the collected particulates are in equilibrium with the gas phase upon sampling, excess adsorption of gaseous compounds can occur only when atmospheric conditions change in favor of adsorption. Both factors imply that the presence of particulate matter cannot significantly affect the positive artifact. Instead, the presence of particles can cause volatilization losses, as will be discussed in connection with negative artifacts. Are the quartz filters of the same type really identical? Surprisingly, it was found that sampling artifacts were also affected by different lots of the

Sampling

13

same type of quartz fiber filters having significantly different adsorption capacities in spite of their identical appearance [Kirschtetter et al., 2001]. This effect—which is very difficult to avoid except when precautions are taken well before sampling—can be as high as a factor of 3 in terms of the corrected organic carbon concentrations. The issue whether both filters can reach saturation by the end of sampling or not at any location seems to be also dependent on the duration of sampling. Kirschtetter et al. [2001] run four quartz filter pairs over a period of 14 hours to study this effect. One sample covered the whole period of sampling whereas the other three were collected for subsequent segments of this period, for 2, 4 and 8 hours. The sum of the corrected total carbon load of the latter three samples exceeded that of the sample collected over the entire period by more than 50 %. If the simple artifact correction had been valid the two values should have been the same. For shorter sampling times the backup quartz filter was found to adsorb less gaseous organic compounds than the front filter, therefore simple subtraction did not correct fully for the adsorption artifact. Regarding the chemical identity of adsorbed species, qualitative and quantitative comparison of thermograms of the front and backup quartz fiber filters sampled in an aircraft revealed interesting features [Novakov et al., 1997b]. Detailed description of the thermal method can be found in subchapter 2.2. It was found that the lowest temperature peak (< 200 °C) was very similar in shape and intensity in the thermograms of both filters, indicating that these compounds likely result from adsorption of volatile organic species, most likely as a positive sampling artifact. A peak centered at about 250 °C was on both filters only at low total carbon concentrations otherwise it was only prevalent on the backup filter. It was hypothesized that less volatile polar organic species adsorbed from the gas phase on the active sites of the quartz filters which were more numerous on the backup filter. The question remains why these polar species—which are expected to bind strongly to the active polar sites of the quartz filter—were displaced by more volatile and possibly non-polar species when the load was higher. The dual filter strategy also seems to eliminate most of the face-velocity dependence of the apparent concentration of particulate organic carbon collected on quartz fiber filters [McDow and Huntzicker, 1990]. However, it was speculated that due to the larger pressure drop on the second filter less vapor adsorbs on the backup filter than on the front one [Huebert and Charlson, 2000]. Unless a decrease in temperature do not compensate for the difference, the use of the back filter to correct for the “positive artifact” on the front filter would actually underestimate its magnitude and overestimate the mass concentration of particulate-phase organic species. It is not possible to assess this bias for lack of data on temperature and pressure changes within

14

Carbonaceous Aerosol

the filter matrix, which causes that the uncertainty of organic sampling remains in the order of 30–50 % even with the dual filter strategy. Correction methods for positive sampling artifact—Quartz-quartz (QQ) and Teflon-quartz(TQ) filter pairs Because of the shortcomings of the dual filter strategy a more sophisticated, albeit more accurate sampling method was developed for artifact correction. In this setup sampling is performed in parallel on two quartz filters in one line and a quartz filter behind a Teflon filter in the other. The flow rate and face velocity in both lines should be the identical. The reason for using two parallel lines is that Teflon filters are not amenable to thermal analysis above 275 °C due to thermal breakdown. Teflon filters have surface area several times smaller than that of quartz filters, and unlike quartz, they are chemically inert. Since the Teflon filter was assumed to have greater than 99.999 % collection efficiency for all particles above 10 nm and negligible adsorptive capacity for most volatile organic compounds, the backup quartz fiber filter only adsorbs organic species from the gas phase. In other terms, the Teflon front filter has negligible effect on any adsorption process occurring downstream nor it is a source of contamination for the backup filter downstream, thus refuting the conclusions of Appel et al. [1989]. In this setup significantly more organic carbon can be found on the backup quartz fiber filters behind Teflon filters than on backup quartz fiber filters behind quartz fiber filters [McDow and Huntzicker, 1990]. The mean difference can be as much as 60 % when sampling at a rural site for 12 hours [Turpin et al., 1994]. For sampling times of 24 hours positive artifacts may exceed 50 % at low particulate organic carbon concentrations, but remain below 20 % above about 7 µgC m−3 [McDow and Huntzicker, 1990]. If adsorption equilibrium is attained quickly on both Teflon and quartz front filters then adsorption on both backup filters should be the same. On the contrary, the results imply that equilibrium was not attained even for several hours on quartz fiber filters, but was established quickly on Teflon filters. For this reason it was suggested that organic carbon concentration measured on a backup quartz fiber filter placed behind a Teflon filter is a more accurate measure of the positive artifact. Detailed examination of the carbon thermograms of both types of backup filters supported this conclusion [Kirchstetter et al., 2001]. The organic carbon load on the front quartz fiber filter (10.3 µg m−2) can be reduced to 6.0 µg m−2 after correction with the second quartz fiber filter, whereas the above approach reduces the estimate further to 5.3 µg m−2. Such a low value cannot be attained when

Sampling

15

using a denuder upstream of the quartz filter, since no denuder has 100 % collection efficiency. This implies that though a denuder might promote volatilization losses from collected particles by depleting gas-phase concentrations, the effect may not be significant. More can be learned about the nature of sampling artifacts from the thermograms of front and backup quartz fiber filters collected in the quartzquartz and Teflon-quartz filter-pair arrangement at a marine location, as shown in Figure 2.1 a and b, respectively [Novakov et al., 1997a].

Figure 2.1. Thermograms of (a) front and (b) back 24-hour filter samples collected March 15– 16, 1995, at Cape San Juan. Carbon concentration of peaks A, B and C are determined from the corresponding areas defined by the vertical lines. Dashed lines are thermograms of field blanks (after Novakov et al. [1997a]).

16

Carbonaceous Aerosol

The thermogram of the front filter showed three peaks labeled A, B, and C appearing in the order of increasing thermal stability, whereas that of the backup filters exhibited only two peaks, closely corresponding to the first two peaks seen in the thermogram of the front filter. Peak A, with similar magnitude on both the front and the backup filters was attributed to species adsorbed on the filter substrate up to saturation. The interpretation of peak B is much more speculative, as it was more pronounced on the backup filter, indicating that the species involved were only partially retained on the front filter but they were more efficiently collected on the backup filter. Since the magnitude of this peak on quartz filters placed behind Teflon filters was systematically higher than on those behind another quartz fiber filter, the most probable explanation was the adsorption of polar organic species on the active sites of the quartz fiber filter, which were more numerous on the backup filter. The straightforward assignment of peak C was that filterable particulate organic carbon gave rise to it. However, the incomplete resolution of the peak made its quantification somewhat uncertain. The average concentration of filterable particulate organic carbon was found to be 391 ± 106 ng m−3 which can be contrasted with the total organic carbon of 1298 ± 510 ng m−3 determined on the front quartz fiber filter. This finding clearly indicated that 24-hour sampling in such pristine environment with a single quartz fiber filter was clearly inadequate to represent aerosol carbon concentrations. Secondly, it also suggested that even the filter-pair sampling method was unable to fully correct for positive sampling artifact, though it is still probably one of the best methods to assess the magnitude of positive artifacts. Thirdly, the results implied that only a combination of sampling and thermal measurement methods has the potential to determine the “true” concentration of particulate carbon for short sampling times at pristine locations. b) Negative artifact As opposed to the positive sampling artifact, which involves adsorption of semi-volatile organic substances onto the surface of quartz fiber filter, largely the same semi-volatile species may also desorb from the collected particulate phase during the course of sampling—especially when ambient conditions change in favor of desorption. The resulting loss of these compounds is called negative artifact (also referred to as evaporation or volatilization artifact) in terms of the particulate phase concentration of organic carbon as well as of semi-volatile species, leading to possible underestimation of their mass. In general, negative artifacts are less significant than positive artifacts, due to the much lower total surface area of particulates on one hand, and to

Sampling

17

the fact that the semi-volatile compounds on the surface of the particles are ab ovo in equilibrium with air on the other. Only substantial changes in equilibrium conditions—such as a rise in temperature or the ongoing sampling of less polluted air—could result in desorption of semi-volatile species which have been originally in the particulate phase. The parameter that affects volatilization artifact most is temperature. A change of 10 °C in temperature typically induces a threefold change in the partitioning coefficient [Mader et al., 2003]. To minimize losses, an increase in temperature during sampling should be avoided, if possible. Consequently, negative artifact can be particularly significant in aircraft sampling when the sampling device is typically located inside the heated cabin and temperature difference between ambient and inside air can reach as high as 60 °C at high altitudes [Mader et al., 2001]. In addition, substantial negative artifact can occur when gas-phase species are removed upstream of the quartz fiber filter using a denuder in the sampling line to minimize positive artifacts. When gaseous semi-volatile organic compounds are removed by the denuder, the same compounds originally absorbed by the particulate phase will re-evaporate from the particles in an attempt to re-establish gas-particle equilibrium. This can occur already during transport of particles through the denuder [Kamens and Coe, 1997], and when the denuded air passes through the particle-laden filter [Liang and Pankow, 1996]. Additionally, the pressure drop across the filter may also induce subsaturation and evaporative losses of volatile species. This negative artifact depends on the volume of denuded air passing through the filter, the gas-particle partition coefficients of the semi-volatile compounds, and the amount of particulate carbon. If a backup quartz fiber filter is applied together with a denuder upstream, the organic carbon found on the backup filter is normally regarded as evaporated from the particles. Its magnitude, however, is difficult to assess, since—as pointed out above—the denuder never has a 100 % collection efficiency. In spite of this, XAD-coated denuders effectively remove the semi-volatile compounds that would adsorb to quartz thereby minimizing positive artifacts [Mader et al., 2002]. As much as 30 % of the particulate organic carbon could evaporate during sampling with a denuder-filter sampler [Mader et al., 2001]. The thermogram of a denuded continental aerosol sample was found to be quantitatively similar to that obtained by subtracting the thermogram of a quartz fiber filter behind a Teflon filter from that of a front quartz fiber filter [Kirchstetter et al., 2000]. In the thermogram of marine aerosol, however, there was no corresponding peak centered at about 250 °C on the front quartz fiber filter, which was pronounced on the quartz fiber filter behind Teflon filter. In the denuded samples only a small fraction (14 %) of the organic

18

Carbonaceous Aerosol

carbon remained on the quartz fiber filter, indicating that the denuder effectively removed not only the most volatile fraction but nearly all organic compounds. When interpreting aerosol thermograms in terms of carbon volatility, we should recall that liquid-phase adsorption on quartz seems to shift the evolution temperature toward higher values compared to adsorption from the gas phase [Kirchstetter et al., 2000]. With this in mind, it can be surmised that the condensed phase species giving rise to the peak at the rear in the thermogram of the front quartz filter evaporated from the Teflon filter then condensed on the backup quartz fiber filter from the gas phase, and evolved at a lower temperature. c) Reaction artifact Oxidation of the organic species deposited on the filter by strong oxidants can occur during prolonged sampling [Pio et al., 2001a]. This can either be positive or negative artifact depending on whether an organic compound is formed or consumed in the reaction. In terms of total organic carbon, the sign of this artifact depends on whether the reaction products are more or less volatile than their precursors. Currently very little is known about this type of artifact, which can be quite significant for certain compound classes, such as polycyclic aromatic hydrocarbons (PAHs). 2.1.1.3 Simultaneous Characterization of Positive and Negative artifacts— Denuder-filter Combinations Denuder-filter combinations are generally used in sampling systems designed to quantify both positive and negative artifacts at the same time [Eatough et al., 1993]. An example of such an apparatus is the multiple sampling system, multichannel parallel plate diffusion denuder sampler, the BYU Organic Sampling System (BOSS). The system consists of three samplers, attached to a common Teflon-coated inlet manifold after a cyclone with 2.5 µm nominal particle size cut. The scheme of the sampling system is shown in Figure 2.2.

Sampling

19

Figure 2.2. Schematic of the BOSS sampling system. Sampler 1 (denuder), sampler 2 (filter/denuder) and sampler 3 (filter pack) (after Eatough et al. [1993]).

Sampler 1 consists of a diffusion denuder, followed by two quartz fiber filters and a charcoal impregnated filter. The gaseous compounds are removed in the denuder, whereas the semi-volatile compounds that volatilize off the particulates or break through the denuder are adsorbed on the quartz fiber filters and the charcoal. The denuder is not expected to induce significant particle losses in terms of aerosol mass, since only the particles < 0.1 µm can be lost on its walls. In Sampler 2 the quartz fiber filters precede the denuder, which collects gas-phase compounds, including those volatilizing off the particles. At the end a charcoal impregnated filter is used to determine denuder breakthrough. Sampler 3 has only two components, a quartz fiber filter pair in front and a charcoal impregnated filter as backup. The multichannel parallel plate denuder absorption efficiency is about 87 ± 2 %, and particle losses inside are assumed to be negligible. The quantities measured are the particulate carbon collected directly on the front quartz fiber filter, the amount of carbon on the second quartz fiber filter (called quartz artifact, without a statement related to its sign), and that lost from the particles during sampling. As regards quartz artifact, the amount of carbon in the filter pack (sampler 2 or 3) can be compared with that found on the second quartz fiber filter after the denuder. These two amounts were found to be almost identical, irrespective of the passing of the air sampled through the denuder. This finding was interpreted that it resulted from the volatilization of semi-volatile compounds from the particles during sampling, i.e. a negative sampling artifact. In other terms, it has to be added back in to yield ambient particulate carbon concentrations. This is in contradiction with the assumption behind the dual filter strategy, which deems quartz artifact as a

20

Carbonaceous Aerosol

positive one which has to be subtracted. This concept, however, has not received widespread recognition. It is generally understood that the quartz artifact is largely positive, resulting from the adsorption of gas-phase organic species, and has to be subtracted from the amount of carbon on the front filter to estimate ambient particle phase concentrations. Mader et al. [2001] recently suggested a high-volume particle trap impactor/coated honeycomb denuder sampler for airborne sampling of carbonaceous aerosol. A particle trap impactor is a type of virtual impactor in which particles are impacted into a cavity having a depth to width ratio >>1. Such impactors do not require grease or oil to provide a sharp size cutoff with a maximum collection efficiency which is also constant at high particle loadings. The problem with conventional inlets such as plate inertial impactors or cyclones is that they are subject to particle bounce and reentrainment at high particle loadings unless they are coated with greases or oils—however such coatings would severely confound subsequent organic analysis. On the other hand, a honeycomb denuder has a much greater density of channels per a unit cross-section than parallel plate or annular denuders therefore it can be operated at high flow rates without loss of collection efficiency. The denuder is coated with XAD-4 stationary phase. The low residence time of the particles in the denuder minimizes evaporation losses of semi-volatile organic compounds during transport. The design of the particle trap impactor/honeycomb denuder sampler is shown in Figure 2.3.

Figure 2.3. Particle trap impactor/denuder sampler (after Mader et al. [2001]).

Sampling

21

2.1.2 Impactor Sampling Impactor sampling of organic aerosol is generally free of adsorption artifacts, as commonly used substrates such as aluminum foils do not have adsorption capacity for gaseous organic species. There are, however, severe artifacts in impactor sampling, especially in the size classification of organic aerosol. Originally these artifacts, namely particle bounce and re-entrainment at high particle loadings, are not specific to organic particles. By applying grease or oil on impactor substrates these artifacts can be greatly reduced and collection efficiency can be improved, especially for larger particles. As noted above, what makes these artifacts specific to sampling of organic aerosol is the fact that such coatings cannot be applied if aerosol samples are collected for subsequent organic analysis. The collection efficiency of ungreased plate impactors was shown to decrease dramatically for larger particles due to particle bounce [Pak et al., 1992]. This makes the determination of size distribution of organic aerosol by plate impactors highly biased towards small particle sizes. Additionally, the large pressure drop in the lower stages of the impactor—unless accompanied by significant adiabatic cooling—may induce a large negative artifact, evaporation of particulate-phase semi-volatile species. Combined with the particle bounce this would result in quite substantial losses since larger particles may also be affected.

22

Carbonaceous Aerosol

2.2 Methods for Determination of Main Carbonaceous Aerosol Components The determination of total carbon (TC) in aerosol collected on carbonfree filter substrates or impactor plates is straightforward. It is usually done by Evolved Gas Analysis (EGA), which—when combined with programmed temperature ramping—also provides complementary information regarding the thermal-oxidation stability of carbon species. Because of this extra feature, this technique is discussed separately in Chapter 4. The TC measurements usually have good enough sensitivity and accuracy typically associated with most routine analytical measurements. The question to what extent the measured TC is indeed particulate carbon is left for the detailed discussion on sampling artifacts in sub-chapter 2.1. Within TC the division between organic (OC) and elemental carbon (EC) has its inherent uncertainty due to the smooth transition between the two in any of their properties. In relative terms this uncertainty is much more critical for the determination of EC, which generally makes up only a minor fraction of TC. EC is often thought to be the same as BC or soot: several arguments against this assumption are discussed in sub-chapter 3.1. Nevertheless, there is one specific case when such a statement can be operationally justified: if a method of EC and BC determination is calibrated with soot of known mass concentrations. In this case the EC or BC reading will directly give the mass concentration of soot. Of course, ample bias may result if atmospheric soot differs from that used for calibration which is quite probable for particles in most part of the troposphere. The measurements methods for elemental carbon can be divided into sensitive and specific ones. All sensitive methods take advantage of the strong absorption of the visible radiation by elemental carbon1. According to the generally accepted nomenclature, these optical methods give a measure of BC. However, optical methods are non-specific and must be calibrated with soot of known optical properties. A unique, but still non-specific method is the photoacoustic method, which measures directly the absorption of suspended particulates and is therefore free of the problems associated with filter-based methods. Since it is based on absorption of light, it also determines BC. Another subset of less sensitive, but also non-specific methods is capable of measuring total carbon as well using thermal or more recently thermo-optical methods. By definition, they measure EC concentrations. For 1

See sub-chapter 3.1 for discussion on the chemical structure and composition of soot particles.

Measurement Methods

23

this reason, these methods are usually not calibrated with soot, but give the absolute mass of carbon that falls above a certain temperature limit. In simple thermal methods this limit is preset during method development and remains fixed for all measurements. In thermo-optical methods the limit is variable and set by optical control measurements for each sample. These latter methods are now understood to be the most reliable off-line measurement techniques for the OC/EC split. However, there are still several methods in use for EC/BC analysis, some of them are outdated, and most of them are hardly comparable. As Chow and co-workers [1993] put it: “It is often stated that the various carbon analysis methods do not agree. The actual situation is worse than that: there is not even agreement on what the disagreements are.” 2.2.1 Methods for BC Determination 2.2.1.1 Optical Methods a) Smoke shade method The “smoke shade” method was originally used in the U. K. [British Standard, 1964], and later adopted by the World Health Organization (WHO) as well as the Organization for Economic and Cooperative Development (OECD). It consists of the measurement of the darkness of stain by a reflectometer on a Whatman No. 1 filter paper. The sampler collects only particles with diameter < 4 µm. By means of an appropriate calibration curve, the results are expressed as µg m−3 of “standard smoke” [Ball and Hume, 1977]. Later the linear relationship between the light absorption of the filter stain in the smoke shade method and the amount of EC on the filter was demonstrated, by thermal analysis of parallel filter samples [Edwards et al., 1983]. The conclusion was that the “smoke shade” method could be used as a surrogate for EC measurement within about a factor of two. Their measured ratio of EC to “smoke” was 0.13 under less polluted conditions. Other measurements conducted at heavily polluted locations yielded a much higher ratio (0.52) [Penner et al., 1993]. The relationship between reflectance and surface concentration of absorbing species is strongly dependent on the particle load on the filter and the penetration depth into the filter material. Therefore it can be shown that the proportionality factors between BC and smoke are markedly different under polluted and less polluted conditions. Optical measurements aimed at TSP concentrations were basically modified versions of the British smoke shade method, but largely unsuccessful in their

24

Carbonaceous Aerosol

original purpose. When transmittance was measured through sequentially collected deposits on a filter tape, the results were expressed in coefficient of haze (COH) units. An alternative technique with reflectance measurement yielded the reflectance unit of dirt shade (RUDS). b) Aethalometer The instrument that uses an optical technique to measure the concentration of aerosol BC in real time was first described by Hansen et al. [1984]. After the Greek word “αιθαλουν”, which means “to blacken with soot”, the instrument was named aethalometer. The aethalometer operates on the principle of continuously measuring the attenuation of a light beam transmitted through a filter through which air is drawn. Provided that the flow rate is constant, the rate of deposition of BC onto the filter is proportional to its concentration in the aerosol and gives a corresponding rate of increase in optical attenuation. The light source of the instrument is a stabilized incandescent lamp with an effective operational wavelength of 880 nm. The instrument has a dual light path, which means that only a small spot on the filter is exposed to aerosol, the rest is used as the reference. Attenuation through the filter (ATN) is expressed as:

ATN = −100 ln (I I 0 )

(2.1)

where I0 and I are the intensities of the transmitted light through the loaded and blank filters, respectively. The aerosol absorption coefficient (crosssection, σAP) can be derived from the attenuation (ATN) measured by the aethalometer by the equation [Bodhaine, 1995]

σ AP (t ) =

1 A ATN (t ) 1.9 V (t )

(2.2)

where A is the area of the sample spot (cm2), V is the sample volume (m3) and t is the sampling time. The operation of the instrument gave a minimum resolving time of 6 seconds at a BC concentration of 1 µg m−3 up to a maximum loading of 8 µg cm−2. For a 1-h integration time, a detection limit of 4 ng m−3 can be attained [Sharma et al., 2002]. The aethalometer is calibrated by performing quantitative thermal analysis on the quartz filter removed. Although the

Measurement Methods

25

optical absorption cross section of BC particles2 depends markedly on their morphology, a value of 10 m2 g−1 was found compatible with other reported values [Hansen et al., 1984]. At maximum sensitivity the aethalometer can have a time-resolution in the order of a few seconds for aerosol having an absorption coefficient of 10−5 m−1. Subsequent studies under different conditions modified the value of optical absorption cross section of BC and established the manufacturer’s value of 19 m2 g−1 [Gundel et al., 1984a]. When the instrument was commercialized, this value became the standard for BC determination. However, scientific debate continued over the constancy of optical absorption cross section of BC which was renamed to specific attenuation cross section (σ) [Liousse et al., 1993]. This term was defined as

ba = σ (BC )

(2.3)

which directly relates the absorption coefficient of the aerosol (ba, m−1) to the mass concentration of BC (g m−3). The range of σ found in the literature varies over an order of magnitude, between 2 and 25 m2 g−1. The calculation of σ from optical attenuation data requires the determination of BC by an independent, most often thermal or thermo-optical method which in itself is also subject to considerable biases and uncertainties. For this requirement, the use of filter-based attenuation measurements is preferred, which also provide samples for BC determination. In these instances, however, the attenuation characterizes the mixed medium of quartz fibers and aerosol particles, and the σ values will likely be significantly different from the mass absorption efficiencies of suspended BC particles. In a fibrous medium, multiple reflection of light occurs which leads to an enhancement of the optical path. In addition, filter attenuation shows some dependence on relative humidity. The combination of these effects may cause an enhancement of σ, as a measurement artifact, by up to a factor of 3 [Patterson and McMahon, 1984]. Secondly, when BC particles are embedded in a quartz fiber matrix having a refractive index of 1.5, their complex refractive index decreases from m = 2 − i to m = 1.33 − 0.66i. Calculations showed that this also enhances σ over the entire size range by 10–15 % [Liousse et al., 1993]. As a result σ will be much higher than the mass absorption efficiency (αabs) of suspended BC particles [Jennings et al., 1993; Bodhaine, 1995]. Significant reduction of 2

Since the method is calibrated against thermally determined EC, it should give a measure of EC rather than BC. However, optical methods are generally accepted to give a measure of BC, irrespective of the method of calibration.

26

Carbonaceous Aerosol

this effect can be achieved when the filters are immersed in oil of a similar refractive index during the measurement [Ballach et al., 2001]. Liousse et al. [1993] concurrently determined aerosol absorption using an aethalometer and BC mass concentration by a thermal method in different environments. Inherent to the filter-based optical measurements, the measured σ is enhanced relative to that of suspended BC particles, which the authors explicitly admit in their paper. Although the authors stated that the optical measurement method applied did not introduce any variability to measured σ values, a statement obviously understood by many that the correct absolute values were determined. The authors were correct in using the term “specific attenuation cross section, sigma (σ)” in the title and throughout the paper, since it was intended to refer to the absorption of BC on quartz fiber filters. Unfortunately, this fact and especially its limitations have been simply overlooked in many subsequent publications dealing with atmospheric absorption of BC for radiative transfer calculations, in which σ has meant to represent “mass absorption efficiency of BC” and its values have been used accordingly (see sub-chapter 5.1). A more correct method to translate BC mass concentration values, as measured by the aethalometer to atmospheric light absorption would be to use an absorption efficiency of about 10 m2 g−1 for BC aerosol which corresponds to a calibration factor of 1.9 between filter attenuation and atmospheric light absorption [Bodhaine, 1995; Arnott et al., 2003]. The σ values obtained were different for different environments, as shown in Table 2.1 [Liousse et al., 1993]. Table 2.1. Variability of σ values obtained for different regions (data obtained through aethalometer measurements) (after Liousse et al. [1993]). Environmental location Grenoble Paris Savannah areas Western Mediterranean Remote areas

Sigma, σ (m2g−1) 12.0 ± 0.7 12.0 ± 0.7 20 ± 1.3 18 ± 0.9 5 ± 0.5

The relative variations of σ were explained by the differences in morphology and mixing state of BC related to different sources and aging processes. The high value for biomass burning BC, i.e. its enhanced absorptive capacity was assumed to result from the fact that in a cooling plume BC particles rapidly acquire a coating of organic compounds. On the other hand, in remote regions aging processes increase the size of the particles which—according to theoretical considerations—leads to a reduction in σ.

Measurement Methods

27

Upon drawing these conclusions the thermal method—which was used throughout the study as a reference method—was deemed to be accurate and absolute. It is known, however, that simple thermal methods with fixed temperature cut tend to erroneously classify organic compounds as elemental to a variable degree (see sub-chapter 2.2.2). In the light of this information— though most of the above conclusions might also be valid—it can be supposed that most variations in σ actually reflected measurement biases in the thermal BC determinations. c) Particle Soot Absorption Photometer (PSAP) A commercially available instrument measuring light transmission through a filter is the Particle Soot Absorption Photometer (PSAP). It uses Pallflex filters in a dual-path setup, and focuses sample on a small spot (0.5 cm in diameter). It operates on a similar principle as the aethalometer, with the exception that it uses a 565 nm LED for light source as compared to incandescent light source for the aethalometer. The sketch of the instrument is shown in Figure 2.4.

Figure 2.4. Cross-section of the filter setup in the PSAP. The sample is drawn through one of the holes, shown on the left, and the particles are deposited on the filter. Filtered air is drawn through the hole shown on the right for a reference measurement (after Bond et al. [1999a]).

The exposed sample spot is also smaller in the PSAP (0.5 cm versus 1.1 cm in diameter). The combined result of these differences is its much lower detection limit (0.8 ng m−3 for 1-minute integration time). The PSAP relies on

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Carbonaceous Aerosol

the optically diffuse properties of the filter medium itself to minimize sensitivity to forward scattering, and allows particles to become embedded in a fiber to reduce sensitivity to backscattering. As these measures against scattering are far from being perfect, a correction formula (Eq. 2.4) was developed to account for the enhanced light scattering by the filter material as well as side- and backscattering by nonabsorbing material on the filter [Bond et al., 1999a]. Without this correction it was shown that about 2 % of the light scattered by pure non-absorbing aerosol would be interpreted as absorption3 and absorption would be overestimated by as much as 22 %.

σ ap = (σ adj − 0.02 × σ sp ) 1.22 = 0.82 × σ adj − 0.016 × σ sp

(2.4)

where σap and σsp are the absorption and scattering coefficients, respectively, σadj is the uncorrected absorption coefficient.

This formula, which was developed for laboratory test aerosol, has been successfully verified for ambient aerosol [Wex et al., 2002; Bundke et al., 2002]. The specific attenuation coefficient for graphitic carbon, as determined by Raman spectroscopy from the filter has been found to be 14.8 ± 2.7 m2 g−1 for a high-elevation rural site [Mertes et al., 2004]. At high relative humidity the PSAP shows erroneously high light absorption, since the cellulose membrane is very hydrophilic and strongly absorbs water [Arnott et al., 2003]. When RH decreases again, the PSAP reading shoots erroneously negative. This underlines the importance of operating PSAP at a constant low RH for routine measurements. d) Integrating plate method

The integrating plate method was first introduced by Lin et al. [1973]. In this method the transmittance through an aerosol sample collected on Nuclepore filter is measured against a blank filter. The special optical arrangement which integrates the light scattered by the filter and the particles ensures that transmission will be solely due to absorption. Furthermore, it is assumed that the aerosol particles are contained in a column, the height of which (x) is determined by the ratio of the sample volume and the area of the deposit, and that the deposition onto the filter does not alter the optical properties of the suspended particles. The integrating plate method uses an opal glass diffuser to transmit forward-scattered light and minimize the 3

i.e. the apparent single scattering albedo of pure white particles would be 0.98.

Measurement Methods

29

sensitivity of backscattered light by matching the refractive index of the filter material to that of the particles. However, it was shown that if the aerosol contained nonabsorbing particles, either internally or externally mixed, the light absorption coefficients calculated were too high, e.g. by a factor of 3 for aerosol with a single scattering albedo of 0.95 [Horvath, 1995]. e) Integrating sphere method The integrating spheres for aerosol measurements were first introduced by Fischer [1970], but the integrating sphere method for the determination of BC was originally described by Heintzenberg [1982]. In this method the aerosol samples are collected on isopore polycarbonate filters which are soluble in chloroform. The schematic view of an integrating sphere photometer is shown in Figure 2.5.

Figure 2.5. Schematic diagram of the integrating sphere. The dashed line is the mounting for a liquid sample. For the transmission measurements on undissolved filters, the filter is placed in a holder at the entrance port of the light beam. If liquid samples are used, the detector is placed at the bottom of the sphere, that is, at right angles to the incident light. The entrance port for the light is then covered with a second diffusor (after Hitzenberger et al. [1996]).

The interior of the sphere is coated with spectraflect which reflects virtually all of the incident light in a perfectly diffuse way. The light source is a stabilized halogen lamp provided with an opal glass plate and an interference filter with maximum transmission at 550 nm. The detector is a silicon diode linear radiometer. The homogeneous samples can be either filter deposits or liquid suspensions. Filter samples are either placed outside the sphere at the entrance or in the center of the sphere.

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In the first arrangement the sphere acts as a diffuse detector and the change in transmission with respect to blank is taken as caused by absorption assuming negligible scattering losses. In the second arrangement the sample is irradiated with the light directly from the source and with diffusively distributed light within the source. Due to the presence of absorbing substance the signal decreases. The instrument is calibrated with homogenized suspensions of standard soot thus giving an absolute concentration of BC without the need of assuming any specific absorption efficiency. Losses of scattered light and changes in the optical properties of the filters reduce signal even if the sample is non-absorbing, leading to an “apparent specific absorption” at maximum in the order of 20 %. Compared to the systematic error of the integrating plate method, this overestimation of absorption is by a factor of 2–3 lower. In absorption measurements there is a possibility that the transparent coating on soot particles enhances absorption. In the case of the integrating sphere detector—even though the inorganic constituents might not dissolve in chloroform—their relative refractive index in chloroform is typically very low (e.g. 1.06 for (NH4)2SO4) [Hitzenberger et al., 2000]. Measurements of standard soot samples with the integrating sphere method yielded specific absorption coefficients of 6.56 ± 0.97 m2 g−1, in reasonable agreement with the results of Mie calculations assuming a refractive index of m = 1.96 − 0.66i and a density of 1.8 g cm−3 [Hitzenberger et al., 1996]. The αBC values for urban aerosol were found to be between 0.36 and 1.84 m2 g−1. It should be noted how low these values are when compared to σ values determined in urban environment (12.0 ± 0.7 m2 g−1) [Liousse et al., 1993]. First, αBC values are normalized to the mass of soot, which can be significantly higher than its EC content. Secondly, αBC is measured in suspended soot particles, contrary to σ which refers to BC particles embedded in a quartz fiber matrix. f) Photoacoustic spectrometry A photoacoustic spectrometer measures aerosol light absorption in situ, in the suspended state of the particles, thereby eliminating disadvantages inherent to filter-based measurements. In this instrument, light from an intensity modulated source is absorbed by the aerosol in a chamber, resulting in periodic heating and subsequent increase in pressure within the chamber. This in turn generates a sound wave which can be sensitively detected with a microphone. In the past, widespread application of this technique was limited by its low sensitivity, the inefficiency of available light sources and problems

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with calibration and window noise. The recent development of highly efficient solid-state lasers and improved acoustic resonator design greatly reduced window noise and raised instrument sensitivity to about 0.5 Mm−1 [Arnott et al., 1999]. The instrument is usually operated in the red at 685 nm or in the green at 532 nm using AlGaInP diode laser and diode-pumped frequency-doubled Nd:YAG laser as a light source, respectively. At both wavelengths gaseous light absorption can be neglected with minor contribution of NO2 absorption at the lower wavelength, which can be eliminated using appropriate metal inlets. A unique feature of photoacoustic instrument is that it can be calibrated using a gas, something that cannot be done with any of the filter methods. However, photoacoustic measurements of light absorption indicated systematic decrease above a relative humidity of 70 %, possibly due to the contribution of mass transfer to the photoacoustic signal [Arnott et al., 2003]. g) Raman spectroscopy In contrast to any other methods, which are non-specific, Ramanspectroscopy is specific exclusively to graphitic structures which are minor but strongly light-absorbing components of atmospheric aerosol. In its quantitative application for atmospheric graphitic carbon (GC) measurements, the mass loading of filter samples is obtained by integrating the band at 1600 cm−1 using a carbon black for calibration [Mertes et al., 2004]. Due to the specificity of the Raman method an accuracy of 10 % can be attained, down to a detection limit of about 0.08 µg cm−2.

2.2.2 Methods for EC Determination 2.2.2.1 Thermal Methods Although thermal methods have a long history in OC/EC determination in atmospheric aerosol, possibly the most widely used method was a two-step combustion procedure developed by Cachier et al. [1989]. The method was carefully optimized to minimize the major artifacts of thermal procedures, namely the charring of organic compounds and the untimely loss of soot carbon during the volatilization step. Prior to the determination of organic and elemental carbon carbonates are removed by exposing the filters to HCl fumes. During the first step, the precombustion step, the filter samples are heated rapidly to 340 °C in an oxygen atmosphere then held there for 2 hours.

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The rapid heating is necessary to favor desorption of organic compounds at the expense of charring, and the role of oxygen is also to minimize charring. Since prolonged exposure even to temperature as low as 320 °C was shown to gradually oxidize all dark-grey aerosol components, indicating that most of them are not pure graphitic carbon, the duration of the precombustion step is set to 120 min which was found to be at its optimum. The carbon fraction that remains on the filter after the precombustion step is referred to as EC. The filters are then combusted at 1100 °C. Both TC and EC determinations are performed by coulometric titration with a Ströhlein Coulomat®702C. The method was tested with carbonaceous matrices which were known to be free of soot. The results of these experiments are shown in Table 2.2. Table 2.2. Charring tests: T=340 °C; t=120 min (after Cachier et al. [1989]). Pure organic compounds Total carbon on the filter (µg cm−2) Glucose 9 15 20–40 pyrogallol 88 Thermally refractory natural mixtures (powders) Carbon % in the mixture humic acids 55 ± 5 (from marine sediment) fulvic acids 49 ± 3 (polysaccharides) dry leaf < 160 µm 45 ± 1 < 50 µm 45 ± 1 45 ± 1 < 20 µm Green leaf < 20 µm 45 ± 2 marine sediment < 20 µm 3.2 ± 0.2 fluvial sediment 4.3 ± 0.1 < 20 µm soil (humus) 9.7 ± 0.6 < 20 µm

Remaining carbon (%) 0 6.9 ± 0.6 10.0 ± 3.0 3.2 ± 0.2 Remaining carbon (%) 37 ± 3 29 ± 2 26 ± 1 20 ± 2 10 ± 1 15 ± 1 17 ± 3 16 ± 1 11 ± 1

Significant fraction of carbon was found to undergo charring in various organic matrices, up to 37 ± 3 % in humic acids, though the extent of charring tended to become smaller with decreasing particle size. The results implied that the extent of pyrolysis of the organic carbon was in the range of 5–20 %, which would cause significant overestimation of the EC concentration in

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aerosol. For example, it was later shown that up to 46 % of the carbon content of fungal spores can be erroneously assigned to EC by this method [Bauer et al., 2002]. However, the authors argued that the loss of refractory carbon which also occurred during the precombustion step exactly balanced charring (assuming a mean value of 10 %) when the EC/TC ratio was 0.24. In cases when soot concentration is either much higher or very low (as it is the case over most of the troposphere), correction needs to be taken. Unfortunately, this caution from the authors themselves has been overlooked in subsequent applications, and the method has gained a widespread application for measurement of EC at rural, remote or oceanic locations without any correction. In some studies, however, the thermal method of Cachier et al. was modified for the measurements of low-level aerosol samples [Lavanchy et al., 1999a]. In their setup, a combustion unit consists of two serial ovens at operating temperatures of 340 °C and 650 °C, followed with an oven filled with CuO catalyst and held at 950 °C to ensure complete oxidation to CO2. The CO2 evolved during each combustion step is trapped in a molecular sieve held at room temperature. Upon completion of each combustion step, the trap is flash-heated to 200 °C and the desorbed CO2 is detected with a nondispersive infrared analyzer. By definition, OC is the fraction that evolves at the pre-combustion temperature of 340 °C for 42 min, and EC is the fraction oxidized at 650 °C for 32 min. Such long combustion times were made possible by the trapping of evolved CO2 on the molecular sieve which yielded a sharp peak upon desorption. Furthermore, in order to minimize charring of organic compounds, prior to pre-combustion the sample is first introduced into the oven at 650 °C for one minute. The method was tested for EC artifacts with model organic compounds such as glucose, starch and humic acid. With the exception of the humic acid no artifact formation was observed. The charring of the highly polymerized humic acid (∼14 %) could not be avoided, but was reduced as compared to the original method. The use of bulk samples instead of spiked filters, however, increased the fraction charred to 33 %. The method was also compared to the German reference method VDI, which minimizes charring by solvent extraction of the sample prior to combustion [Petzold and Niessner, 1995; Kuhlbusch, 1995]. VDI methods In the VDI 2465/1 method a solvent extraction in a toluene-2-propanol mixture of 50:50 % (v/v) is performed for 24 hours to remove extractable

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organic compounds. After drying the filter for 4 hours under nitrogen then for 24 h in vacuo, the non-extractable organic compounds are removed by thermodesorption at 200 °C for 1 min then at 500 °C for 7 min. Eventually the remaining fraction—which is assigned as EC—is combusted at 650 °C and the evolving CO2 is determined by coulometry. The VDI 2465/2 method uses only thermal fractionation between the carbonaceous fractions. The sample is first heated under a flow of an inert gas at temperatures of 350 °C and 620 °C to remove organic compounds. Then oxygen is introduced into the gas flow (at least 20 % v/v) and the sample is heated to 700 °C to determine “EC”. The evolved gases pass a copper/cerium(IV)oxide catalyst to convert the gaseous carbon compounds into CO2 which is detected by non-dispersive infrared spectrophotometry (NDIR). Thermal manganese oxidation (TMO) method The thermal manganese oxidation method uses MnO2 in contact with the sample throughout the analysis as the oxidizing agent. Only the temperature is relied upon to distinguish between OC and EC: carbon evolving at 525 °C is classified as OC, and carbon evolving at 850 °C is classified as EC [Fung, 1990]. It is important to note that this method was optimized for the properties of polycrystalline graphite. Any less refractory high molecular weight material, even if it is virtually elemental carbon by composition, would be classified as organic by this method. a) On-line monitoring thermal techniques for EC determination High-temperature volatility technique A high-temperature volatility technique was developed to infer the presence of elemental carbon in aerosol [Jennings et al., 1994]. It is the volume of elemental carbon that is estimated from the fall off in number concentration between the critical onset temperature of 730–735 °C and 860 °C. There is also a gradual decrease in the number concentration between about 300 °C and 700 °C, which is attributed to the volatilization of organic carbon or soot carbon. Ambient Carbon Particulate Monitor (ACPM) The commercially available Ambient Carbon Particulate Monitor (ACPM) performs quasi-continuous analysis of carbonaceous aerosol on the

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principles of thermal analysis. For sampling, it uses a cartridge with an impactor plate with a 50 % cutoff of 0.14 µm at a flow rate of 16.71 l min−1 [Höller et al., 2002]. In order to further minimize adsorption of gas-phase organic compounds, the impactor plate is held at 50 °C which may induce substantial losses of semi-volatile organic species. The inlet is provided with a cyclone with a 2.5 µm cutoff. After the sampling cycle is completed, the cartridge is cut off from ambient air and is immediately subjected to analysis. Carbon analysis is performed in a closed cycle, encompassing the impactor plate, an oven, an afterburner held at 750 °C to convert all carbon species to carbon dioxide, and a non-dispersive infrared CO2 detector. Combustion can be performed in air at four temperature levels which can be pre-set within one analysis cycle. The default value for the OC-EC split is 340 °C. One sampling-measurement cycle lasts for a minimum of 2 hours. Detection limit for a 4-h cycle is 0.01 µgC m−3 [Matsumoto et al., 2003]. A major shortcoming of the instrument is that particles smaller than 0.1 µm are not sampled by the impactor, though this is the size range in which a considerable fraction of carbonaceous particles may be found. This negative bias can be most severe for BC particles whose mass concentrations may be underestimated by a factor of 1.9–3.1 [Höller et al., 2002]. The magnitude of the negative bias for OC particles is not known. Additional bias may result from particle bounce since no coating is applied on the impactor plates. Using two ACPMs with and without denuder revealed considerable positive artifact in spite of the use of the impactor, particularly for substances evolving below 200 °C [Matsumoto et al., 2003]. It was suggested that adsorption of gasphase organic species on the wall of the cartridge may be a major contributing factor. 2.2.3 Thermo-optical Methods of OC/EC Analysis 2.2.3.1 Thermo-optical Reflectance (TOR) Method The thermo-optical reflectance method of carbon analysis was developed by Huntzicker et al. [1982]. The filter is first combusted at 350 °C in 2 % O2–98 % He mixture, the volatilized and partially oxidized carbon is oxidized to CO2 in the MnO2 bed held at 950 °C, reduced in the methanator to CH4 and measured by a flame ionization detector. The combustion zone is then purged with helium, then the temperature is raised to 600 °C and the carbon is volatilized into the helium carrier gas. For EC measurement the combustion zone temperature is dropped to 400 °C, the carrier gas is switched back to 2 % O2–98 % He, then the oven is heated first to 500 °C for 120

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seconds and 600 °C for 200 seconds. During the measurement the reflectance of the filter is monitored continuously by a He-Ne laser of 633 nm wavelength. Significant degree of charring was observed in the second stage of organic carbon determination (600 °C in pure helium), especially for biological samples such as wood fiber, leaf, pollen, for which 45 %, 64 % and 18 % of total carbon were converted to BC. In this method “volatile organic carbon” is defined as carbon that evolves at 350 °C. “Residual organic carbon” is the fraction of carbon that evolves between 350 °C and 600 °C in pure helium, plus that which evolves in 2% O2/98 % He at 400 °C up to the point when the intensity of the reflected light regains its initial value. The fraction which subsequently evolves at ≥ 400 °C in 2% O2/98 % He is defined as EC. After extraction of ambient aerosol filters in a 2:1 mixture of chloroform/methanol, the extent of charring was reduced by 80 %. Carbonization (charring) seemed to extend completely through the filter, as the back side of the filter also showed coloration. It was suggested that oxidative removal of organic compounds at 340 °C minimizes charring and does not affect EC [Johnson et al., 1981]. Pre-extraction of urban aerosol in benzene-ethanol (4:1 by volume) also decreased the apparent EC by 30 % [Cadle et al., 1982]. Slightly lower reduction in apparent EC concentration was observed after oxidative removal in air at 340 °C. However, in urban aerosol it was found that ∼6 % of the absorbance at 550 nm was due to extractable organic compounds. The DRI thermal/optical reflectance (TOR) analyzer is essentially a modified version of the first instrument [Chow et al., 1993]. The scheme of the instrument is shown in Figure 2.6. The TOR analyzer works on the principle of 1) liberating carbonaceous compounds successively in pure He and oxidizing atmosphere at different temperatures; 2) converting all carbon species to CO2 over MnO2 held at 912 °C; 3) reducing CO2 to CH4 in a methanator (firebrick impregnated with Ni catalyst under H2 at ∼550 °C); quantification of CH4 by flame ionization detector. The pyrolysis correction is made by continuously monitoring the filter reflectance via a He-Ne laser at a wavelength of 632.8 nm and a photodetector. The reflectance, which is largely dominated by elemental carbon and light-absorbing organic carbon, decreases as a result of charring during pyrolysis and increases as light-absorbing carbon is burned. The optical pyrolysis correction assumes that the light extinction per unit mass of pyrolytically produced carbon is the same as the light extinction per unit mass of carbon removed in the EC1 segment until the reflectance regains its initial value. Since this condition is unlikely to be met, there is an inherent bias in either direction in the demarcation between light-absorbing and elemental

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carbon. The possible magnitude of this effect cannot be estimated but it is certainly much less than having no optical correction at all. A typical TOR thermogram is shown in Figure 2.7. This thermogram can be divided into seven carbon fractions as a function of temperature and atmosphere. The temperature in pure He ramps as rapidly as possible from room temperature to 120 °C (OC1), 120 °C to 250 °C (OC2), 250 °C to 450 °C (OC3), and 450 °C to 550 °C (OC4). It is critical to ensure that no traces of oxygen be present in the He. Traces of oxygen can be evidenced by the continuing downward drift after the end of the OC4 segment. Ramping to the next temperature starts when the detector response returns to baseline or attains a constant value, but is constrained by the condition that the time spent in each segment should be between 80 and 580 s (850 s in the case of EC1). After this condition has been reached in the OC4 segment, the atmosphere is changed to 2 % O2 /98 % He, and the peaks are integrated at constant temperatures of 550 °C (EC1), 700 °C (EC2), and 800 °C (EC3), subject to the conditions detailed above [Chow et al., 1993].

Figure 2.6. DRI thermal/optical reflectance (TOR) analyzer block diagram (after Chow et al. [1993]).

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Carbonaceous Aerosol

Figure 2.7. Example of DRI thermal/optical reflectance carbon analyzer thermogram. Reflectance and FID output are in relative units. Reflectance is normalized to initial reflectance and FID output is normalized to the area of the reference peak. This is an ambient sample collected on 17 May 1989 in Yellowstone National Park (after Chow et al. [1993]).

Using the TOR method it is also possible that the reflectance actually increases due to the volatilization of light-absorbing organic carbon. This was a rare occurrence, suggesting that light-absorbing organic species do not commonly volatilize at or below 550 °C. Thus any light-absorbing organic carbon is classified as elemental carbon by the TOR method. “Elemental carbon”, as defined by the method, thus includes elemental carbon, lightabsorbing organic carbon, any refractory organic carbon evolving after the pyrolysis correction, and the bias of the pyrolysis correction. Carbonate carbon can be removed prior to the carbon analysis by acidification. Without this initial step, carbonate carbon is detected as either OC or EC, since carbonates could evolve in almost any segment of the thermogram. Most carbonate is likely to evolve around 600 °C, which would put them in the EC2 segment of the thermogram. This would also mean that

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even a small relative amount of carbonate carbon could introduce a substantial bias into the determination of elemental carbon unless it is removed by acidification [Chow et al., 1993]. Acidification, however, may also liberate volatile organic acids from their salts, which would be correspondingly missing from the fraction of organic carbon, such as HCl + CH3COONa → NaCl + CH3COOH↑

(2.5)

Therefore “carbonate carbon” might be more properly termed as “acidreleased” carbon. Oxalic acid, a major aerosol constituent tends to form insoluble calcium oxalate with calcium present in sea salt aerosol. Since upon heating this species decomposes to CaCO3 which is thermally refractory up to 650 °C, and CO, only half of its mass will be determined as OC, and the other half will appear as EC in the thermal/optical method [Neusüss et al., 2002]. 2.2.3.2 Thermal/optical Transmission Method (TOT) The thermal/optical transmission method (TOT) differs from TOR in that it uses transmission instead of reflectance for the optical pyrolysis correction. It was first described by Turpin et al. [1990] and later developed into a semi-continuous carbon analyzer which has been widely used for determination of OC/EC in aerosol. TOR and TOT agreed within 1 % for diesel exhaust particulates, but EC determined by TOR generally exceeded that of TOT for other aerosol, the worst being woodsmoke in which the disparity was four- to sevenfold [Chow et al., 1993]. The fourth-generation thermal/optical reflectance/transmittance (TOR/TOT) carbon analyzer was developed to resolve the difference between TOR and TOT, allowing concurrent reflectance and transmittance measurements [Chow et al., 2001]. It was hypothesized that when the reflectance of the filter surface regains its initial value, there might be still unburned pyrolytic material beneath the surface, causing undercorrection and erroneously high elemental carbon values. Consequently, TOT values are likely to be more correct. 2.2.3.3 Charring in Thermal/optical Methods As it was shown, one of the major contributing factors to the disagreement between thermal EC/BC measurements is charring of some

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organic constituents. Its correction in thermo-optical methods is done by measuring the fraction of EC oxidized that is necessary to return the reflectance or transmittance of the filter back to the initial value before charring occurs. However, large uncertainty is associated with the measurements for aerosol containing a large fraction of pyrolyzable OC and little EC. Unfortunately, charring depends on many factors, including the amount of OC, temperature, heating periods, the presence of inorganic constituents. It was demonstrated that under conditions of the NIOSH standard method to measure EC in diesel particulates [Birch, 1998] significant charring occurred in the water-soluble organic fraction of urban and marine aerosol [Yu et al., 2002]. The char formation was evidenced by the decrease in filter transmittance and by the fact that some OC evolved under the thermal analysis conditions during which aerosol EC would have evolved. There has been disagreement on the chemical nature of the carbon fraction evolving at 850 °C under He in the NIOSH method which assigns it as part of OC. On the other hand, Chow et al. [2001] argued that it is part of EC which is oxidized by the traces of oxygen supplied by the thermal decomposition of the mineral oxides on the filter. Yu et al. [2002] lined up additional evidences suggesting that increasing transmittance is due to the evolution of light-absorbing polymeric OC. The percentage of charring increased with increasing carbon load on the filters up to about 7 µg cm−2 then stabilized at around 39 %. On the contrary, the n-hexane extract of the same aerosol samples showed very little charring (< 2 %). It was calculated that 13– 66 % of the overall char formation of the aerosol came from the WSOC, and the remaining fraction was tentatively assigned to cellulose and humic substances present in the insoluble fraction. The presence of O2 in the carrier gas influences charring as part of the OC can react with O2. Some charring was found to occur and ∼20 % of OC evolved at 850 °C in an oxidative atmosphere [Yu et al., 2002]. It was concluded that a part of aerosol WSOC and EC require the same conditions to evolve in thermal analysis, thus they are indistinguishable from each other. Furthermore, thermal analysis of starch and cellulose in the presence of NH4HSO4 produced 2–3 times the amount of char that was formed in the absence of the inorganic compound [Yu et al., 2002]. In contrast, charring of levoglucosan was reduced by 15 % in the presence of NH4HSO4. Two thermal programs have been used widely in the analysis of particulate carbon, namely the NIOSH and IMPROVE methods which mainly differ in the residence time at each temperature step. The NIOSH method uses fixed residence time, whereas the IMPROVE method prescribes variable residence time to allow the FID peaks to return to baseline and to give rise to

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well-resolved peaks. Charring is more significant with the NIOSH method (by a factor of ∼1.6), indicating that the OC that would otherwise evolve if given more time apparently contributes to the subsequent charring.

Figure 2.8. Thermograms of the analysis of the PM3.5 sample collected in Bakersfield, CA, on December 8, 1995. Panel a) shows the thermogram of the EC/OC analysis using the base temperature program for the EC/OC analysis. Panel b) shows the thermogram of the EC/OC analysis using the alternative 3 temperature program (after Schauer et al. [2003]).

In a recent study a thermo-optical method was tested to determine the sensitivity of the EC/OC split on the temperature program [Schauer et al., 2003]. In this method, OC is defined as carbon that evolves under a heating cycles in the presence of pure He, and EC is operationally defined as the component which evolves in the subsequent heating steps in the presence of oxygen in the purge gas. The samples tested were atmospheric aerosol, wood

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smoke, coal fly ash, secondary organic aerosol from smog-chamber experiments as source samples, and organic matrices such as wood smoke extract, candle wax and motor oil. As expected, all atmospheric aerosol samples pyrolyzed during analysis, as did wood smoke and secondary organic aerosol among the source samples, and wood smoke extract among the organic matrices. Even with the optimized 10-step temperature program, which was used throughout the intercomparison (shown in Figure 2.8 for the analysis of an atmospheric aerosol sample), it was not possible to fully correct for charring. The outcome was that elemental carbon was measured by the thermooptical method in samples in which it was not expected to be present. In other terms the measurement of EC was biased by some organic matrices. For example, in the wood smoke extract, 5 % of the carbon was assigned as elemental by the thermo-optical method. In case progressively lower temperatures were used in all steps in which the sample was heated in pure helium, the bias increased significantly. When the lowest combustion temperatures were used in these steps, charring continued in the subsequent steps when O2 was already present in the purge gas. 2.2.4 Intercomparison of various methods In a recent intercomparison of TC/EC/BC measurement methods 17 laboratories participated in the analysis of urban aerosol samples [Schmid et al., 2001]. A group of them employed the VDI 2465 methods which are standard methods in Germany. The TC measurements showed good agreement, with the interlaboratory relative standard deviations were generally below 10 %. On the other hand, the results of EC/BC determinations showed a wide scatter, interlaboratory RSD values were in the range of 24–46 %. Since no standardized reference sample was available for EC/BC, the results of the round robin test were relied on. Since the results were likely biased in one direction, namely thermal methods without correction for charring (carbonization) tended to overestimate EC, the “true value” was determined from a subset of results which were deemed reliable. A comparison of the results revealed that the VDI methods, in particular VDI 2465/2, determined higher concentrations of EC than the thermo-optical methods. The thermo-optical methods correct for the systematic bias introduced by charring provided that the filters are not overloaded. Without optical correction, EC values higher by a factor of 2–3 are obtained by these methods. Another correction method might be pre-extraction of the filters in a mixture of dimethyl-formamide and toluene (75: 25 % v/v) prior to the

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application of the VDI 2465/2 method, which yielded similar results as the thermo-optical method. Another comparison between different carbon analysis methods revealed interesting differences [Chow et al., 1993]. For example, in smogchamber aerosol which contained no EC, the TOR/TOT method read ∼2% EC, whereas the TMO method yielded 8 %. In other cases, TOR readings of EC were considerably higher than that of TMO, especially when EC constituted only a small fraction of total carbon. The worst disparity was observed in woodsmoke aerosol, where the differences were 8–53-fold. It is likely that most of the EC1 segment of the TOR method was classified as organic carbon by the TMO method, since it used similar temperature in the presence of an oxidant. Since EC1 usually contains most of the light-absorbing carbon, misclassification could be of great consequence. Furthermore, it is known that graphitic carbon in aerosol may evolve at temperatures as low as 440–470 °C in pure O2 [Rosen et al., 1982; Novakov 1982]. The “lighter fluid effect” (used in regenerative particulate traps in automobiles), wherein elemental carbon is ignited through the combustion of organic carbon, may also cause premature combustion of EC. It is should be remembered that the TMO method has been optimized for the properties of polycrystalline graphite. Thus the excellent agreement between the methods for diesel exhaust particulates is understandable. The light-absorbing carbon in diesel exhaust particulates exhibits a weakly ordered structure approaching that of polycrystalline graphite, and contains very little polymeric organic material. Consequently, the TOR thermogram of the diesel exhaust particulate showed very little EC1 and was almost devoid of the optical evidence of any pyrolysis [Chow et al., 1993]. Simultaneous measurements with a photoacoustic spectrometer and an off-line TOR method revealed high correlation between light absorption and EC, yielding an average EC absorption coefficient of 3.85 m2 g−1 [Moosmüller et al., 1998]. Correlation was poorer between light absorption and integrating plate BC measurements, with high slope and zero offset values. This correlation improved significantly when the correction method suggested by Horvath [1997] was applied to the uncorrected filter data. The correlation between the aethalometer and photoacoustic measurements was satifactory with an absorption efficiency for BC of 10 m2 g−1 at 532 nm, in line with the assumptions by Bodhaine [1995]. At 685 nm, however, the average absorption efficiency of aethalometer BC was found to be about a factor of 2 lower (5 m2 g−1) [Moosmüller et al., 1998]. Fitting a power law to experimental data revealed a wavelength-dependence of λ−2.7. An on-line high-temperature volatility technique was also compared to those measured with an off-line thermal method and an aethalometer [Jennings et al., 1994]. Since the latter two methods are targeted to the

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determination of BC rather than EC, from the comparison it was also possible to infer a measure of the relative abundance of EC in the BC fraction. A minimum percentage of 30 % was obtained for the polluted air mass studied, though BC determinations themselves were subject to a large uncertainty, and differed from each other by a factor of 2.

Chapter 3 MAJOR CARBONACEOUS PARTICLE TYPES AND THEIR SOURCES 3.1 Soot—Black Carbon (BC), Elemental Carbon (EC), Graphitic Carbon 3.1.1 Definitions Carbon is typically the largest elemental fraction of atmospheric aerosol particles, present in many different chemical and physical forms. The total amount of carbon (TC) that can be found in particulate matter can be readily determined by elemental analysis. The accuracy of the measurements is that typical of most analytical determinations. No uncertainty results from the definition itself (as in the case of all other sub-categories), though serious over- or underestimation of TC concentrations may come from sampling artifacts (adsorbed volatile organic species or volatilization losses, see subchapter 2.1). On the other hand, a specific subset of aerosol carbon—generally representing a minor fraction of TC in aerosol—does not even have a common name: it is called black carbon (BC), elemental carbon (EC), graphitic carbon, or soot. Accordingly, there are numerous definitions for this sub-category of carbonaceous matter. In fact, however, such uniform fraction does not exist, or more precisely, it is method-dependent since it cannot be unambiguously separated from organic carbon by any of the methods. Although the carbon fractions referred to by the different definitions overlap to a large extent, they are far from being the same. The differences are also manifested in intercomparisons of various methods of the determination of BC/EC, which often reveal substantial and systematic differences in concentrations. For detailed discussions see sub-chapter 2.2. The only physical and observable carbon particle type which is intended to be represented by all other sub-categories is soot. Soot is often associated with combustion-generated primary carbonaceous aerosol.1 Soot is recognizable by its special morphology by scanning or transmission electron 1

The term “smoke”, particularly in connection with coal burning, was sometimes used as a synonym for soot.

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Carbonaceous Aerosol

microscopy. In chemical terms, soot consists of an array of individual organic compounds, soluble in organic solvents, and a refractory and insoluble carbonaceous matter which is termed “elemental carbon (EC)”, “graphitic carbon” or “black carbon (BC)”. Unfortunately, in the high vacuum of the electron microscope, the organic component of soot may be lost, so what is actually seen is its vacuum-resistant core. Sometimes soot is identified as primary carbon derived from combustion, which may not be correct since this definition ignores other elements associated with organic matter within or on the surface of soot [Novakov, 1984]. Other studies divide soot into two sub-categories: the black carbon soot of either amorphous or graphitic microstructure, and the second, less well-known is the organic soot released in certain stages of biomass burning [Blake and Kato, 1995]. Albeit it is a matter of nomenclature, it has recently been postulated that the latter is actually not primary soot, but a specific subset of secondary organic aerosol, “tar balls” [Pósfai et al., 2004]. The choice and use of these synonyms tend to be separated by the method of determination. In other terms, any definition of organic and black/elemental carbon is operational and reflects the method and purpose of measurement. Elemental carbon (EC) is the preferred term in conjunction with thermal and wet chemical determinations which are deemed suitable for the selective measurement of the refractory component. The term implicitly infers that this component consists entirely of carbon, though none of the methods used are capable of determining carbon as an element. The methods which are capable of determining elemental composition are in fact suitable only for the determination of total carbon (TC). In addition, there is a conceptual ambiguity in the definition of organic carbon (OC) and EC. Since the carbon/hydrogen ratio approaches infinity within the homologous series of polynuclear aromatic hydrocarbons, the transition between organic and elemental carbon is gradual. A synonym for EC is the term “graphitic carbon” which infers the presence of graphitic-like microcrystalline structure. This can only be evidenced by Raman spectroscopy or high-resolution transmission electron microscopy (HRTEM). Whereas these measurements are highly specific and conclusive, they are unsuitable for quantitative purposes since even graphitic black carbon particles have only poorly developed graphitic structure and a wide variety of oxygen-bearing functional groups on their surface. The term “black carbon (BC)” implies that this component is responsible for the absorption of visible light, and is used mostly when optical methods are applied for its determination. It is often unclear whether this term refers to the carbonaceous matter or only to the carbon itself (although the conversion factor between the two is not much different from unity).

EC/BC/Soot

47

Formerly, BC was defined as “black particulate carbon having a graphitic microstructure” [Novakov, 1984]. Nowadays most definitions of BC emphasize the light-absorbing property of carbonaceous aerosol, and consider BC as the only optically absorbing component of soot and smoke [Penner et al., 1993]. However, in the light of recent research, “light-absorbing carbon” would be a more appropriate term for light-extinction budgets or optical measurements than BC. This is because there can also be many organic compounds in the aerosol which do absorb light, and usually they are not black (“brown carbon”) [Bond, 2001; Andreae et al., 2002]. However, the extent to which organic compounds contribute to aerosol absorption has not been determined yet. For lack of generally accepted terminology, the terms BC and EC are still often used interchangeably in the literature, without regard to the above classification. Even if one tries to use these terms consistently, there are many cases when the above tacit classification is difficult to observe. For example, thermo-optical methods, which are now widely accepted as the most reliable methods for OC/EC determination, combine two basically different principles which were traditionally assigned to EC and BC measurements separately. Since this method was developed from a thermal method used for EC determination it seems evident that it still belongs to this family. BC may also be used as a synonym for soot under special conditions, in spite of the fact that BC represents only a sub-fraction of soot by mass. However, this specific sub-fraction is thought to responsible for all lightabsorption by soot, and is defined and measured accordingly. Therefore in light-extinction budgets or radiative transfer models BC does stand for soot. In fact it needs to, since BC is the only form of carbon whose optical properties are relatively well known. Care has to be taken, however, not to take the mass concentration of BC equivalent to that of soot, for example in mass closure calculations. EC may seem to be better suited for this purpose after appropriate conversion using the mass fraction of carbon in soot. This approach, however, has not become common practice: in mass closures only the mass concentration of BC or EC is considered, neither of which are characteristic particle types. Following the recommendations by Penner and Novakov [1996], we will use the term BC as a synonym for EC throughout this book in any general context when the method of determination is not explicitly stated or not important. There are several reasons for supporting this approach. The first is that among the atmospheric effects of soot light-absorption is of utmost importance. Secondly, as pointed out above, BC is the equivalent of soot in radiative transfer calculations. Thirdly, soot (and even BC) does contain other elements besides carbon, so EC is not a good surrogate for soot in any

48

Carbonaceous Aerosol

atmospheric processes. Even heterogeneous reactions on soot surfaces are profoundly affected by functional groups on the carbonaceous surface. However, when referring to individual particle types, soot will be used consistently. 3.1.2 History of Soot Pollution Possibly smoke or soot was the very first pollutant ever recognized by man, but interestingly one of the last that was studied in depth by the atmospheric research community [Penner and Novakov, 1996]. Its documented history dates back into the thirteenth century England [Brimblecombe, 1978]. Faraday [1861] already recognized that soot was composed of carbon and produced by incomplete combustion of carboncontaining fuels. The disastrous effect of coal burning in London on human health and mortality in the early 1950s invoked an urgent need to measure the atmospheric concentrations of soot. These provided the first regular measurements ever for an important type of carbonaceous aerosol, and also outlined some aspects of aerosol-fog interactions. Soot derived from bituminous coal or oil combustion was known to be of micrometer sized particles which can pollute large areas far from its sources [Wilkins, 1954]. Its residence time was estimated to be 2–3 days, resulting from its low deposition velocity [Brooks, 1954]. Unfortunately, in the U.S. where the air pollution problem known as the Los Angeles smog had already emerged that time, scientific misinterpretation of its nature led to an almost complete neglect of black carbon for a long time. Interest in carbonaceous aerosol in cities at that time was fuelled by the need for determination of organic compounds in aerosol (in particular with respect to carcinogenic compounds), and for a methodology that infers total suspended particulate (TSP) concentrations, as regulated in air quality standards, from simple optical measurements. The first systematic measurements of long-range transport of soot were initiated in Sweden, under the suspicion that much of the air pollution originated from sources outside the country [Rodhe et al., 1972; Brosset et al., 1975]. However, BC was first put into a global perspective in studies on the Arctic haze [Rosen et al., 1981]. Then the issue of nuclear winter scenarios [Crutzen and Birks, 1982; Turco et al., 1983] and the long-range transport of biomass burning smoke [Andreae, 1983] gave impetus for the study of carbonaceous aerosol. By the late 1980s, biomass burning smoke was recognized to have global climatic effects [Penner et al., 1991].

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49

3.1.3 Properties and Chemical Composition of EC/BC/soot 3.1.3.1 Morphology It is widely accepted that primary soot particles do not have a separate existence—rather they cluster together to form aggregates that are their most stable form. Soot particles freshly emitted from Diesel engines or aircraft combustors often exhibit ramified and apparently irregular agglomerate structures. Using transmission electron microscropy, the average diameter of primary spherules was determined to be 22.6 nm ± 6.0 nm [Wentzel et al., 2003]. Other studies determined somewhat larger size of primary particles between 30 and 50 nm, obviously affected by a number of parameters, including fuel characteristics, engine types and operating conditions [Clague et al., 1999]. It was shown that spherical and compacted particles are associated with the smoldering stage of biomass combustion [Martins et al., 1998]. These particles seem to be more stable with respect to atmospheric aging processes. In contrast, flaming combustion is conducive to the formation of spherules of about 50 nm diameter, which aggregate to produce nonspherical particles. In fresh smoke these clusters tend to form open structures, which are then transformed by aging processes—either by cloud processes or by interaction with water vapor—into more closely packed particle types. These particle types are typical of biomass burning, as illustrated in Figure 3.1.

Figure 3.1. Examples of some possible mixtures between black carbon (BC) and nonabsorbing materials in smoke particles (after Martins et al. [1998]).

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Carbonaceous Aerosol

In aged wood smoke single spherical particles were visible by scanning electron microscopy. They can be readily identified, as they retain their individuality (do not spread) on the sampling substrate [Rogers et al., 1991a]. Lammel and Novakov [1995] combusted diesel oil in an alcohol lamp and collected freshly generated diffusion flame aerosol on Nuclepore filters, which was black and fluffy by visual inspection. Scanning electron microscopy revealed agglomerates forming branched chains as well as large, more compact secondary structures. The primary particles were spherical, with diameters of 60–120 nm. The uncontrolled combustion of fossil fuel, such as oil fires in Kuwait were shown to produce highly nonspherical particles in the size range of 1–4 µm diameter, and relatively few particles below 0.4 µm diameter, the size range typical of biomass burning [Weiss et al., 1992]. Soot aggregates collected on the polluted North Atlantic are either compact or chain-like associations of spherules having diameters between 10 and 50 nm [Pósfai et al., 1999]. The spherules are attached to one another and coated by a carbonaceous film. The internal structure of the spherules of combustion soot revealed by high resolution transmission electron microscopy (HRTEM) showed curved, disordered graphitic layers. The spacing between the chaotically arranged layers was typically larger than in pure graphite (0.335 nm) [Pósfai et al., 1999]. On the other hand, soot aggregates from the Southern Pacific consisted of spherules of only 10–20 nm, and did not seem to be glued together by carbonaceous film. The individual spherules had more ordered structure with wavy but roughly concentric graphitic layers which formed an onion-like structure. 3.1.3.2 Chemistry The initial chemical composition of soot depends strongly on its sources: some sources can produce almost pure elemental carbon, while others produce soot of which 50 % by mass is organic matter. Graphite, the ultimate form of pure elemental carbon in the atmosphere, is an absolutely inert material under atmospheric conditions. It can be oxidized in air only at temperatures exceeding 600 °C, will react in solution with strong oxidizing agents such as fuming nitric acid or KMnO4, and reacts explosively with F2. In spite of being inert, the surface impurities of soot can initiate adsorption and catalyze chemical reactions [Charlson and Ogren, 1982]. Besides the σbond between each carbon atom, the extra electron on each C atom is available for π-bond. Each carbon-carbon bond achieves only one third of

EC/BC/Soot

51

double bond character, unlike a strictly aromatic or olefinic linkage. The resonance of the π-electrons among various configurations accounts for its electrical conductivity and broad-band light absorption. This latter property is of utmost importance in radiative transfer in the atmosphere, and serves as a basis for several methods of its determination, thus it is discussed in depth in sub-chapters 5.1 and 2.2. Graphitic structures in which carbon atoms occupy lattice sites in a twodimensional honeycomb network have intense Raman modes but very weak IR vibrational absorption. These Raman modes enable unambiguous identification of graphitic structures in atmospheric aerosol, possibly as the only method available at a molecular level [Rosen and Novakov, 1977]. In the Raman spectra of various forms of graphite, the first-order transitions lie between 1200 and 1700 cm−1 while the second order transitions are near 2700 cm−1, extending to about 3500 cm−1. The shapes and intensities of Raman transitions are functions of the graphite crystal size, morphology and sample composition [Sze et al., 2001]. The morphology, in turn, is closely related to the formation process of the graphitic component. The matrix in which the graphite crystals are embedded also has some effect. The structure of hexane soot was studied with successive extraction in a suite of solvents combined with instrumental analyses [Akhter et al., 1985]. The soot extracts was found to amount up to 32.9 % of the mass of soot, and consisted of polyaromatic hyrocarbons (PAHs), oxygenated polyaromatic hydrocarbons, and a small fraction of aliphatic compounds. The aromatic to aliphatic ratio was found to be at least 9:1. Although on an elemental basis the overall composition was still dominated by carbon, concerning the oxidation state of carbon it was far removed from elemental carbon. Based on these findings, the chemical structure shown in Figure 3.2 was envisaged. Interestingly, the infrared spectrum of soot did not change qualitatively after extraction. A recent high-resolution transmission electron microscopic study on kerosene soot particles revealed that they appear to be almost amorphous, with some signatures of short-range order [Ferry et al., 2002]. The distance between the graphite planes was somewhat larger than in pure graphite crystals, and onion-like particles were made of crystallites having a very small number of graphite planes, usually 2–4, having a lateral extension of about 2– 3 nm. The average elemental composition of kerosene soot was 92.2 % (m/m) C and 7.8 % (m/m) O. Diesel soot usually show the onion-shell microstructure of about 7–10 graphite lattice planes in single domaines. Therefore Diesel soot particles are often referred to as having a turbostratic structure [Clague et al., 1999].

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Carbonaceous Aerosol

Figure 3.2. Representation of hexane soot segment as formed in flame (after Akhter et al. [1985]).

Therefore any soot particle may be regarded as a complex threedimensional organic polymer with the capability of transferring electrons, rather than merely an amorphous form of elemental carbon [Chang et al., 1982]. The relatively low mass fraction of oxygen in soot may be deceiving since most of them are actually found on the surface of soot in various functional groups. The relative amounts of these surface complexes depend on the thermal history of carbon particles. Nitrogen species are also present on the soot surface. The interaction of NH3 with soot was shown to produce ammonium-like species [Chang et al., 1982]. At low temperatures the surface carboxyl or phenolic functional groups may act as Brönsted acids in their interaction with NH3, forming carboxylic or phenolic ammonium salts by proton exchange. Ammonia may also adsorb by hydrogen bonding to surface −OH or carboxylic groups. From the chemistry and structural features of soot particles it follows that the term “elemental carbon” is actually a misnomer in atmospheric sciences. The existing particle type is soot whose chemical composition is

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53

indeed dominated by carbon, but not necessarily in its zero oxidation state: the share of organic species can be substantial. In addition, due to the abundance of surface functional groups, soot particles do not really behave as pure elemental (graphitic or amorphous) carbon in most atmospheric processes. 3.1.3.3 Mixing State of Soot in the Atmosphere Transmission electron microscopic studies revealed that soot occurred internally mixed with sulfate particles in the remote marine troposphere [Pósfai et al., 1999]. Most soot aggregates collected on the polluted North Atlantic contained sulfur, and some also traces of K, Na, Si or Ca. As for the sulfate particles, about 50 % of the smallest (0.1 µm) and as much as 90 % of the largest (1 µm) sulfate particles had soot inclusion or attachment in the remote marine troposphere [Pósfai et al., 1999]. When soot occurred at the edge of a sulfate particle, it is likely that it was forced to the periphery when sulfate crystallized from hydrated aerosol droplet. Independent soot particles were not found in the size range covered by sampling on grids. In general, the fraction of soot-bearing sulfate apparently increased with altitude. Internally mixed particles may form when soot particles get activated in contrails behind cruising aircrafts, primarily through the condensation of H2SO4 on the surface of soot [Rogaski et al., 1997; Yu and Turco, 1998]. Heterogeneous processes on the surfaces of soot particles during transport could further increase the number of internally mixed sulfate-soot particles. It is likely that activated soot particles provide a suitable aqueous medium for SO2 oxidation to proceed. The importance of this process, however, is subject to controversy in the literature [Novakov et al., 1974; Chang et al., 1981; Mamane and Gottlieb, 1989]. However, some soot particles in the pristine Southern Hemisphere, especially those which did not contain measurable K, occurred without being aggregated with sulfate [Pósfai et al., 1999]. These particles were likely emitted by aircrafts, and escaped the heterogeneous processes which would have made them internally mixed. The varying efficiency of these processes was also indicated by the highly variable soot/sulfate ratio in the range of 0.01–0.25. An important consequence of the chemical purity of the particles may be that they are poor cloud condensation nuclei, as opposed to soot from biomass burning or coal combustion which are active CCN [Grğic et al., 1993].

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3.1.3.4 Scavenging efficiency for BC Due to its chemical composition and hygroscopic properties, BC is expected to be scavenged much less efficiently than highly water-soluble species such as sulfate, provided that they are externally mixed. Since this condition appears not to be valid over most of the troposphere, the parameter used for the quantitative characterization of scavenging, the mass scavenging efficiency (ε), must be viewed with caution. The mass scavenging efficiency is defined as the ratio of BC mass concentration in cloud water per unit volume of air to total BC concentration in the same volume. In spite of the fact that it is based on measurable quantities, ε remains a hypothetical parameter since it refers to an ideal state which barely exists in the troposphere. The concept of scavenging efficiency implicitly assumes that scavenging from the gas phase does not contribute significantly to the concentration of the species in cloud water. Data for εBC in fog or clouds are rather sparse in the literature. A comparison of the average scavenging efficiencies for BC and SO42− for various locations is presented in Table 3.1. Table 3.1. Average scavenging efficiencies ε for BC and SO42− at various locations (after Hitzenberger et al. [2001]).

Rax, 2000 Sonnblick, 1997 [Hitzenberger et al., 2000] Po Valley, Italy [Hallberg et al., 1992] Kleiner Feldberg, FRG [Hallberg et al., 1994] Great Dun Fell, UK [Gieray et al., 1997] Spitsbergen (Norway) [Heintzenberg and Leck, 1994]

ε for BC

ε for SO42−

0.54 ± 0.25 0.74 ± 0.19

0.78 ± 0.23 0.91 ± 0.08

0.06

0.18

0.15

0.51

0.57

0.77

0.80

0.80

In general, available data tend to support the expectations based on differences in hygroscopic behavior and size distribution only in the proximity of sources. For example, in the heavily polluted Po Valley, Italy, εBC in fog was found to be 0.06, as compared to εSO42− of 0.18 [Hallberg et al., 1992]. At a rural site in Germany εBC = 0.15 and εSO42- = 0.51 were measured [Hallberg et al., 1994]. In a fog study at Berkeley, California, εBC = 0.26 was found, resulting in a BC concentration of 0.16 µg ml−1 in fog water [Gundel et al., 1994].

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In contrast to ground fogs, the difference between scavenging efficiencies of sulfate and BC tend to diminishes in clouds which usually form on more aged aerosol. For example, in clouds of continental origin at Great Dun Fell, Great Britain, εBC = 0.57 and εSO42−= 0.77 were measured [Gieray et al., 1997]. Hitzenberger et al. [2000] found the average scavenging efficiencies to be 0.74 ± 0.19 for BC, 0.57 ± 0.21 for TC and 0.91 ± 0.08 for SO42−, respectively, at the Mt. Sonnblick. The scavenging efficiencies of all three species showed some dependence on the cloud liquid water content, in particular for LWC < 0.4 g m−3. As an extreme, at the remote Spitzbergen identical value for εBC and εSO42− was reported, suggesting an internal mixture of the two components [Heintzenberg and Leck, 1994]. The scavenging ratios (W) for BC were also determined experimentally in a few studies. Expressed as the ratio of mass mixing ratio of BC in precipitation divided by the mass mixing ratio of BC in air, these values spanned a wide range at different locations and precipitation types. For example, in the Artic, the values were found to be between 60–160 for snow [Clarke, 1989], whereas in urban locations for rain and rural location in snow they found to vary between 18–650 and 49–1260, respectively [Dasch and Cadle, 1989; Cadle and Dasch, 1988]. However, the mean values were reported within a factor of three, irrespective of the location and type of precipitation [Penner et al., 1993]. There are two effects both of which tend to make εBC converge to εSO42− when moving from polluted fog to clouds formed in remote areas. Ageing processes produce internally mixed aerosol, i.e. soot particles become associated with sulfate which then drive them into cloud droplets. This process can be indirectly evidenced by the simultaneous determination of BC and sulfate mass size distribution. It is also possible that the surface of soot particles becomes more hydrophilic upon aging, and even if they do not associate with sulfate particles, they can activate at lower supersaturations. The change in the mixing state of aerosol upon aging could be more important than modification of the soot surface. Yet in atmospheric models ageing is usually represented by gradual transformation of hydrophobic BC into hydrophilic one, with prescribed time-constants. 3.1.4 Emission Sources of EC/BC/soot Soot can be generated by the pyrolysis of aromatic or aliphatic compounds, and even methane can produce soot, demonstrating that even small hydrocarbon molecules can produce graphitic carbon. Another possible gas-phase production mechanism involves the equilibrium

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Carbonaceous Aerosol

CO + CO ↔ C + CO2

(3.1)

which is shifted to the right at T >600 °C. Furthermore, pyrolysis of carbonaceous droplets or particles may also produce soot, as in oil burners or diesel engines [Charlson and Ogren, 1982]. All of these processes tend to produce primary particles with radii from 0.02 µm to 0.1 µm. The larger particles are emitted preferentially when the residence time and/or concentration of primary particles are sufficiently large to permit coagulation. Albeit the small initial particle size promotes coagulation, but at the same time the rigidity and inertness of the solid phase ensures that the specific surface area is maintained as the particles agglomerate. 3.1.4.1 Biomass burning Globally biomass burning is an important source of soot particles. Of the major biopolymers of woody tissue, soot formation is preferred from lignin (22 % and 29 % in hardwood and softwood, respectively), whereas holocellulose (cellulose and hemicellulose) primarily promotes volatilization of organic species during combustion [Mühlbaier and Williams, 1982]. As regards the type of burning, soot particles are primarily produced in intense flaming fires in which the oxygen supply is deficient, or parts of the flame become quenched before the oxidation of carbon radicals is complete. On the other hand, under smoldering conditions generally little soot is produced and the carbonaceous fraction of aerosol consists predominantly of tarry material2 [Andreae et al., 1998]. In addition, some charred material may become airborne by wind erosion long after the fire has been extinguished and thus may contribute to the atmospheric concentrations of coarse BC particles [Kuhlbusch et al., 1996]. Emission factors are usually available for the forms of carbon that can be readily measured and quantitated: either BC or EC. The mean emission factor for BC in savanna and grassland fires was found to be 0.59 ± 0.19 g kg−1 dry matter [Andreae et al., 1998]. For comparison, the emission factors for PM3.0 values were typically in the range of 2.2 to 4.1 g kg−1 dry matter as derived from airborne measurements. In broad agreement with these emission factors, the BC content of the savanna and grassland smoke aerosol for PM3.0 was found to vary between 10 to 27 %. While BC is expected to be largely conserved upon aging (which might not be the case, see discussion in sub2

Although these particles are not soot by definition, they can contribute to light-absorbing carbon (BC) to a variable extent.

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57

chapter 5.3), it is possible that semi-volatile compounds making up the bulk of the aerosol re-evaporate upon dilution of the smoke with ambient air due to reduction in their partial pressures [Liousse et al., 1995]. 3.1.4.2 Fossil Fuel Combustion A major source of submicron soot particles are emission from Diesel engines. Motor vehicle emission factors are often measured under different conditions, either on chassis dynamometers or road conditions. Much depends on the cycles used (steady-state versus transient), fuel characteristics as well as maintenance of the engines. The techniques for exhaust sampling also have a large impact on the measured values. Another problem is that since most regulatory measurements are made for total mass emissions, it is nearly impossible to obtain submicron particle-specific information from the data. On top of this, it was shown that in spite of legislative efforts on the road about 10 % of poorly maintained vehicles are responsible for over 80 % of total emissions, thus questioning the applicability of emission factors obtained from the very limited number of direct measurements involving wellmaintained engines under controlled conditions. All these factors make estimation of Diesel emission highly uncertain, both in terms of emission factors and total emissions. With this in mind, BC emission factor is estimated to be in the order of 2 mg kg−1 for the transport and domestic sectors in the developed countries, contrasting to the very low value for gasoline-powered vehicles (0.03 mg kg−1) [Cooke et al., 1999]. In the developing countries, emission factors for the transport sector and gasoline-powered vehicles are estimated to be a factor of 5 higher. Although not high in absolute terms, emission of soot by commercial air traffic at altitudes between 9 and 13 km is of utmost climatic relevance. The relatively long residence time of soot particles in the lower stratosphere and upper troposphere (50 days and 10–15 days, respectively), combined with the light-absorbing nature and possible IN-activity of soot and its role in heterogeneous processes, make it important out of proportion to the source strength [Petzold et al., 1999]. The mechanism of soot production in aircraft engines is very complex and poorly known, since it is influenced by the inhomogeneous flow and temperature field in the combustion system, the injection system and combustor technologies, as well as the type of fuel burned, to name just a few. The BC mass emission indices were found to be about 0.011 g kg−1 and 0.11 g kg−1 under cruise conditions for a modern and older engine, respectively [Petzold et al., 1999]. These values are in good agreement with those which can be estimated from number concentrations measured in follow-up flights

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Carbonaceous Aerosol

(0.016–0.064 g kg−1) [Anderson et al., 1998]. The older engine emitted larger particles than the modern one. For the latter, a primary BC mode at mass median diameter of 25 nm (σ= 1.55) and a coagulated mode at 150 nm (σ=1.65) were identified. The fleet average BC emission factor was calculated to be 0.038 g kg−1. The vertical distribution of BC emission and fuel consumption is shown in Figure 3.3.

Figure 3.3. Vertical distribution of aircraft-related BC emission and fuel consumption (after Petzold et al. [1999]).

During ground operation, take-off and climb up to 3 km altitude, 16 % of total fuel is consumed but 35 % of aviation-related BC is emitted, whereas 67 % of total fuel consumption and 40 % of BC emission occur in the altitude band of 9–13 km. 3.1.5 Atmospheric Observations of BC/soot 3.1.5.1 Surface Concentrations of BC As it was pointed out in the introduction, the determination of BC is loaded with several uncertainty factors which combine into a substantial overall uncertainty. Of them, the most important is the lack of standardized methodology of BC determination which makes measured BC concentrations difficult to compare, even on a relative scale. The biases and uncertainties in various optical methods of determination are treated in depth in sub-chapter 2.2, to say nothing of conceptual differences, e.g. when EC is determined with

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a thermal method. Therefore it would not really make sense to compile a long list of BC concentration data available from the literature, since they are not readily comparable and highly uncertain. Below, only a few reported values are given for illustration of the order of magnitude of typical BC concentrations in various environments. Concentrations of BC typically varied between 200 and 800 ng m−3 in rural regions of the Northern Hemisphere [Rosen et al., 1982]. In more remote areas, for example over Nova Scotia, Canada, surface BC concentrations were found to fall between 10 and 360 ng m−3, with an average value of 110 ng m−3. At the same site, higher BC concentrations (220 ± 30 ng m−3) were observed above the altitude of 1 km, suggesting that most of the BC observed at higher altitudes originated from long-range transport [Chýlek et al., 1999]. The BC to nss-SO42− mass ratios, on the other hand, were found to be independent of altitude. In background tropospheric aerosol at the Mauna Loa Observatory, Hawaii, BC concentrations of 2.5–12 ng m−3 were measured [Clarke et al., 1984]. These values were of similar magnitude than those suggested for remote oceanic areas, 5–20 ng m−3 [Rosen, 1984]. The exception was the Artic in winter, where BC concentrations could be as high as 300 ng m−3, accounting for 40 % of TC in late February [Rosen et al., 1982]. The lowest BC concentration ever detected in aerosol 1 ng m−3 was recorded in the South Pole [Hansen et al., 1988]. 3.1.5.2 Soot in the Upper Troposphere and Stratosphere Understandably, in the upper troposphere and stratosphere not BC mass concentrations which are measured, but instead, soot concentrations are inferred from electron microscopic observations of aerosol samples collected on grids. These methods may have their own bias: for example, very small soot particles are underrepresented due to the decreasing collection efficiency of wire impactors toward smaller particle size. In addition, when using scanning electron microscopy, soot entrained within sulfuric acid droplets cannot be detected, because it is a surface imaging technique. For the above reasons, these estimated values cannot be compared to BC mass concentrations reported for lower altitudes. A strong covariance was found between the distribution of observed soot concentrations and fuel usage data at 10–11 km altitude, indicating that a perennial soot layer exists at this altitude [Blake and Kato, 1995]. Nearly half of the total worldwide fuel amount (about 4.4 × 1010 kg) is used in the latitude band of 40° N–60° N, a substantial portion of which lies poleward of the polar night jet, in an area where the tropopause dips several kilometers toward

60

Carbonaceous Aerosol

higher latitudes. As a consequence, an estimated 10–40 % of the total soot emitted is deposited directly in the stratosphere. For the lower and higher estimates it yielded a soot concentration of 1.39 ng m−3 and 5.57 ng m−3, respectively. The average observed soot concentration in this latitude belt was closer to the lower estimate (1.73 ng m−3). The residence time of soot in the stratosphere, which depends on the assumed amount of soot injected, was 1 year for the lower estimate (probably more realistic) and 4 months for the higher one. Interestingly, at an altitude of 20 km, soot concentrations showed correlation with those of lower altitude. The total surface area of soot was in the same order of magnitude as that of sulfuric acid aerosol during volcanically quiescent periods (∼2.5 × 10−7 m2 m−3). Therefore soot particles in the upper troposphere and lower stratosphere can play an important role in heterogeneous reactions, as discussed in details in sub-chapter 5.3. 3.1.5.3 BC Concentrations in Cloud Water and Precipitation In cloud water of marine stratus clouds the range of BC concentrations was found to be between 8 and 80 µg kg−1 [Chýlek et al., 1999]. With a mean droplet radius of marine stratus clouds between 6 and 8 µm, these data implied that less than 10 % of the cloud droplets (and the corresponding CCN) incorporated BC particles. Similar calculations for BC aerosol suggested that 67–96 % of their number remain interstitial. In cloud water of stratocumulus clouds off the coast of southern California, BC concentrations varied between 23 and 79 µg kg−1 [Twohy et al., 1989]. The mass percentage of BC incorporated into clouds varied between 2 and 31 %, with an average of 9 %. This implies than about 90 % of the mass of BC was not mixed internally with soluble particles which are active CCN, i.e. BC was not effectively affected by nucleation scavenging. It should be noted that this conclusion referred to a polluted scenario, in harmony with those drawn from direct comparisons of εBC and εSO42−. The average BC concentration found in cloud water at Mt. Sonnblick, Austria, in 1997 was 1.07 µg ml−1, and accounted for 30 % of the TC [Hitzenberger et al., 2001]. The BC concentrations in cloud water correlated with those of SO42−, and to a lesser extent also with those of TC. These correlations may imply that the initial aerosol on which the clouds form was mostly internally mixed, at least in the size range where most of the CCN can be found. For the interstitial aerosol, on the other hand, BC/TC ratio was found to be 0.19 ± 0.24 on average. This might indicate that BC and OC were to some extent separated, with BC tended to be more closely associated with SO42− than the OC.

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At urban and rural locations in Canada, rainwater contained BC at average concentrations of 4.0 and 2.6 µg kg−1, respectively, without any trend with time in prolonged rain events [Chýlek et al., 1999]. It was suggested that incorporation of BC into raindrops primarily occurred through nucleation and/or in-cloud scavenging. In snow samples collected at the same urban site, BC concentrations were found to be higher (average 11 µg kg−1). This is possibly a consequence of the more effective below-cloud scavenging by snowflakes. 3.1.5.4 Historical Records of BC Concentrations For short-lived aerosol particles, such as BC, historical record of regional emissions is only available near the sources, in glaciers, back to only a few hundred years. The concentrations of water-insoluble OC and BC were determined from a high-alpine ice core covering the period between 1755 and 1982 [Lavanchy et al., 1999b]. Unfortunately, these data have been erroneously presented in the latest IPCC report, stating that the time-trend for organic carbon (OC) is given [Penner et al., 2001]. It should be stressed that WSOC was removed prior to measurements, so the data refer only to the water-insoluble and filterable OC fraction, i.e. the particulate OC fraction. Dating of the ice-core was performed with stratigraphic markers such as Saharan dust events, the atomic bomb horizon and identified volcanic eruptions. The uncertainty of dating was ∼2 years over the last 100 years, and ∼5 years for the period between 1755 and 1882. The Saharan dust events are often associated with significant concentrations of CaCO3, which causes positive artifacts in EC measurements. Since the mineral particle content of the ice core was several orders of magnitude higher than that of BC, correction had to be applied in the optical measurements, too. The time-trends of the measured concentration of BC, EC, (waterinsoluble) OC and TC in the ice-core between 1755 and 1975 are shown in Figure 3.4. As shown in Figure 3.4, there are large fluctuations in the concentrations in all records. The variations are more pronounced for more recent segments, due to the shorter time-averages reflecting the different efficiencies of vertical transport during summer and winter. Significant increase in concentration of all species can be observed after 1890, which compares well with the observed increase in sulfate concentrations from the same ice core [Doscher et al., 1995]. It is interesting to note that whereas the correlation between SO42− and TC was rather high (r2 = 0.43) for the period of 1890–1975, it dissipated for the period prior to 1890 (r2 = 0.02). The

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correlation between BC and SO42− concentrations was even higher (r2 = 0.56) for the more recent period.

Figure 3.4. Ice core concentrations (dashed line) and 10-year averages (solid line) of organic carbon (OC), elemental carbon (EC), BC, and total carbon (TC) for the time period 1755 to 1975. BC concentrations were calculated using σBC = 9.3 m2 g−1 (after Lavanchy et al. [1999b]).

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From the comparison between the corresponding EC and BC concentrations it was possible to derive the time-trend of the specific attenuation coefficient. The σBC varies between 2.1 and 25.2 m2 g−1, its mean value was 10.5 ± 4.8 m2 g−1 over the entire time period. The first half of the twentieth century experienced much higher σBC values (18–25 m2 g−1) than the nineteenth century (8–12 m2 g−1), possibly due to the change in dominant fuel in energy production in the region, from predominantly wood to coal combustion, around the end of the nineteenth century. The mean value, however, was similar to those obtained for high-alpine atmospheric aerosol (9.3 ± 0.4 m2 g−1) [Lavanchy et al., 1999a]. This agreement is somewhat suspicious since the assumed water-soluble coating on carbonaceous particles, which is understood to enhance light absorption, was removed in the ice-core measurements. The very high correlation between water-insoluble OC and EC implies common combustion sources for the two, and underlines the importance of the organic component of soot. 3.1.6 Global Models of BC The first global inventory of BC emissions was based on the inventory of global SO2 emissions and long-term monitoring data on ground-level ambient BC and SO2 concentrations [Penner et al., 1993]. Very high correlations can be expected if the sources of both BC and SO2 are sulfurcontaining fossil fuel combustion. This may be the case in source-dominated urban areas, where differences in removal efficiencies are not expected to influence the measured BC/SO2 ratios because of the proximity of the sources. In this scenario, the BC emission inventory can be derived by multiplying the known SO2 emissions with the region-specific BC/SO2 ratio. The estimated global BC emissions amounted to 24 TgC yr−1 (for the year 1980). The highest emission densities (> 200 kgC km−2 yr−1) were seen in China, Eastern Europe and the eastern coast of the United States. This BC inventory was significantly higher than earlier estimates based on conservative emission factors and data on the worldwide use of fossil fuels. With a prescribed dry deposition rate of 0.1 cm s−1 and removal coefficients of 2.5 and 0.7 cm−1 for stratiform and convective precipitation, respectively, the global distribution of BC mass concentrations was modeled on the basis of the emission inventory. The most notable deviation from observations was the severe underprediction of seasonal variations in the Artic (up to a factor of 5). Measured BC concentrations at the South Pole were similarly underpredicted by the model, possibly due to its incomplete biomass burning inventory.

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Another global model for the distribution of carbonaceous aerosol also addressed BC separately [Liousse et al., 1996]. In the model emissions from biomass and fossil-fuel burning were both considered. The biomass burning inventory comprised emissions from savanna, forest, agricultural and domestic fire sources, with emission factors estimated separately for fine particles, where most carbonaceous aerosol can be found. The detailed BC inventory is given in Table 3.2. Table 3.2. Inventory of the yearly emissions of black carbon particles and organic matter (after Liousse et al. [1996]). Products, Tg/yr Organic Matter, Tg of mass/yr Black carbon, Tg C/yr Biomass burning 44.6 5.63 Savannas 15.5 2.17 Tropical forests 16.6 1.93 Agricultural fires* 3.1 0.53 9.3 1 Domestic fuels# Fossil fuel 28.5 6.64 Natural sources 7.8 ----Total 81 12.3 * Agricultural fires: included are wheat, barley, rye, corn, rice, and sugar cane. # Domestic fuels: fuel wood, bagasse, charcoal, and dung.

In the model, biomass burning emissions were injected to the first 2000 m altitude, whereas other emissions were distributed up to the altitude of 1000 m. The prescribed removal coefficients (2.1 cm−1 and 0.6 cm−1 for stratiform and convective precipitation, respectively) were slightly different from those of Penner et al. [1993]. A single value was assigned for the scavenging ratio for BC particles, in spite of the indications that there are considerable differences between the hygroscopic behavior of BC particles derived from biomass burning and fossil-fuel combustion. The predicted annual zonal average concentrations of BC are shown as a function of altitude in Figure 3.5. An important feature that can be observed in the figure is that concentrations decrease steeply with altitude, except in polar regions of the Southern Hemisphere. The calculated total dry and wet deposition flux of BC were 3.0 and 9.3 Tg yr−1, respectively, and the total BC burden amounted to 0.13 Tg. The combination of these figures yields an average residence time of 4–4.5 days, in agreement with previous estimates [Ogren and Charlson, 1984].

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Figure 3.5. Predicted annual average and zonal average concentrations of black carbon (in ng m−3) (after Liousse et al. [1996]).

Another global inventory for BC originating from fossil fuel emission and biomass burning was developed and put into a global model of a resolution of 10 × 10 × 100 hPa [Cooke and Wilson, 1996]. An important difference compared to the model by Liousse et al. [1996] was that all BC was assumed to be hydrophobic when emitted, then transformed at a rate of 5 % per 2-hour time step into a hydrophilic form which wet deposited at the same efficiency as sulfate. Comparison of the model results with available measurement data revealed that the model overestimated surface mass concentrations at remote locations. Similarly, at altitudes between 5.5 and 11 km modeled BC concentrations (10–25 ng m−3) were consistently higher than observations (0.7–7.0 ng m−3) [Pueschel et al., 1992]. Even higher but random deviations were observed between predicted and observed BC concentrations in precipitation, which can be explained partly by the inability of the model to represent the subgrid-scale spatial and temporal variability of the precipitation events.

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The global average residence time for BC particles derived from the model was 7.85 days for the base case, but decreased to 5.57 days if hydrophilic BC emission was assumed. The calculated annual global BC emissions were 7.96 Tg from fossil fuel emission and 5.98 Tg from biomass burning, making a total of 13.94 Tg. Boreal forest fires, which were included in the model, were shown to contribute as much as 20–50 % to the observed BC concentrations in the summer Arctic. Cooke et al. [1999] incorporated their inventory of 1° × 1° resolution into a global model (ECHAM). Upon parameterization of the properties of carbonaceous aerosol, they assumed that 80 % of BC was hydrophobic, whereas OC was divided equally between hydrophilic and hydrophobic particles. The aging process of both BC and OC was represented by an assumed conversion of hydrophobic to hydrophilic aerosol with a timeconstant of 1.15 days. The dry deposition velocities for hydrophilic aerosol were assumed to be 0.025 cm s−1 and 0.2 cm s−1 over dry and wet surfaces, whereas hydrophobic aerosol deposited at a constant velocity of 0.025 cm s−1. The tropospheric burden of BC was found to be 0.077 TgC, and its estimated range was expected to increase to 0.15–0.25 TgC when biomass burning BC was included. The burden of OC was calculated to be 0.237 TgC, which was projected to be in the order of 1 TgC if biomass burning emissions and secondary production of OC was considered. The model predictions were compared with available measurements at remote locations. For Mace Head, the model results agreed well with the measurements, except in the summer months when overprediction by a factor of 2 occurred. This was likely due to the inadequate representation of atmospheric conversion from hydrophobic to hydrophilic aerosol. The Artic haze phenomenon was seriously underpredicted by the model, possibly due to incorrect representation of the transport processes. The agreement between the measured and modeled concentrations of BC was reasonable up to the altitude of 6 km, but broke down in the upper troposphere. Furthermore, the model underpredicted wet deposition of BC by 5–93 %. This finding may imply that BC emitted from fossil fuel burning was not completely hydrophobic. For the stratosphere, there was another model which predicted the distribution of pure soot particles, as shown in Figure 3.6 [Bekki, 1997]. The model was in qualitative agreement with measurements in that it predicted maximum concentrations between 50° N and 60° N, though quantitatively it was at the lower end of the measured mass concentrations.

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Figure 3.6. Model calculated 1992 annual average (a) concentration (ng m−3) and (b) surface area density (µm2 m−3) of black carbon soot as a function of latitude and pressure for run B (after Bekki [1997]).

A time-trend of fossil-fuel BC emission has been established based on estimates for six regions representing about two-third of present day emissions, and extrapolation to the rest of the world [Novakov et al., 2003]. This reconstructed emission trend is shown in Figure 3.7. Despite the uncertainties inherent in the estimations, the qualitative feature of the time-trend is believed to be realistic, showing a steep increase in the last decades of the nineteenth century, followed by a leveling off in the first half of the twentieth century then a boost after the Second World War, largely determined by the development of the South-East Asian region. Although not conclusive, the trend also indicates some downward trend for

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the last decade, contrary to the IPCC scenario which predicts further increase in BC emission for the next 50 years [Houghton et al., 2001].

Figure 3.7. Estimated fossil-fuel BC emissions (after Novakov et al. [2003]).

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3.2 Primary Organic Aerosol Primary organic aerosol particles are generally understood to be those particles which are released directly from various sources. However, because of the vast diversity of sources and emission processes, the actual definition may not be as simple as that. A part of primary organic aerosol may be released as individual particles which are recognizable by their characteristic morphology. Examples for these particles are bioaerosol particles or vegetation waxes from natural sources. Another fraction of primary organic aerosol components is released on primary particles which are not carbonaceous in themselves, such as surface active organic matter on sea salt aerosol. Internally mixed particles of predominantly primary origin could also be released by nucleation of inorganic species on carbonaceous cores, as it is the case in aircraft wakes. On the other hand, semi-volatile organic species can condense onto inorganic or soot particles to produce internally mixed primary particles, such as in smoke plumes. The formation of internally mixed primary particles raises the issue of how to make a clear distinction between primary and secondary organic aerosol in, for example, biomass burning plumes. In fact, there is a gradual transition between the combustion regime and ambient air within the plume which develops in space and time. Traditionally, in addition to particulate organic matter, primary organic species were understood to be those which are released into the gas phase by the sources then partition into the aerosol phase without previously undergoing gas-phase chemical reactions in the atmosphere. Conversely, compounds which are not released directly but form in photochemical reactions from precursors prior to partitioning into the aerosol phase are considered to be secondary aerosol components. The exception is if secondary processes are very fast and take place within seconds away from combustion, the aerosol formed may well be considered primary. This concept may help differentiate between primary-secondary organic aerosol on a molecular level, but it is absolutely unsuitable for visualizing bulk emission of primary organic aerosol from the burning process. According to the above definition, secondary organic aerosol formation could start in an air parcel of a plume well before condensation of primary species is complete. More importantly, however, this traditional approach ignores organic compounds which may be formed in multiphase reactions. Although originally such reactions were deemed unimportant in SOA formation, recently there have been mounting evidences that they can be significant, in close analogy with sulfate formation. For a detailed discussion of these processes see sub-chapter 3.3. Apart from these conceptual ambiguities,

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however, the major sources of carbonaceous aerosol are basically well understood. In this chapter the major types and sources of primary carbonaceous aerosol are discussed, with the exception of soot, to which a separate subchapter (sub-chapter 3.1) is devoted. However, the discussion will be restricted only to an illustrative description of the emission processes and their basic characteristics, as well as to the bulk characterization of the particles emitted. The primary reason for this is that there are excellent books available on these topics which the interested reader may consult with. In addition, other aspects, such as chemical composition or physical properties of the particles will be treated in depth in subsequent chapters. 3.2.1 Natural Primary Organic Aerosol Sources 3.2.1.1 Vegetation The direct emission of organic particulates by vegetation was first indicated in the early 1970s [Arpino et al., 1972; Schnell and Vali, 1973]. A possible source can be small particles of epicuticular wax which can be removed by certain stress conditions such as leaf expansion. Electron microscopic studies indicated that they are predominantly needle-like particles up to a length of 200 nm and a width of 30 nm [Beauford et al., 1975]. Such particles resemble wax rodlets which typically occur on leaf surfaces. Signatures of plant wax particles were found at a distance of 5000– 6000 km from the nearest sources [Gagosian et al., 1982]. 3.2.1.2 Bioaerosol Bioaerosol is defined as airborne particles, large molecules that are living, contain living organisms or were released by living organisms. Key viable atmospheric particles include fungi, bacteria, pollen, algae, yeasts, molds, mycoplasma, viruses, phages, protozoa and nematodes. Bioaerosol particles also include non-viable cells, as well as cell fragments which are usually several times smaller than the original cells or spores. The size of the bioaerosol particles may vary from 10 nm to 100 µm. The shape of the bioaerosol particles can be extremely diverse, for example spherical, dodecahedral, needle-like and flakes. Many viruses are pleomorphic and change their shapes. Among all bioaerosol particles, bacteria play the most important role in atmospheric processes. Viruses are too small to contribute measurably to

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organic aerosol mass concentrations. On the other end of the scale, pollens have very low number concentrations and are confined in a size range which has considerable settling velocity and consequently relatively short residence time to be atmospherically significant. The presence of airborne bacteria has classically been shown by colonial growth on culture media, though most airborne bacteria are either dead or viable but not culturable. Epifluorescent microscopic methods of counting total bacteria revealed that only 0.02–10.6 % of their total number were actually culturable [Lighthart, 1997]. A typical size distribution of airborne bacteria is shown in Figure 3.8.

Figure 3.8. Bargraph showing the aerodynamic particle size category distribution of bacteria associated particles in the outdoor atmosphere (left panel) and the size disparity between nominal single bacterial cells (ca. 1 µm) and the large bacteria associated particles in the outdoor atmosphere (right panel) (after Lighthart [1997]).

The count median diameter of bacterial aerosol was about 3.6 µm AED. The reason for sizes larger than one bacterium (∼1 µm AED) is that bacteria are usually adherent on particles such as amorphous plant, fungal or soil debris, or as bacterial-slurry residue of quasi-spherical shape upon evaporation. The great majority of the genera are Gram-positive (73–90 %), the Gram-negative genera are represented by pseudomonads and xanthomonads [Lighthart, 1997]. Most airborne bacteria are primarily derived from plants, and to a lesser extent from soil surfaces, though virtually any surfaces, including water surfaces can be sources. The mechanism of their entrainment into the atmosphere from terrigeneous sources are thought to be either direct wind

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action or indirect, abrasive dislodgment by leaf-to-leaf contact. Under nonwindy conditions, thermal convection and thermal or photophoresis may be at work. From water bodies they are emitted primarily by bubble bursting and wave action. Most bioaerosol particles are released by plants into the atmosphere as pollen and spores. Other sources include human activities such as industry (textile mills), agriculture (fertilizing), and municipal activities (sewage plants). The atmospheric number concentration of culturable bacteria is expressed in CFU m−3 (colony forming unit), and typical values are in the order of 200 and 600 for rural and urban atmosphere, respectively [Schaffer and Lighthart, 1997]. In polar and marine atmosphere they can be as low as 0–18 and 1–31 CFU m−3, respectively [Pady and Kelly, 1953]. The most prevalent genus at all continental locations was Bacillus (12–45 %). Since the atmosphere is a harsh environment for bacteria, primarily due to dessication and solar radiation, it is not surprising that a highly resistant spore-former represented the most prevalent genus [Schaffer and Lighthart, 1997]. Over the Southern Ocean, however, mainly rod-shaped and presumably Gram-negative bacteria were observed by electron microscopy [Pósfai et al., 2003]. Although these genera were possibly ejected by the bubble-bursting mechanism that generates sea-salt particles, they were found not aggregated with sea-salt aerosol. Bioaerosol particles larger than 0.2 µm in radius were studied in urban aerosol [Matthias-Maser and Jaenicke, 1994]. For particles with equivalent radii larger than 2 µm light microscope was used in combination with staining with a protein dye, whereas in the size range of 0.2 and 2 µm scanning electron microscope was deployed equipped with an energy-dispersive X-ray detector. There exist particles—such as phytoliths—which are biogenic but present in the atmosphere in mineralized form. These particles are not assigned as bioaerosol particles by the staining method. There were basically two types of bioaerosol particles found based on morphology, elemental composition and behavior upon irradiation. The first included rod-like, elongated or curved particles with primarily background EDX-spectra, sometimes with traces of P, S, K or Ca. They shrank or even vanished upon irradiation leaving a residue or an empty sheath. The second group included special shapes such as spheres with spines or dents, showing high background EDX spectra sometimes with peaks of Si, P, S, K, and Ca, but no change upon irradiation [Matthias-Maser and Jaenicke, 1994]. In the aerosol studied, bioaerosol particles contributed about 30 % to the number concentration of all particles above 0.2 µm equivalent radius. The carbon contents of fungal spores typically range between 42 and 66 % (average 51 %) on a dry mass basis [Bauer et al., 2002]. This translates into

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an average value of 13 pgC per spore or 0.38 pgC µm−3 of spore. By determining number concentrations of fungal spores in air at Mt. Rax, Austria it was calculated that they accounted for 2.4 % to 5.4 % of total carbon in the size range between 2 and 7 µm. 3.2.1.3 Soil Soil may be an important direct source of organic aerosol, though little is known about the organic carbon content of soil-generated aerosol. In Saharan dust collected off the coast of West Africa only ∼3 % (m/m) OC was found [Lepple and Brine, 1976]. Apart from the bulk composition, there might be a chemical fractionation of elements by soil particle size. For example, substantial enrichment of organic carbon was observed in the relatively organic-poor soil-derived aerosol from dust storm areas in southern high plains in Texas [Zenchelsky et al., 1976; Delany and Zenchelsky, 1976]. While organic carbon contributed only 0.2–4 % of the bulk mass concentration of total suspended particulates, their contribution in the size range of 0.4–1 µm was found to be as high as 5–20 %. It should be noted that carbonate is typically present in mineral dust and was found to be a significant component of fine particulate matter in some locations during the ACE-Asia experiment [Andronova et al., 1993; Gomes et al., 1993]. The mechanism of selective enrichment of organic matter compared to mineral dust was postulated in studies on the sublime and poorly understood dust haze events observed in Australia [Boon et al., 1998]. The high concentrations of organic matter associated with background dust (10–60 µg m−3) were suggested to originate from the wind erosion of rural soils in spite of the fact that Australian soils are distinctive for their very low organic content. The high relative share of organic matter in dust could reflect the socalled winnowing effect of winds which preferentially entrain low mass organic matter rather than the heavier mineral fraction of dust. This postulated mechanism was also supported by the fact that in the case of occasional dust entrainment events, accompanied with much higher mass concentrations (up to 130 µg m−3), the mass fraction of organic matter was usually lower since stronger winds are less effective in winnowing out organic matter from soil. The global flux of organic carbon from soil was estimated to be 11 TgC yr−1, mostly in coarse particles [Duce, 1978]. In the tropics alone, on the basis of stable carbon isotope studies soil erosion was suggested to release fine carbonaceous aerosol at a flux of 0.8 TgC yr−1 [Cachier et al., 1985].

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3.2.1.4 Ocean The presence of surfactants, such as protein degradation products in marine aerosol was shown as early as 1964 [Blanchard, 1964]. It was found that surface-active organic material is highly enriched on the sea-salt aerosol particles relative to the bulk seawater (up to a factor of several thousands). Although part of this enrichment can be due to accumulation of surface-active organic material on the bulk air-water interface, it was postulated that the major part was from organic material adsorbed or scavenged by the bubbles as they rise toward the water surface. Since the ejection heights of the various jet drops depends on their kinetic energy which in turn derive from the bubble surface free energy, suppression of the surface free energy by dissolved surface-active material lowers the ejection height which can be observed experimentally [Blanchard and Syzdek, 1974]. Aerosol particles of 70–500 nm were observed over open water between ice floes in the Arctic Ocean, on calm, sunny days [Leck et al., 2002]. It was speculated that these particles results from film drops ejected by bursting bubbles which have been released by decomposition of the organic matter or respiration of ice algae. Alternatively, air bubbles released as ice melts could provide another source at the sunlit edges of the leads. These film drops consisted mainly of organic material which—being surface active— concentrates in the surface layer in the open leads. It was observed that these particles generally had lower CCN activation threshold than (NH4)2SO4 at a given supersaturation, which is in accordance with the presence of surfactant material on the surface of the particles. This assumption seems to be corroborated by measurements of potassium, known to be associated with organic material. Particles with diameters below 130 nm showed an enhancement of K+ concentration by a factor of 30 relative to that measured on cloudy days. Even at wind speed above 12 m s−1, when usually jet drops predominated, organic material still contributed to aerosol volume concentration up to an estimated 20 %. 3.2.2 Anthropogenic Primary Organic Aerosol Sources 3.2.2.1 Biomass Burning Albeit smoke from biomass burning has ancient historical record contrary to fossil fuel combustion which dates back only to the medieval times, the latter has attracted much more scientific interest for a long times. Biomass burning serves a variety of purposes, such as clearing of forests and

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brushlands for agricultural use, control of pests and weeds, prevention of litter accumulation to preserve pastures, production of charcoal, for cooking and heating, among other things. Two types of forest clearing is practiced in the tropics: shifting agriculture, where the land is cultivated for a few years then allowed to return to forest vegetation during a fallow period, and permanent conversion of forests to pastures or croplands [Crutzen and Andreae, 1990]. Originally, crop and fallow periods lasted for 2–3 and 10–50 years, respectively, but the growing population and shrinking forest area have shortened the fallow period to the extent that the land cannot recover to its productivity, causing shifting agriculture to decline. During the dry season, the undergrowth is cut and trees are felled then left to dry for a few months to improve burning efficiency then set on fire. The efficiency of the first burn is variable but relatively low, since large trunks containing most of the biomass remain largely unaffected by the first burn. The remaining unburned biomass is either left to rot or collected and set on fire again. Altogether, an estimated 40 % of the biomass is consumed, amounting to 500–1000 TgC each year [Crutzen and Andreae, 1990]. Tropical savannas and brushlands consist of a largely continuous layer of grass interspersed with trees and shrubs, are burned every 1 to 4 years during the dry season with highest frequency in the humid savannas. There only dry grass, litter, weeds and shrubs are burned, whereas the larger trees of fire-resistant species suffer little damage. In savanna fires 300–1600 TgC are burned each year. In the developing countries, fuel wood and agricultural wastes are the primary energy sources for cooking and domestic heating. In the developing countries an estimated 35 % of energy derives from biomass fuels, burned in quantities of 600–1200 TgC yr−1, assumed to be distributed evenly between firewood and dung and crop residues [Crutzen and Andreae, 1990]. Burning of agricultural wastes in the field exposes annually some 300–600 TgC of firewood and 500–800 TgC of agricultural wastes to fire, mostly in the developing countries. The prescribed burning practiced in North America and Australia and the wildfires occurring sporadically in temperate and boreal forests, are estimated to contribute much less (150–300 TgC) to the amount of carbon burned annually. The global significance of biomass burning lies not only in the amount of primary particles released, but also in the areal extent they spread out in the atmosphere. The smoke plume rises in the atmosphere and dilutes with ambient air. Clouds are formed frequently on the plume but usually reevaporate without causing precipitation. Upon loosing buoyancy, the plume drifts horizontally in relatively thin layers with the prevailing winds, extending over a thousand kilometers or more. During the dry season the

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height of this layer is usually limited by the trade wind inversion at about 3 km in the tropics. The smoke-laden air then disperses by the large-scale circulation, which is dependent on the season and the region from which the plume originates. In tropical Africa, the plumes usually travel in a westerly direction toward the equator. Upon approaching the Intertropical Convergence Zone, the intensifying vertical convection destroys the layered structure causing the plume to dissipate. A part of the smoke may reach the middle and upper troposphere, perhaps even the stratosphere. Plumes from the biomass burning regions in South America usually travel south or southeast, because of the effect of the Andes on large-scale circulation. Here again they may become entrained in the seasonal (austral spring) South Atlantic Convergence Zone [Crutzen and Andreae, 1990]. a) The process of biomass burning Biomass burning begins with a drying/distillation step, releasing water and volatile species, followed by pyrolysis. Pyrolysis results in the formation of char of high carbon content, tar of intermediate molecular weight and volatile compounds in the form of a flammable white smoke. Above 450 K the process becomes exothermic and at about 800 K glowing combustion begins, accompanied with the release of tar and gaseous products. When these substances are diluted with air and ignite, flaming combustion occurs. When the supply of volatiles depletes in the near-surface region of the fuel, flaming combustion ceases and smoldering phase begins. The temperature in this stage usually remains below 850 K, and vast amounts of partially oxidized pyrolysis products are emitted. The primary particles emitted by smoldering combustion are primarily tan organic droplets and have a very low BC content [Patterson et al., 1986]. Open vegetation fires are typically dynamic fires, with a moving fire front passing over an area of either savanna or forest. In this case, all combustion stages are present at any time, and the plume includes their combined emission. It was shown that the vast majority of emissions (> 99 %) takes place within 4–8 hours after ignition [Andreae and Merlet, 2001]. b) Emission factors for various types of biomass burning In biomass burning inventories emissions are typically represented in the form of spatiotemporally resolved fields, where emission per unit area and time is provided at specified spatial and temporal resolution. Emission data

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can be given in the form of emission ratios, when the excess concentration of a species in the plume is reported with reference to the excess concentration of a simultaneously measured reference gas, CO2 or CO. For aerosol, these emission ratios are usually given in unit aerosol mass per kg of CO2 carbon [Andreae and Merlet, 2001]. An advantage of emission ratios is that knowledge of the fuel composition or the amount of fuel burned is not required, so they are ideal for airborne studies where such information is not available. Another frequently used parameter is the emission factor, which is the amount species released per the amount of dry fuel consumed, expressed in units of g kg−1. The carbon content of biomass varies only over a fairly limited range (typically between 37 % and 54 %), so the emission factors for carbonaceous aerosol are primarily a function of the type of the combustion process. The mean total particulate (TP) emission factors were takes as 8.1 ± 5.5 g per kg of fuel and 18 ± 10 g per kg of fuel for savanna and forest fires, respectively [Liousse et al., 1996]. In the developing countries, nearly half the fuel wood is estimated to be burned for heating and cooking, while the other half is for charcoal making. The efficiency of burning for charcoal making was assumed to be ∼20 %, with particulate emission factor of 17 g kg−1 of fuel. If, however, the charcoal is burned in the households, the assumed emission parameter is 10 g per kg of fuel [Liousse et al., 1996]. c) Global estimates of POC from biomass burning Quantitative estimates on smoke emissions from biomass burning can be retrieved from satellite observations. Detection of smoke plumes is possible over low-albedo surfaces1 such as water or forests, by the visible (0.58–0.68 µm) or near infrared (0.725–1.1 µm) channel of the satellite [Liousse et al., 1997]. The yearly flux of carbonaceous particles emitted by savanna fires was estimated on the basis of satellite observations. Taking into account that in the region 85 % of the total suspended particulate is carbonaceous aerosol, the annual flux was estimated to be 6.5 TgC. Of this, BC emission is about 1 Tg, given that the BC/TSP ratio is found to be about 10 %. These figures are accompanied by a factor of 5 uncertainty. This satellite-based estimate agrees fairly well with the value obtained from ground-based measurements [Cachier et al., 1995].

1

Over surfaces of low-albedo (A < 0.1), the backscattering of the incident radiation by the atmosphere dominates, so in the presence an aerosol layer an increase in albedo can be observed, i.e. ∆A is positive.

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A combination of the biomass burned and estimated emission factors, the global emission strength from biomass burning was 30–100 TgC yr−1 and 6–30 TgC yr−1 for particulate OC and BC, respectively [Crutzen and Andreae, 1990]. A detailed inventory of emission factors and global emission of TC, OC and BC from biomass burning is provided for various ecosystem types, as listed in Table 3.3 [Andreae and Merlet, 2001]. Table 3.3. Global emission of selected pyrogenic species in the late 1990s (in mass of species per year; Tg a−1) (after Andreae and Merlet [2001]).

TC OC BC

Savanna and grassland

Tropical forest

Extratropical forests

Biofuel burning

Charcoal making

Charcoal burning

Agricultural residues

Total

11.7 10.6 1.5

8.7 7.0 0.88

5.3 5.8 0.36

14.0 10.7 1.6

-------

0.24 0.18 0.06

2.2 1.8 0.37

42.2 36.1 4.8

3.2.2.2 Fossil Fuel Burning Globally, coal burning is one of the major anthropogenic sources of primary organic particles. When burning bituminous coal the initial burning phase is dominated by the combustion of devolatilized organic matter, giving rise to highly absorbing, dark, thick smoke and high flames [Bond et al., 2002]. Average particulate matter emission factor was found to be 12 ± 17 g kg−1. It is possible that upon adding coal chunks to the fire, the chunks become hot enough for devolatilization to occur, but the temperature remains below that needed for the combustion of the volatile organic matter. Under these circumstances, called “lukewarm ignition”, mostly tar escapes in the form of a thick cloud of light-absorbing smoke, practically without fire. Typically, these periods of high emissions usually accompany with rather weak light absorption in most cases. The size-distribution of primary emissions from domestic coal burning is shown in Figure 3.9 [Bond et al., 2002]. The size distribution for bituminous coal combustion shows significant primary emission into the accumulation mode, preferentially occurring during the gas-phase combustion of volatiles. Similarly, emission from lignite combustion tends to have maximum number concentration in the accumulation size range and not in the nucleation mode, presumably for lack of high-sooting orange flames. In contrast, coal briquette combustion typically releases very small particles that do not interact efficiently with solar radiation, explaining the low scattering and absorption from this type of burning.

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Figure 3.9. Distributions weighted by the amount of fuel burned. All concentrations are in the undiluted exhaust gas (after Bond et al. [2002]).

The devolatilized matter during combustion is partly taken up by the clay binder of the briquette. Furthermore, the clay serves as a catalyst to crack the coal tar to carbon and hydrogen. Among the large particles emitted by high-temperature fuel oil combustion, there are large spheroidal carbonaceous particles of porous or spongy morphology which are unambiguous markers for industrial emissions in sediment records since they are not susceptible to post-depositional alterations [Rose et al., 2003]. Although these particles may occasionally be transported over large distances, their atmospheric effects are insignificant. Emission of primary organic particles from aircrafts is of high importance since they are released in the upper atmosphere where they have long residence time and can participate in important atmospheric processes. It was shown that a large number of ultrafine particles can be released in the wake [Yu et al., 1999a]. When low-sulfur fuel is used in the aircraft engines, organic compounds in the exhaust dominate the mass of the volatile ultrafine particles. Even in the case of medium to high sulfur content, the role of organic species can still be significant. These processes can explain the observed organic particulate emission index of 23 mgC kg−1 of fuel. The mechanism by which these ultrafine particles may form was also postulated. Aldehydes (mostly formaldehyde and acetaldehyde) detected in aircraft exhaust have larger proton affinity than water, consequently they readily react with hydronium to form protonated core ions in exothermic reactions. Polar organic molecules and water molecules are likely to attach to these positive ions, forming charged clusters such as CH2OH+(CH2O)n(H2O)m. On the other hand, the electrons produced via chemiionization are captured by

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O2 molecules to form O2−, which subsequently reacts with sulfur and nitrogen species to form the stable negative ions of NO3− and HSO4−. By further uptake of water, H2SO4 and HNO3 molecules cluster ion such as HSO4−(H2SO4)n(H2O)m are formed. Due to the electrostatic effect, the charged nascent aerosol grows much faster than neutral aggregates [Yu et al., 1999a]. When such positive and negative ions recombine, mixed acid/organic microparticles are formed in which internal chemical reactions may take place. Sulfuric acid is known to react with aldehydes and alkenes to yield highly polar compounds of low volatility. Such organic species can also be taken up from the gas phase due to their high acid solubility and reactivity. The chemiion-induced volatile aerosol mode can bifurcate into two submodes, associated with positive and negative ions. Both modes consist of volatile particles of mixed composition, and the bifurcation is most pronounced when exhaust sulfur emissions are low and organic emissions are more prominent [Yu et al., 1999a]. A detailed inventory of primary carbonaceous aerosol emission from fossil fuel combustion was presented based on fossil fuel consumption data from the United Nations (for the year 1984), and emission factors for various combustion processes [Cooke et al., 1999]. The differences in emission factors of the same combustion processes were taken into account for countries classified into three stages of development. Domestic and industrial combustion processes were also considered separately: the lower temperature and lack of control leads to a higher production of (mostly organic) aerosol particles in domestic combustions, whereas in industrial combustion processes the BC emissions are relatively more important. A valuable feature of the inventory was the separation of bulk and submicron emissions, though size distribution data are mostly available for coal and diesel combustion only. The range of estimated OC emission factors for the major source types are summarized in Table 3.4. The global emission fluxes based on the bulk emission factors were 10.1 Tg yr−1 for OC, from which submicron aerosol emission represented 7.0 Tg yr−1. The “hot spots” appeared to be East-Asia, which accounted for 35–40 % of the global emissions. The EU and North America had similar share in the emissions (7–9 and 6–10 %, respectively). Coal combustion dominated global carbonaceous emissions, even for the submicron aerosol, as shown in Figure 3.10.

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Table 3.4. Emission factors for bulk and submicron organic carbon aerosol (after Cooke et al. [1999]). Fuel type Hard coal, hard coal briquettes, coke oven coke, gas coke, brown coal coke Lignite brown coal, lignite briquettes Peat Peat briquettes Diesel

Aviation gasoline Jet fuel Kerosene Liquid petroleum gas Residual fuel oil Motor gasoline Natural gas and other gases∗ Hard coal, hard coal briquettes, coke oven coke, gas coke, brown coal coke Lignite brown coal, lignite briquettes

Sector

Underdeveloped, Semideveloped, g kg−1 g kg−1 Bulk OC Emission Factors combined 3.09 1.76 domestic 9.54 9.54 industrial 1.0 0.325 combined domestic industrial combined domestic industrial transport domestic industrial all all all all

11.5 24.5 5.4 2.71 6.07 1.21 5.0 1.0 0.25 1.15 0.45 0.022 0.02

Developed, g kg−1 1.08 5.83 0.2

6.6 24.5 1.77 2.71 6.07 1.21 5.0 1.0 0.07 1.15 0.45 0.022 0.02

4.02 15 1.08 2.71 6.07 1.21 1.0 1.0 0.05 1.15 0.45 0.022 0.02

all 0.1 all 0.73 combined 2.2 domestic 11.2 industrial 0.432 Submicron OC Emission Factor combined 2.29 domestic 4.77 industrial 1.10

0.1 0.73 2.2 11.2 0.432

0.1 0.07 2.2 11.2 0.432

1.19 4.77 0.298

0.66 2.92 0.149

combined domestic industrial

4.45 12.3 1.61

2.46 7.5 0.804

8.55 12.3 5.94

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Figure 3.10. Relative contribution of various sources to global carbonaceous aerosol emission from fossil fuel sources (after Cooke et al. [1999]).

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3.3 Secondary Organic Aerosol (SOA) The formation of secondary organic aerosol from volatile organic precursors is an important process in the troposphere, especially in regions where photochemical ozone formation is significant. On one hand, it represents a major sink for the semi-volatile photooxidation products of a suite of anthropogenic hydrocarbons and biogenic volatile organic compounds (VOC), through which organic species are eventually removed from the atmosphere. At the same time, this process is certainly an important yet poorly characterized source of fine aerosol particles. The earliest observation of particle formation from irradiated petroleum fractions dates back to as early as 1873 [Fudakowski, 1873], and was resuscited for a passing moment in the work of Mader et al. [1952]. A hypothesis on the origin of “blue haze” on sunny summer days over forested areas was first put forward by Went in 1960. He suggested that the photooxidation of biogenic hydrocarbons emitted by forests gives rise to a large number of small particles which makes the “blue haze” visible. The earliest smog characterization studies conducted in the South California Air basin were the first to recognize secondary organic aerosol as an important fraction of aerosol carbon. Using the chemical element balance method 50 % of total carbon was assigned to be secondary, produced by gasto-particle conversion [Friedlander, 1973]. For the much larger Aerosol Characterization Experiment (ACHEX), the same method yielded a SOA estimate as high as 80 % of total carbonaceous aerosol [Gartrell et al., 1980]. In subsequent studies, other methods based on measured EC/OC ratios gave consistently lower, though still important SOA contributions (see sub-chapter 3.3.5.1). Interestingly, the most recent SOA estimates outside the South California Air basin tend to indicate even higher share of secondary organic species than the very first approaches. Another impetus to studies on secondary organic aerosol formation came from the understanding of secondary inorganic aerosol formation in the 1970s. The pioneer in the chemistry of secondary organic aerosol formation from anthropogenic precursors was Grosjean [1975]. The breakthrough in the chemistry of biogenic secondary aerosol formation may be awarded to Kamens et al. [1981], about two decades after the publication of Went’s idea. Since the early 1990s, inspired by the pioneering work by Novakov and Penner [1993] on the atmospheric significance of organic aerosol, research into the chemistry of secondary organic aerosol has boosted and significant advances have been made in the fields of reaction mechanisms, partition theories and modeling of secondary organic aerosol formation. New analytical instruments have been called in to identify low-volatility reaction

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products which are more likely responsible for new particle formation, and field observations of many reaction products have also been progressed. This chapter will treat secondary organic aerosol from anthropogenic and biogenic precursors separately all the way through to the last but one sub-chapter in which they are combined into global source estimates. Secondary organic aerosol formation in heterogeneous and multiphase reactions will be treated separately in the last sub-chapter. There are two reasons for that: the first is that these mechanisms do not fit into the conventional concept of SOA formation, since their products do not need to change phase by gas-particle partitioning to produce SOA (though usually their precursors need). Secondly, our knowledge on such processes is still rudimentary, and these mechanisms have not been incorporated yet into models of SOA formation. 3.3.1 Emissions of Volatile Organic Compounds (VOC) relevant to Secondary Organic Aerosol Formation Thousands of volatile organic compounds (VOC) are emitted into the atmosphere from various biogenic and anthropogenic sources. Most of them are involved in tropospheric photochemistry, thereby controlling OH concentrations over most of the troposphere. Thus they indirectly affects the mixing ratio of important greenhouse gases, such as methane and carbonmonoxide, therefore they also have an impact on climate. Detailed discussion of the sources, chemistry and atmospheric effects of these species is a subject of recent books (e.g. Reactive Hydrocarbons in the Troposphere, C. N. Hewitt (ed.), Academic Press, 1998). Fortunately, however, of this vast array of VOC only a small suite of compounds is of relevance in secondary organic aerosol formation. The principles by which these compounds are selected will be discussed in details in the subsequent sub-chapters. Consequently, this sub-chapter will implicitly concentrate only on the volatile organic species relevant in SOA formation. 3.3.1.1 Emission of Biogenic VOC Natural sources of VOC include oceans and fresh water, soil and sediments, microbial decomposition of organic litter, geological hydrocarbon reservoirs, plant foliage and woody material. Guenther et al. [1995] compiled best estimates of natural volatile organic compounds emissions from oceans and plant foliage on a global gridded basis, an inventory which has been widely used ever since in atmospheric chemistry models. Emissions of VOC

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85

from other sources were neglected for lack of reliable data, but were thought of representing only a few percent of overall global VOC emission [Zimmerman et al., 1979]. It should be noted that whereas this may be true on a mass basis with respect to all VOC emitted, it is highly uncertain how this would affect the emission inventory of the fraction of natural VOC that is relevant in SOA formation. Among the several hundred volatile organic compounds emitted by vegetation, only monoterpenes, sesquiterpenes and their derivatives are thought to contribute to SOA formation. These chemical classes were restricted further to only 14 individual species which likely dominate secondary biogenic organic aerosol formation [Griffin et al., 1999a]. These compounds include monoterpenes of chemical formula C10H16, α-pinene, βpinene, ∆3-carene and sabinene, bicyclic olefins that differ in the location of the double bond and the number of carbons associated with the secondary ring, cyclic diolefins limonene, α-terpinene, γ-terpinene and terpinolene that differ only in the location of the second double bond, acyclic triolefins myrcene and ocimene with a third double bond in different location, as well as two sesquiterpenes, β-caryophyllene and α-humulene of chemical formula C15H24, and two oxygenated terpenes, linalool and terpinene-4-ol. The chemical structures of these species are shown in Figure 3.11. The sesquiterpenes and oxygenated terpenes, as well as certain other species were classified into “other reactive volatile organic compounds (ORVOC)1 in the global emission inventory of natural VOC compiled by Guenther et al. [1995]. These were identified as compound classes rather than individual compounds, and included terpenoid alcohols, C7–C10 ncarbonyls, aromatics, sesquiterpenes, terpenoid ketones, and higher olefins. In this inventory, foliar emission fluxes are calculated as [Guenther et al., 1995]: F = Dεγ (3.2) where D is the foliar density (kg dry matter m−2), ε is an ecosystem dependent emission factor (µgC m−2 h−1 at a photosynthetically active radiation (PAR) flux of 1000 µmol m−2 s−1 and leaf temperature of 303.15 K), and γ is a dimensionless activity factor accounting for the influence of PAR and leaf temperature. Values of ε were estimated to vary between ecosystem types in the range between 0.2 and 2.4 µgC g−1 h−1 in five discrete steps, the default being the lowest value.

1

By definition these compounds have lifetimes shorter than 1 day in the troposphere.

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Figure 3.11. Chemical structures of the biogenic hydrocarbons relevant in SOA formation. Bonds between carbon atoms are shown with vertices representing carbon atoms; hydrogen atoms bonded to carbon are not explicitly shown (after Griffin et al. [1999a]).

For ORVOC, an emission factor of 1.5 µgC g−1 h−1 was adopted from Guenther et al. [1994]. Monoterpene emission rates are primarily controlled by their vapor pressure which is determined by temperature. The relationship between monoterpene emission rate and temperature can be described as [Guenther et al., 1995]:

γ = exp[β (T − Ts )]

(3.3)

where β (K−1) is an empirical coefficient, T is the leaf temperature, and Ts is leaf temperature at 303 K. The value of β was assigned to be 0.09 K−1, which

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87

was shown to be valid for a range of individual compounds and plant species [Guenther et al., 1993]. Although monoterpene emission rates from some plants were shown to be sensitive to light intensity, this was not included in the calculations for lack of sufficient data and adequate parameterization. Even the age of the needles or relative humidity may have a large impact on monoterpene emission rate. For example, emission of linalool from orange trees increased by an order of magnitude during blooming [Arey et al., 1991]. Physical injuries of plants may also trigger large increase in monoterpene emission. Of the annual global flux of VOC of 1150 TgC estimated for the year 1990 by the global model, the monoterpenes represented 127 TgC, whereas ORVOC 260 TgC. The major fraction of the total flux was attributed to isoprene (503 TgC), and the balance was due to the class of other volatile organic compounds2, neither of which were thought to contribute to secondary organic aerosol formation. As far as the ecosystems are concerned, woodlands were estimated to represent ¾ of global monoterpene emissions (95 TgC yr−1), and another significant source was shrub (25 TgC yr−1). The emission of monoterpenes followed a latitudinal pattern, the two most important parameters being biomass and temperature. In July, the highest emission rates for monoterpenes were predicted for the western United States, eastern Canada, central Europe, and parts of the Amazon basin. In January, maximum monoterpene and ORVOC emission rates were calculated for tropical woods in South America and Africa. The calculated monoterpene emission rates were lower than those reported in earlier estimates, up to a factor of 6. The flux estimates of Guenther et al. [1995] were not corrected for chemical and deposition losses within the canopy, since a reasonable agreement was found between above-canopy concentration gradient and leafenclosure measurements. It is anticipated that the typical tropospheric lifetime of most VOC (>1000 s) is much higher than the time-scale of turbulent diffusion in the canopy (<100 s). Exceptions can be nighttime conditions when the time-scale of diffusion could be much larger, or βcaryophyllene and linalool, which were shown to suffer substantial losses within the canopy already as aerosol particles due to their exceptionally high reactivity. The flux estimates of Guenther et al. [1995], together with those given by Müller [1992], for selected species and compound classes capable of forming secondary organic aerosol, were compiled into a speciated relative emission inventory by Griffin et al. [1999c], as shown in Table 3.5.

2

With lifetimes shorter than 1 day.

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Table 3.5. Estimated species contribution to global emissions (after Griffin et al. [1999c]). MONOTERPENES ORVOC Monoterpene1 Contribution ORVOC2 Contribution α-Pinene 35% Terpenoid Alcohols 9% β-Pinene 23% C7-C10 n-Carbonyls 7% Limonene 23% Aromatics 6% Myrcene 5% Sesquiterpenes 5% Sabinene 5% Terpenoid Ketones 4% ∆3-Carene 4% Higher Olefins 1% Ocimene 2% Terpinolene 2% α- & γ-Terpinene 1% 1 Guenther et al. [1995] estimated a total monoterpene emission rate of 127 TgC yr–1, Müller [1992] estimated 147 Tg yr–1. 2 Only those capable of forming aerosol are included; Guenther et al. [1995] estimated an ORVOC emission rate of 260 TgC yr–1; Müller [1992] estimated 94 Tg yr–1.

3.3.1.2 Emission of VOC from Anthropogenic Sources Although a large number of volatile organic compounds are emitted from anthropogenic sources, of them only the aromatic and large aliphatic species were found to generate significant secondary aerosol [Griffin et al. 1999b; Hallquist et al., 1999; Kamens et al., 1999]. Since the total emission flux of VOC from anthropogenic sources is inferior to that emitted by the biosphere, to date very few detailed inventories are available. Among them the most frequently used inventory is the compilation by Piccott et al. [1992], based on global activity data for 28 specific source groups. These included combustion of various types of fossil fuels, various industrial activities as well as important biomass burning sources. The total VOC flux was estimated to be 110 Tg yr−1, much lower than that of biogenic VOC. Since the data were broken down to compound classes, it is possible to assess their importance in SOA formation, though it was not explicitly done in the paper. The most important compound class in terms of potential SOA formation is the group “other aromatics”, which exclude benzene, toluene, and xylenes. The global emission of this class was only 4.6 Tg yr−1, a minute fraction of even the global flux of monoterpenes. Even benzene, toluene and xylenes (BTX aromatics), which are not expected to give rise to significant SOA at all, had an overall emission flux of only 14 Tg yr−1. As a result of these low values, anthropogenic VOC are expected not to contribute significantly to SOA formation in the global troposphere. On the other hand, they might be important in regions under strong anthropogenic influence, especially in the light of recent studies suggesting

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new routes for SOA formation from aromatic precursors (see sub-chapter 3.3.6). 3.3.2 Gas-phase Chemical Reactions Leading to Condensable Species In assessing the atmospheric transformations of biogenic unsaturated compounds, reactions with hydroxyl (OH) and nitrate radicals (NO3) and ozone (O3) must be considered. The initial step of all these reactions proceeds via the addition of oxidant to the double bond of the hydrocarbon. For the OH radical reaction, H-atom abstraction is also possible, but is assumed to be of minor importance. The rate constants for the gas-phase reactions of OH radicals, O3 and NO3 radicals with biogenic hydrocarbons at 298 ± 2 K are shown in Table 3.6 [Hoffmann et al., 1997]. Table 3.6. Rate constants for the gas-phase reactions of OH radicals, O3, and NO3 radicals with the biogenic hydrocarbons studied at 298 ± 2K (after Hoffmann et al. [1997]). k(OH) × 1012 cm3molecule−1s−1

k(O3) × 1018 cm3molecule−1s−1

k(NO3) × 1012 cm3molecule−1s−1

k(NO3) / k(O3) × 10−5

β-pinene

78.9

15.0

2.5

1.67

d-3-carene

88.0

37.1

9.1

2.45

d-limonene

171.0

200.0

12.2

0.61

Terpinene-4-ol

170.0

250.0

14.6

0.58

α-pinene

53.7

86.6

6.2

0.71

Cis/transocimene

252.0

540.0

22.0

0.41

Linalool

159.0

430.0

11.7

0.27

Transcaryophyllene

200.0

11600

19.0

0.02

The reaction rate constants given in Table 3.6 can be combined with assumed ambient concentrations of OH radicals, O3 and NO3 radicals to estimate the tropospheric lifetimes of the biogenic hydrocarbons with respect to gas-phase reactions with these oxidants. These estimated lifetimes are shown in Table 3.7 [Hoffmann et al., 1997].

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Table 3.7. Estimated tropospheric lifetimes of the hydrocarbons studied with respect to reaction with OH radicals, O3, and NO3 radicals (after Hoffmann et al. [1997]). Lifetime due to reaction with OH O3 NO3 daytime daytime nighttime [min] [min] [min]

Rel. contribution of O3-reaction to tropospheric degradation daytime [%]

β-pinene

132

902

1107

13

d-3-carene

118

365

305

24

d-limonene

61

68

228

47

Terpinene-4-ol

61

54

190

53

α-pinene

194

156

451

55

Cis/trans-ocimene

41

25

126

62

Linalool

66

31

237

68

Trans-caryophyllene

52

1

146

98

During the day ozonolysis represents the major degradation pathway for most of the terpenes, as indicated in the last column of Table 3.7. Among the numerous reaction pathways possible between the already limited range of volatile organic compounds and major oxidizing species in the troposphere, the hydroxyl (OH) and nitrate radicals (NO3) and ozone (O3), only those will be considered which yield reaction products that have been identified in the aerosol phase in smog-chamber studies. Even with this restriction, this sub-chapter will be illustrative rather than exhaustive, strongly biased towards the most abundant species for which considerably more smog-chamber studies are available. Generally, the secondary aerosol formation increases with increasing share of the ozone reaction, except for β-pinene and linalool, pointing to the importance of ozonolysis of biogenic hydrocarbons as a major source of nonvolatile products. Reactions with nitrate radicals are possibly more important for β-pinene and ∆3-carene [Hoffmann et al., 1997]. However, the understanding of the formation mechanisms of condensable—in particular low-volatility—species from biogenic hydrocarbons is still in its infancy.

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3.3.2.1 Reactions of Biogenic VOC a) α-pinene The gas-phase chemistry of α-pinene oxidation has been studied for two decades, and remarkable progress has been made in elucidating the products and reaction pathways. A major reaction product of α-pinene oxidation which may partition into the aerosol phase—though it is unable to form new particles due to its high vapor pressure (5.1 Pa at 298 K) [Hallquist et al., 1997]—is pinonaldehyde. It can form in reactions with OH and NO3 radicals and O3, in variable but high yields [Atkinson and Arey, 2003; and references therein]. The possible formation pathways in reactions with OH radicals are illustrated in Figure 3.12.

Figure 3.12. The proposed mechanism leading to the formation of acetone and pinonaldehyde from the oxidation of α-pinene (after Wisthaler et al. [2001]).

Hoffmann et al. [1998] conducted reaction chamber experiments to detect condensable species in the α-pinene/ozone reaction system, with particular emphasis on new particle formation processes. Their experimental setup removed excess ozone with a diffusion denuder prior to collection of the reaction products on adsorption tubes in order to exclude secondary formation of artifacts on the adsorbent surface. For the detection of mostly oxygen-containing polar species they used the on-line method of liquid chromatograpy-mass spectrometry in the atmospheric pressure negative chemical ionization mode (APCI), since these species are known to stabilize

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negative charges due to their high gas-phase acidity. Using this technique they confirmed the presence of pinonic acid, a ketocarboxylic acid, which had been known to form from the ozonolysis of α-pinene as suggested by Hatakeyama et al. [1989]. α-pinene is first converted to high-energy Criegee biradicals by the exothermic addition reaction of O3 into the double bond of α-pinene. The Criegee biradicals can either be collisionally stabilized or convert to several substituted four-member cyclic compounds through ring cleavage of the original six-member ring of α-pinene. The reaction scheme is shown in Figure 3.13.

Figure 3.13. Formation of pinonic acid during the ozonolysis of α-pinene (after Hoffmann et al. [1998]).

An analogous compound may also be formed between the peroxy radicals formed in the reaction between α-pinene and NO3 radical [Hallquist, 1999]. The formation of a cyclic diacid, cis-pinic acid was also evidenced by its mass spectra, which had been previously identified as an ozonolysis product of α-pinene [Hoffmann et al., 1998]. Furthermore, a diacid with a molecular weight of 172 (C8H12O4) was detected, and tentatively assigned to norpinic acid (2,2-dimethyl-cyclobutane-1,3-dicarboxylic acid). The structures of these two species are shown in Figure 3.14. The carboxylic acids are likely formed by the autoxidation of carbonyls in the liquid phase, though the atmospheric significance of these mechanisms is yet to be confirmed [Jang and Kamens, 1999].

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Figure 3.14. Chemical structures of α-pinene-O3 reaction products. I – nor-pinonic acid; II – cis-pinonic acid; III – 2,2-dimethylcyclobutane-1,3-dicarboxylic acid; IV – cis-pinic acid; V – 2,2-dimethylcyclobutane-1-carboxylic acid-3-carboxaldehyde; VI – nor-pinonaldehyde; VII – cis-pinonaldehyde; VIII – 2,2-dimethylcyclobutane-1,3-dicarboxaldehyde; P1 – cis-3carboxyformyl-2,2-dimethylcyclobutaneacetic acid; P2 – cis-3-carboxymethylol-2,2dimethylcyclobutaneacetic acid; P3 – cis-3-carboxylic acid-2,2-dimethylcyclobutaneacetaldehyde (after Jang and Kamens [1999]).

For the first time, Hoffmann et al. [1998] also observed a stable dimer of cis-pinic and norpinic acids, which could be separated by liquid chromatography. However, it was not clear whether such adduct actually formed in the reactions or was merely an artifact in the ion source of the mass spectrometer. In addition, adduct formation was evidenced only in experiments with relatively high α-pinene and ozone concentrations, while it is known that such reactions depend strongly on the mixing ratios of the reacting species.

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When the experimental setup was complemented with an activated charcoal denuder to remove gas-phase compounds, a decrease in abundance was observed for the carboxylic acids but not for the dimer. This strongly support that such adduct does form in the reaction and it has extremely low vapor pressure, pinpointing to its potential role in new particle formation from biogenic VOC oxidation. The formation of such dimers as gas-phase clusters may be the first step in homogeneous nucleation processes, which then proceed by subsequent addition of further molecules to form stable trimers, tetramers, etc., kept together by strong intermolecular forces. However, to date no experimental evidence for the presence of such oligomers have been obtained. The radical initiated oxidation of pinonaldehyde in the presence of excess NO2 yields a PAN-analogue3 compound, 3-acetyl-2,2-dimethylcyclobutane-acetyl peroxynitrate, often referred to as α-pinonylperoxynitrate (αP-PAN), in the reaction sequence shown in Figure 3.15, where X can be either OH or NO3 radical in the atmosphere [Wängberg et al., 1997; Nozière and Barnes, 1998].

Figure 3.15. Formation of α-pinonyl peroxynitrate from pinonaldehyde, where X=OH or NO3 (after Nozière and Barnes [1998]).

The yield of αP-PAN, which has a molecular weight of 245 g mol−1 and very low equilibrium vapor pressure, was found to be 100 % and 81.3 ± 16 % for reactions with NO3 and OH radicals, respectively [Nozière and Barnes, 1998]. Albeit secondary aerosol formation from αP-PAN was implicitly inferred from smog-chamber measurements, it is still unclear whether it can be significant under atmospheric conditions at much lower αP-PAN concentrations. 3 PAN is peroxyacetyl nitrate, an adduct of peroxyacetyl radical and nitrogen dioxide, of the chemical formula of CH3CO.O2NO2, a well-known and abundant component of photochemical smog.

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Secondary aerosol formation from NO3 oxidation of α-pinene appears to be controlled by the concentration of the nitrooxy peroxy radicals, which are in equilibrium with the thermally unstable nitrooxy peroxy nitrates [Barthelmie and Pryor, 1999]: C10H16 + NO3 → C10(ONO2)H16 + O2 → C10(ONO2)H16OO C10(ONO2)H16OO + NO2 ↔ C10(ONO2)H16OONO2

(3.4) (3.5)

An important step is the reaction of nitrooxy peroxy radicals with NO to yield dinitrate which is assumed to partition into the aerosol phase with about the same efficiency as pinonic acid [Barthelmie and Pryor, 1999]: C10(ONO2)H16OO + NO → C10(ONO2)H16ONO2

(3.6)

b) β-pinene β-pinene/OH product studies showed that the only major product was nopinone with a yield of 0.79, in analogy with the reaction of β-pinene with ozone [Hatakeyama et al., 1989, 1991; Hakola et al., 1994]. The postulated formation mechanism of nopinone in the reaction of β-pinene with OH radicals is shown in Figure 3.16 [Wisthaler et al., 2001].

Figure 3.16. The proposed mechanism leading to the formation of acetone and nopinone from the oxidation of β-pinene (after Wisthaler et al. [2001]).

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Since the equilibrium vapor pressure of nopinone is too high, it can only partition into pre-existing organic aerosol phase, in the gas-phase it readily decomposes to acetone. A chemical model fitted to smog chamber data showed that 68, 10 and 22 % of the β-pinene reacted with OH radicals, ozone and NO3 radicals whereas the aerosol yield was distributed as <0.1 %, 12 % and 88 % among the OH-, O3- and NO3-products, respectively [Hoffman et al., 1997]. The observed delay in the onset of the evolution of aerosol concentration pointed out to the importance of secondary oxidation reactions of the first generation oxidation products. c) Other monoterpenes and sesquiterpenes The reaction of ∆3-carene with OH was shown to form a ketoaldehyde (3-caron aldehyde, 2,2-dimethyl-3-(2-oxopropyl)-cyclopropane acetaldehyde) as the main product (α = 0.34) [Hakola et al., 1994]. This substance has a chemical structure almost identical to that of pinonaldehyde (VII in Figure 3.14), thus it is expected to behave similarly in partitioning between the gas and aerosol phase.

Figure 3.17. Major oxidation products from the atmospheric degradation of d-limonene and linalool. IPOH – 3-isopropenyl-6-oxo-heptanal; MVT – 5-methyl-5-vinyltetrahydrofuran-2-ol (after Calogirou et al. [1999a]).

Sabina ketone (5-isopropyl-bicyclo[3.1.0]hexan-2-one), is the only known oxidation product of sabinene with O3 and OH, but its partitioning behavior is not known [Calogirou et al., 1999b]. In spite of the fact that cyclic diolefins, d-limonene, α-terpinene, γterpinene and terpinolene have been shown to generate secondary aerosol at the highest yields after sesquiterpenes in smog-chamber experiments, very little is known about the chemistry of their reactions with OH or O3 [Griffin

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et al., 1999a]. For d-limonene, it was found that about 70 % of the organic aerosol yield resulted from the gas-to-particle conversion of primary products, while the balance can be attributed to secondary products [Hoffmann et al., 1997]. A major primary product in both OH and O3 reactions is limonon aldehyde (3-isopropenyl-6-oxo-heptanal, in Figure 3.17) [Calogirou et al., 1999a]. The acyclic triolefins, myrcene and ocimene were shown to be relatively less important in generating secondary aerosol in smog-chambers, as they are likely to produce primarily volatile decomposition products in their reactions with O3. The same conclusions may apply to oxygenated terpenes, linalool and terpinene-4-ol. The structure of the major oxidation product of linalool, 5-methyl-5-vinyltetrahydrofuran-2-ol (MVT), is shown in Figure 3.17. Sesquiterpenes are by far the most efficient in producing secondary organic aerosol due to their high molecular weights. The major primary products in the reaction of β-caryophyllene with O3 were the ketoaldehydes β-caryophyllon aldehyde (3,3-dimethyl-γ-methylene-2-(3-oxobutyl)cyclobutyl-butanal, a in Figure 3.18) and keto-β-caryophyllon aldehyde (3,3dimethyl-γ-oxo-2-(3-oxobutyl)-cyclobutyl-butanal, b in Figure 3.18) [Calogirou et al., 1999b]. However, secondary reactions were found to be even more important in aerosol formation from sesquiterpenes, producing high molecular weight carboxylic acids [Hoffmann et al., 1997].

Figure 3.18. Major gas-phase ozonolysis products of β-caryophyllene. a – β-caryophyllon aldehyde (3,3-dimethyl-γ-methylene-2-(3-oxobutyl)-cyclobutyl-butanal); b – keto-β-caryophyllon aldehyde (3,3-dimethyl-γ-oxo-2-(3-oxobutyl)-cyclobutyl-butanal) (after Calogirou et al. [1999b]).

Contrary to the commonly held view that aerosol formation from isoprene photooxidation is negligible [Pandis et al., 1991], this possibility has recently been raised by the identification of two diastereoisomers of a polyol, 2-methylthreitol and 2-methylerythritol in forest aerosol [Claeys et al., 2004]. The structures of these species are shown in Figure 3.35. A

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possible gas-phase oxidation mechanism has also been postulated as shown in Figure 3.19.

Figure 3.19. Proposed formation of the 2-methyltetrols from isoprene by reaction with OH˙O2 followed by self- and cross-reactions of peroxyradicals. The intermediate 1,2-diols have been reported in chamber experiments with isoprene under low NOx conditions [Ruppert and Becker, 2000] (after Claeys et al. [2004]).

3.3.2.2 Anthropogenic VOC Generally, aromatic compounds were shown to account for the bulk of secondary aerosol formation from whole gasoline vapor [Odum et al., 1997a]. In terms of secondary aerosol formation potential, aromatic species can be divided into high-yield and low-yield compounds. The high-yield species are those with one or fewer methyl and one or fewer ethyl substituent (e.g. toluene, ethylbenzene, ethyltoluene) as well as n-propylbenzene, whereas the low-yield species are those containing two or more methyl substituents (e.g. xylenes, trimethylbenzenes, tetramethylbenzenes) [Odum et al., 1997a].

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In spite of the fact that many individual aromatic species generate similar or even identical oxidation products, it was speculated that the singlesubstituted compounds might generate a higher ratio of ring-retaining to ringcleavage products than multiple-substituted species. Because of their larger molecular weight, ring-retaining products are expected to have lower vapor pressure resulting in higher aerosol yields [Odum et al., 1997a]. However, a detailed study on the mechanisms of aromatic photooxidation reactions did not support this simple hypothesis [Forstner et al., 1997b]. While nearly equal split was observed between the identified ring-retaining and ringfragmentation products in the reaction between 1,2,4-trimethylbenzene and OH radicals, in the case of ethylbenzene a significant portion of identified species were ring-fragmentation products. It is likely that the products— which are produced in secondary reactions and cannot be identified—play a significant role in secondary aerosol formation from aromatic species. For toluene, the fraction of ring-retaining products was found to be less than 40 %. The predominant species identified in the aerosol phase were 3methyl-2,5-furandione, dihydro-2,5-furandione, 2-methyl-4-nitrophenol, 2,5furandione, 3-methyl-4-nitrophenol, and benzoic acid [Forstner et al., 1997b]. The most abundant species 3-methyl-2,5-furandione, also known as citraconic anhydride, is rather volatile (its estimated vapor pressure is 1 mmHg at 47.1 °C), so it is likely to absorb into preexisting organic aerosol phase. The mechanism leading to its formation is shown in Figure 3.20.

Figure 3.20. Mechanisms of toluene-OH reaction leading to 3-methyl-2,5-furandione and 3methyl-2(5 H)-furanone (after Forstner et al. [1997b]).

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Another major and volatile oxidation product is 2,5-furandione, which can be formed in the reaction between toluene and OH radicals as shown in Figure 3.21.

Figure 3.21. Toluene-OH pathway leading to 2,5-furandione (after Forstner et al. [1997b]).

This compound may transform into dihydro-2,5-furandione (succinic anhydride) in photolytically-induced intermolecular hydrogen abstraction in the particle phase, as shown in Figure 3.22.

Figure 3.22. Photolytically induced mechanism from 2,5-furandione to dihydro-2,5furandione (after Forstner et al. [1997b]).

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The nitrophenols are formed in the reactions of NO2 and methyl hydroxyhexadienyl radicals resulting from OH addition on toluene, as shown in Figure 3.23.

Figure 3.23. Reactions of NO2 with the methyl hydroxycyclohexadienyl radical resulting from toluene-OH reaction leading to observed organic aerosol products (after Forstner et al. [1997b]).

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Among the products of the reaction between m-xylene and OH radicals, the ring-fragmentation products predominated (∼75 %), the most abundant being 3-methyl-2,5-furandione, m-toluic acid, and 2,5-furandione. In secondary aerosol formed from 1,2,4-trimethylbenzene 4-methylphtalic acid, 3-methyl-2,5-furandione, 3,4-dimethylbenzoic acid, and 2,5-furandione were identified. The species identified among the ethylbenzene oxidation products were acetophenone, 3-methyl-2,5-furandione, 2,5-furandione, dihydro-5-methyl-2-furanone, benzaldehyde, 3-ethyl-2,5-furandione, and ethyl nitrophenol. This last compound was assumed to be formed in the reaction between ethylbenzene-OH adduct and NO2, as shown in Figure 3.24.

Figure 3.24. Suggested pathways of ethylbenzene-OH adduct leading to ethyl nitrophenolic compounds (after Forstner et al. [1997b]).

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Another group of volatile organic compounds of anthropogenic origin which is capable of producing secondary aerosol comprises higher molecular weight alkenes. The molecular composition of secondary organic aerosol formed from the photooxidation of higher alkenes, 1-octene and 1-decene was also studied in details [Forstner et al., 1997a]. The identified fraction consisted primarily of heptanal, heptanoic acid and dihydro-5-propyl-2(3H)furanone from 1-octene. The corresponding compounds for 1-decene were nonanal, nonanoic acid and dihydro-5-pentyl-2(3H)furanone. The aldehydes are understood to form via 1,2,3-trioxolane adduct which then decomposes into two excited Criegee biradicals and two aldehydes: heptanal/nonanal and formaldehyde, as shown in Figure 3.25.

Figure 3.25. Mechanism of O3 addition to 1-octene/1-decene leading to heptanal/nonanal (after Forstner et al. [1997a]).

Model calculations proved that the initial decay of the alkenes was due to OH radicals, then reactions with ozone and O(3P) became predominant in the case of 1-octene and 1-decene, respectively. Aldehydes may be further oxidized into the corresponding acids via peroxy radicals (Figure 3.25) or alternatively they may form from the isomerization of either the thermally stable or excited biradical (see Figure 3.26).

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Figure 3.26. Mechanism of aldehyde oxidation to its carboxylic acid (after Forstner et al. [1997a]).

The lactones were postulated to form via initial OH attack on the aldehyde, a subsequent oxygen addition and the internal cyclization of the αcarbonyl alkoxy radical, as shown in Figure 3.27 [Wang et al., 1992].

Figure 3.27. Possible mechanism leading to dihydro-5-propyl-2(3H)-furanone and dihydro-5pentyl-2(3H)-furanone from heptanal and nonanal (after Forstner et al. [1999a]).

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3.3.2.3 Secondary Organic Aerosol Formation from Other Compounds There are a host of other compounds which—at least in part—are regarded secondary in origin. However, these species are usually ignored in modeling of either anthropogenic or biogenic secondary organic aerosol formation, for one or more of the following reasons: 1) The compound class is of mixed origin, and the relative contribution of primary and secondary sources is spatiotemporally highly variable and generally poorly understood. 2) Even if a group of compounds is believed to be secondary in origin, it is unclear to what extent their precursors originate from anthropogenic and biogenic sources. 3) The origin and primary/secondary nature of homologues show a marked variability by carbon number, often with overlap between the different subsets. A universal example for all the three points above could be the class of α,ω-dicarboxylic acids in the range of C2–C27. Their occurrence in atmospheric aerosol will be discussed in details in Chapter 4. Some aspects of their chemistry, however, are highlighted here for illustration. 1) The abundant short-chain homologues, C2–C4 α,ω-dicarboxylic acids, oxalic, malonic, and succinic acids can either be emitted directly in vehicular exhaust [Kawamura and Kaplan, 1987], or formed in the atmosphere by photooxidation of aromatic hydrocarbons [Kawamura and Sakaguchi, 1999]. Similarly, longchain dicarboxylic acids in the range of C10–C24 might have an unspecified primary source or may also be attributed to the atmospheric degradation of ω-hydroxy fatty acids found in vascular plants [Simoneit and Mazurek, 1982]. 2) The C8–C9 α,ω-dicarboxylic acids, suberic and azelaic acids are understood to be photooxidation products of unsaturated fatty acids, such as oleic (C18:1), linoleic (C18:2), and palmitoleic (C16:1) acids. These unsaturated fatty acids, however, could originate from microbial sources [Simoneit et al., 1988], or released in great abundance from anthropogenic sources, such as meat cooking operations [Rogge et al., 1991]. 3) In addition to the homologues discussed above, C5–C7 α,ωdicarboxylic acids, glutaric, adipic, and pimelic acids are considered to be formed from the photooxidation of cyclic olefins [Hatakeyama et al., 1985], whereas the higher homologues above C20 are assumed

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to have direct marine origin, and they are possibly released in their salt forms [Stephanou and Stratigakis, 1993]. Note that this latter range overlaps to some extent with the primary and/or secondary sources described in point 1) above. These points taken together clearly illustrate the marked dependence of primary/secondary and/or natural/anthropogenic origin on the carbon number of the homologues, as well as its uncertainties, which make modeling of secondary aerosol formation involving these compounds highly ambiguous. The class of ω-oxocarboxylic acids could be another example for points 2) and 3) above. In spite of the fact that for the above reasons most of these compounds have not been included in atmospheric models of secondary organic aerosol, some established pathways of their atmospheric formation will be illustrated below. The envisaged reaction scheme producing C4–C9 diacids, then predominantly oxalic and malonic acids, according to Kawamura and Sakaguchi [1999], is shown in Figure 3.28.

Figure 3.28. Reaction mechanisms proposed for the photochemical production of (a) C4–C9 diacids from unsaturated fatty acids and their oxidation intermediates and (b) oxalic (C2) and malonic (C3) acids from succinic (C4) acid in the marine atmosphere (modified from Kawamura and Gagosian [1990a] and Kawamura and Ikushima [1993]. Although hydroxymalonic acid has not been identified, it is postulated to exist in the marine aerosols (after Kawamura and Sakaguchi [1999]).

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The reaction of oleic acid with OH radicals or O(3P) are assumed to yield specifically 9,10-epoxyoctadecanoic acid (A), 1-nonanal (B), nnonanoic acid (J), and 9-oxononanoic acid (C), as shown in Figure 3.29 [Stephanou and Stratigakis, 1993].

Figure 3.29. Reaction mechanism for oleic acid photooxidation (after Stephanou and Stratigakis [1993]).

As shown in Figure 3.29, the unstable 9,10-epoxyoctadecanoic acid is expected to produce 1-nonanal and 9-oxononanoic acid, which are further oxidized to 1,9-nonanedioic acid (I) and n-nonanoic acid (J). In the reaction of oleic acid with ozone, in addition to 1-nonanal and 9-oxononanoic acid, stable Criegee biradicals are formed (E and F in Figure 3.30), which in the presence of water give rise to 1,9-nonanedioic acid (I) and n-nonanoic acid (J), as shown in Figure 3.30 [Stephanou and Stratigakis, 1993]. Again, 1-nonanal and n-nonanoic acid may have other sources than photochemical formation, namely direct emission from plants and biomass burning [Greenberg et al., 1984], which further complicates the issue. The possible precursors and gas-phase formation mechanisms of multifunctional, water-soluble organic compounds were evaluated in a comprehensive retrosynthetic4 study [Pun et al., 1999]. This novel approach was the first attempt to bridge the gap between smog-chamber observations and WSOC speciation measurements (see Chapter 4). Many of these compounds were found to be possible first- or second generation products of 4

The theoretical approach is termed retrosynthetic due to the backward direction of the analysis, from products to precursors.

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common biogenic and/or anthropogenic VOC such as alkanes, alkenes, cyclic alkenes, and aromatics.

Figure 3.30. Reaction mechanism for oleic acid ozonolysis (after Stephanou and Stratigakis [1993]).

For example, a carboxylic group may form in the reactions between alkenes and O3, and aldehydes and OH under low NOx conditions. Thus, dicarboxylic acids can either form directly from unsaturated carboxylic acids, or from (ω)-oxo-carboxylic acids. Unsaturated carboxylic acids may be directly produced from dienes, or through the oxidation of the aldehyde group of an unsaturated aldehyde, which also likely forms from dienes. Retrosynthetic analysis of the other pathway (i.e. via (ω)-oxo-carboxylic acid intermediates) identified alkanes, alkenes, alcohols, hydroxyalkenes, and aromatics as possible primary precursors. Some of the intermediates such as hydroxyaldehydes and unsaturated aldehydes may also be directly emitted. These precursors may derive from a multitude of biogenic and anthropogenic sources. An illustrative example of this mechanism is 3-hydroxypropanoic acid, which was identified in tropospheric aerosol samples [Pun et al., 1999]. The scheme of the retrosynthetic mechanisms is shown in Figure 3.31.

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Figure 3.31. Retrosynthetic analysis of 3-hydroxypropanoic acid (after Pun et al. [2000]).

As shown in the figure, possible precursors of 3-hydroxypropanoic acid include (Z)-3-hexen-1-ol (also called leaf alcohol), which is emitted by vegetation, simpler unsaturated alcohols such as 3-buten-1-ol, as well as 1,3propanediol. To evaluate the relative importance of the identified pathways, a SOA formation index (SFI) was introduced, defined as a ratio of the stoichiometric coefficient (a) of the condensable product to the characteristic time scale (τ) of the reaction for any first-generation product: SFI = a/τ

(3.7)

High SFI values, hence efficient SOA production result from high yields of condensable products and/or smaller time constants (faster reactions). The observed SFI values were quite significant for several pathways, implying efficient SOA formation, especially for high oxidant concentrations. It is important to note that the characteristic time-scale for most of these reactions is in the order of days. Therefore they were not detected in smog-chamber studies, and the aerosol forming potential of low molecular weight species was assumed to be essentially zero. As a result, current SOA models do not attribute aerosol forming potential to most of these precursors. On the other hand, when multifunctional and water-soluble species are

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formed, they can readily partition into aqueous aerosol particles. For example, the lowest diacid, oxalic acid has so low volatility that it is present entirely in the aerosol phase. 3.3.3 Gas-to-particle Partitioning of Reaction Products 3.3.3.1 Aerosol Yield—Fractional Aerosol Coefficient The simplest and most convenient parameter to estimate the fraction of reactive volatile organic species that is converted to aerosol is the so-called fractional aerosol coefficient (FACi) [Grosjean and Seinfeld, 1989]. This dimensionless parameter is defined as the ratio of secondary aerosol formation rate (kg day−1) to the emission rate of parent reactive organic gases (ROG) (kg day−1).

FAC i =

(

aerosol from ROGi kg day −1 ROGi kg day −1

(

)

)

(3.8)

Thus an “emission parameter” can be calculated for secondary organic aerosol as the product of FACi and the emission rate, allowing secondary organic aerosol to be modeled as if they were primary emissions. Grosjean and Seinfeld [1989] compiled available literature data for organic aerosol formation for a large number of volatile species on the basis of the concept of dual “reactivity-volatility” dependence [Grosjean, 1976]. This concept assumes that efficient aerosol formation requires relatively fast gas-phase reactions yielding products of sufficiently low vapor pressure. Secondary aerosol formation can thus either be “reactivity-limited” (low rates of atmospheric reactions) or “volatility-limited” (the products are not formed in sufficient quantities to reach saturation). It is important to note that here saturation is a prerequisite for aerosol formation, though recent studies on gas-particle partitioning showed that absorption into the organic phase can occur well below saturation. Besides for compounds with available experimental parameters, the focus was extended to dozens of other volatile organic species. Reaction rates with O3 were estimated from available ionization potentials, and for reactions with OH radicals, a cut-off value of kOH = 5 × 10−12 cm3 molecule−1 s−1 was adapted as a limit for reactivity to eliminate slowly reacting compounds. Similarly a vapor concentration of 1 ppb was set as a limit for saturation concentration of condensing molecules. Thus a consolidated list of fractional aerosol coefficients was established.

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In general, good agreement was found between experimental observations and “reactivity-volatility” considerations. No aerosol formation can be observed from compounds having less than seven carbon atoms. Significant fractional aerosol coefficients can be assigned to higher alkenes (12 and 15 % for 1-undecene and trimethyldecene, respectively), especially to cycloalkenes (8.5–15 %). Even higher n-paraffins exhibited low values (5 % for n-pentadecane), the coefficients for cycloalkanes were similarly low. Normalized to the number of carbon atoms, aromatic compounds possessed a relatively high fractional aerosol coefficient, though they were not high in absolute terms (e.g. o-xylene 6.3 %). It was the class of the monoterpenes which had by far the largest aerosol production yield, estimated to be in the range of 30–50 %. Contrary to first-glance expectations, FAC had a marked timedependence when interpreting smog-chamber studies, starting from FAC=0 in the induction period then gradually increasing to its maximum value reported in the consolidated list. It should be acknowledged that the estimates provided by Grosjean and Seinfeld [1989] were based on literature data available at that time, and since then tremendous scientific efforts—both theoretical and experimental—have been put into smog-chamber studies including detailed organic speciation and the development, as well as refinement and validation of gas-particle partitioning models. Initially, the parameterization of the results of smog-chamber experiments was based on experimentally determined fractional aerosol yields (Y). The fractional aerosol yield (Y) is defined as the fraction of a reactive organic gas (ROG) reacted that is converted to aerosol:

Y=

∆M 0 ∆ROG

(3.9)

where ∆M0 is the organic aerosol mass concentration (µg m−3) produced for a given amount of ROG reacted, ∆ROG (µg m−3). In principle, fractional aerosol coefficients and aerosol yields are meant to represent the same, except that the former is a theoretical parameter deduced from reactivity and product volatility considerations for any given volatile organic compound. To some extent, it requires a fundamental understanding of the mechanism of the oxidation reactions, estimates of reaction rates and stoichiometric constants, identification of the reaction products and estimates of their volatilities. On the contrary, experimental determination of aerosol yields requires only the measurement of aerosol volume or mass concentration in smog-chamber studies.

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For about 25 years considerable efforts have been put into the experimental determination of aerosol yields in smog-chamber experiments, but the results for any individual compound showed a large scatter between and even within laboratories. As an example, the aerosol yield for α-pinene was found to be vary from zero up to 90 % [Hallquist et al., 1999; Hatakeyama et al., 1989; Hull, 1981], pointing to the weakness of this simplified concept. The major shortcoming of this concept is the explicit assumption that any vapor-phase product starts to condense upon preexisting seed particles or nucleate homogeneously only when its gas-phase concentration exceeds its saturation value. Our current understanding now tells us that even homogeneous nucleation may occur at mixing ratios well below saturation, if two or more species are involved (bimolecular homogeneous nucleation). The breakthrough in the understanding secondary aerosol formation was the theory of absorption partitioning by Pankow [1994], which was then developed into a comprehensive model to explain secondary organic aerosol yields [Odum et al., 1996]. 3.3.3.2 Absorption Partitioning Theory The concept of absorption partitioning suggests that once semi-volatile organic compounds have begun to condense and an organic layer has formed on the surface of the particles, more volatile products with mixing ratios well below saturation values will partition in part into this condensed organic phase. Gas-particle partitioning can be parameterized with the partitioning constant Kp,i as

K p ,i =

Fi TSP Ai

(3.10)

where Kp,i is the temperature-dependent partitioning constant (m3 µg−1), TSP is the mass concentration of total suspended particulates (µg m−3), Fi (ng m−3) and Ai (ng m−3) are the concentrations of species i in the aerosol and gas phase, respectively. The higher of the value of Kp,i, the more of compounds i will be absorbed in the aerosol phase per unit mass concentration of particulate matter. In spite of its name, Kp,i is actually not constant but depends on a number of parameters as will be shown below. It should be noted that symbols and units are used differently in various papers related to secondary aerosol formation. Here we follow the derivation of Pankow [1994], but will use a simplified notation and SI units that have been adapted in subsequent studies.

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If there are pre-existing organic particles, equilibrium gas-particle absorptive partitioning takes place governed by the following equilibrium

pi = X i , omξ i p L0 , i

(3.11)

where pi (Pa) is the gas-phase partial pressure of species i, Xi,om is the mole fraction of species i in the organic aerosol phase, ξi is the activity coefficient of species i in the absorbing organic mixture, and pL0,i (Pa) is the equilibrium vapor pressure of the pure liquid i (subcooled if necessary). Calculating pL0,i from the vapor pressure of the solid (pS0,i) at a given temperature can be performed with the following equation:

ln

p L0 , i p

0 s, i

=

∆S melt , i  Tm − T  R  T 

(3.12)

where ∆Smelt,i (J mol−1 K−1) is the entropy of melting (fusion), R is the universal gas constant (= 8.314 J mol−1 K−1), Tm (K) is the melting point of compound i, T (K) is the temperature of interest. The value of ∆Smelt,i is in the order of 80 J mol−1 K−1 for associating, polar molecules that most likely make up secondary aerosol. When T < Tm, pL0,i will always be larger than pS0,i, that is a subcooled liquid is at a higher energy level than the solid. The molar concentration of species i in the gas phase, ni,g,/V (mol m−3) can be expressed as

ni , g V

(mol m ) = −3

pi RT

(3.13)

which can be transformed into mass concentration Ai by the relation

(

)

Ai ng m −3 =

ni , g V

MWi 10 9

(3.14)

where MWi is the molecular weight of compound i (g mol−1). Combining equations 3.13, 3.15, and 3.16 we obtain that

(

Ai ng m

−3

)=

X i , om ξ i p L0 , i RT

MWi 10 9

(3.15)

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In a unit volume of air (1 m3), Xi, om will be

X i , om =

ni , om ntot , om

(3.16)

where ni,om is the number of moles of species i absorbed into the organic aerosol phase, and ntot,om is the total number of moles of all organic compounds in the aerosol. ntot,om equals to the mass of organic phase in unit volume of air divided by its mean molecular weight:

ntot , om =

f om TSP MWom 10 6

(3.17)

where fom is the mass fraction of the absorbing organic phase in the total suspended particulates (TSP), MWom is its mean molecular weight (g mol−1). In 1 m3 of air ni, om will be

ni , om =

Fi , om

MWi 10 9 Combining equations 3.17–3.20 we have

Ai =

Fi , om MWomξ i p L0 , i 10 6 f omTSP ⋅ RT

(3.18)

(3.19)

When gas/particle partitioning is dominated by absorption, the partitioning constant Kp,i will be

K p ,i =

Fi , om TSP Ai

=

f om RT MWomξ i p L0 , i 10 6

(3.20)

If pL0,i is expressed in units of torr and R = 8.2 × 10−5 m3 atm mol−1 K−1, then equation 3.20 will take the form of

K p ,i =

Fi , om TSP Ai

=

f om 760 RT MWomξ i p L0 , i 10 6

(3.21)

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which is the original formula derived by Pankow [1994]. This quantity incorporates vapor pressure, activity coefficient and molecular weight to provide a single equilibrium parameter for each compound, which is analogous to Henry’s law coefficient. Analogously, if only the mass of the organic phase is considered, a partitioning coefficient for species i can be defined

K om, i =

Fi , om Ai M 0

= K p , i f om

(3.22)

where like Kp,i, Kom,i has units of m3 µg−1, and M0 is the absorbing organic mass concentration (µg m−3). Note that M0 is generally not equal to total organic aerosol mass concentration in ambient aerosol, since not all organic aerosol components participate in absorptive partitioning. In contrast, in smog-chamber studies, the entire mass of organic aerosol behaves as an absorbing medium. It follows from Equation (3.22) that with constant Kom,i, the mass of a given compound in the aerosol phase actually increases with increasing absorbing organic mass concentrations, M0. In other terms, the fractional aerosol yield Y will be dependent on M0, which accounts for the high variability of Y determined in different smog-chamber experiments. With time-constants in the order of a few hours, formation of semivolatile products and particle deposition occurs relatively slowly as compared to gas-to-particle transport, which typically involves a time-scale in the order of a minute. This ensures that the gas-particle system will remain in almost perfect equilibrium at all times. Only when very small initial aerosol surface is available or when semi-volatile products are formed from very reactive parent compounds, can the system depart from the state of equilibrium. Such non-equilibrium conditions, however, are unlikely to persist since the growth of small particles and/or nucleation increase available surface area and enhance the rate of gas-to-particle transport. a) Effect of temperature In addition of the temperature term represented explicitly in Equation 3.20, the vapor pressure, pL0,i itself depends strongly on temperature. This dependence can be expressed by the Clausius-Clapeyron equation

 − Hi  p i0 = Bi exp   RT 

(3.23)

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where Bi is the preexponential constant of species i (Pa), and Hi is its enthalpy of vaporization (J mol−1). A 10 °C decrease in temperature results in an increase in the yields for high-yield aromatic compounds by 20–60 % at M0 = 50 µg m−3, and by 50–150 % at M0 = 5 µg m−3 [Sheehan and Bowman, 2001]. The secondary aerosol formed from α-pinene appeared to exhibit an even greater sensitivity to temperature. This had a substantial effect on the modeled diurnal variations of the mass concentration of secondary organic aerosol. For a ±10 °C temperature profile the aerosol mass concentrations were found to decrease by 16 % at high daytime temperatures whereas they increase by 22 % at low nighttime temperatures as compared to the constant temperature case. It was concluded that the available constant temperature partitioning coefficients determined at high temperatures are inadequate in representing ambient partitioning behavior. The systematic errors are greatest when partitioning coefficients are measured outside or near the limits of the ambient temperature range [Sheehan and Bowman, 2001]. b) Effects of relative humidity In calculations for low-humidity conditions MWom refers only to organic species, whereas for typical ambient conditions the uptake of water should also be considered [Kamens et al., 1999]. Partitioning of semi-volatile organic species into an inorganic-water zone is likely of little importance given the very high activity coefficients and low solubilities of organic species in such a liquid phase, with the exception of highly polar water soluble species [Jang and Kamens, 1998]. However, humidity may influence secondary organic aerosol partitioning by altering the chemical and physical nature of the absorbing phase, that is shifting its composition away from a pure organic mixture to an aqueous organic-inorganic mixture, with corresponding changes in the mean molecular weight of the absorbing phase, MWom and the activity coefficients, ξi of secondary products [Griffin et al., 1999b]. MWom decreases with increasing water fraction, thereby favoring partitioning into the aerosol phase. For hydrophobic compounds, ξi increases with increasing water fraction, which translates into decreasing solubility, but for polar secondary organic compounds no change in ξi with increasing RH was observed [Jang and Kamens, 1998]. Thus, the overall effect for most secondary species seems to be governed by the decrease in MWom. Since secondary organic aerosol particles exhibit hygroscopic growth with

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increasing RH (for in-depth discussion see sub-chapter 5.2), the estimated effect will be an increase in Kp,i with increasing RH. c) Comparison with partitioning data derived from ambient gas and aerosol phase measurements In order to test the applicability of the absorption partitioning model for understanding ambient secondary aerosol formation, the Kp,i values obtained by theoretical calculations using measured vapor pressures, or by curve fitting to smog-chamber data, were compared to estimated values derived from ambient measurements of gas and aerosol phase components, as presented in Table 3.8 [Andersson-Sköld and Simpson, 2001]. Table 3.8. Estimated Kp values derived from theory, model simulations, measured vapor pressure, and ambient measurements of gas and aerosol phase components (after AnderssonSköld and Simpson [2001]). Species Pinic acid

Pinonic acid

Pinonaldehyde

Reference Kamens et al. [1999]a Barthelmie and Pryor [1999]b Kamens et al. [1999]a

5°C

Kp-value 25°C

40°C

0.43

0.036

0.0069 0.11

7.9×10−3

7.9×10−4

Barthelmie and Pryor [1999]b Kamens et al. [1999]a

5.6×10−5

7.1×10−6

Hallquist et al. [1997]c

1.9×10−5

2.4×10−6

1.7×10−4 3.6×10−3 1.8×10−6

Source and Temperature Kavouras et al. [1999b], T=6°–25°C Yu et al. [1999b], T=18°–32°C Kavouras et al. [1999b], T=6°–25°C Yu et al. [1999b], T=18°–32°C Kavouras et al. [1999b], T=6°–25°C Yu et al. [1999b], T=18°–32°C

Kpvalue ~0.01– infinityd 0.009– 0.04e 0.04– 0.30e 0.009– 0.04f 0.01– 0.05e 0.004– 0.007 f

Barthelmie 0.001 and Pryor [1999]b a From Kamens et al. [1999] species: diacid for pinic acid, pinacid for pinonic acid, and pinald for pinonaldehyde.;bBarthelmie and Pryor [1999] used 40°C when simulating the Hoffmann et al. [1997] experiments, and derived Kp values from curve fitting.; cHallquist et al. [1997] measured vapor pressures.; dEstimated from figures. ;eValues are quoted from Kavouras et al. [1999b], but inspection of figures suggests periods where Kp is essentially infinity (no gas phase) for pinonaldehyde and pinic acid. ;fRange of values estimated assuming M0 = 1–5 µg m−3.

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One characteristic feature of this comparison is the strong temperature dependence of the Kp values, resulting in a 50–100-fold increase from a typical smog-chamber temperature of 40 °C down to 5 °C. Apart from the fact that the estimated Kp values at 40 °C were different by at least an order of magnitude for the various approaches, they were significantly lower than those derived from ambient measurements, pinpointing to the uncertainties and limitations of the application of the thermodynamic equilibrium theory. 3.3.3.3 Aerosol Yield Based on Absorption Partitioning The concentration of an individual reaction product formed by the photooxidation of a reactive organic gas (ROG) is proportional to the amount of ROG reacted (∆ROG)

1000α i ∆ROG = C i

(3.24)

where αi is the proportionality constant relating the mass concentration of the ROG reacted to the total mass concentration of product i (Ci) that is formed and distributed between the gas and aerosol phase. It corresponds to the stoichiometric factor of the reaction multiplied by the ratio of the molecular weight of product i to that of the precursor ROG. The factor of 1000 is needed to convert mass concentration units of µg m−3 for ∆ROG to ng m−3 for Ci. By definition, the fractional aerosol yield of a given species (Yi) is the mass of aerosol formed (Fi) divided by the mass of the species that has reacted (∆ROG) per unit volume. Since a semi-volatile product of the reaction formed in total concentration of Ci is distributed between the gas and aerosol phase (that is Ai+Fi=Ci) according to the absorption partitioning equilibrium (Eq. 3.22), the following equation can be derived [Odum et al., 1996]:

 α i K om , i Yi = M 0   1 + K om , i M 0 

   

(3.25)

where M0 is the total organic mass concentration in µg m−3 and Kom,i the partitioning coefficient as defined in Eq. 3.22. Since SOA formation usually involves several semi-volatile oxidation products, the total SOA yield can be expressed as the sum of all species produced

Secondary Organic Aerosol

 α i K om , i Y = ∑ Yi = M 0 ∑ i i  1 + K om , i M 0 

119

   

(3.26)

It follows that the aerosol yield increases with increasing mass concentration of organic particles. In smog chamber experiments without pre-existing organic particles (M0 = 0), initially when little parent compound has reacted, the yield is negligible, but after the onset of aerosol formation partition equilibrium shifts to the aerosol phase and the yield increases. At low organic aerosol mass concentration and for products that have relatively small partitioning coefficients, the yield will be a linear function of total aerosol mass concentration, M0. On the contrary, for non-volatile products and/or for large organic mass concentrations, the yield will be equal to αi, independently of M0. Apart from this extreme case, however, it is highly unlikely that any ROG will have a unique fractional aerosol yield. Instead, it will exhibit a range of yields over a range of organic aerosol mass concentration. For lack of complete identification and quantification of individual oxidation products, the yield is a convenient overall measure of the aerosol forming potential of the secondary products formed by the atmospheric oxidation of the parent organic molecule. Since the concept of yield is widely used in chemistry in connection with chemical reactions, it is important here to distinguish between stoichiometric coefficients, xi and aerosol yield, Yi, of atmospheric reactions. The stoichiometric coefficients solely depend on the chemical mechanism of gas-phase reactions and are assumed to be constants, representing the total amount of semi-volatile products formed in both the gas and aerosol phase, per unit amount of parent compound reacted. The aerosol yield, on the other hand, refers only to the fraction of semi-volatile products that has partitioned into the aerosol phase, which is not constant but a strong function of the existing organic aerosol mass concentration. It was found that if the fractional aerosol yield (Y) is determined experimentally for any given ROG at various M0 in smog-chamber experiments, and the values are plotted as a function of M0, the function expressed in Eq. 3.26 can be fitted to the data with two assumed products having parameters α1, α2, Kom,1 and Kom,2, which are determined by minimizing the residuals [Odum et al., 1996]. Since typically more than two semi-volatile products form in the reactions, these two products usually do not represent single chemical species, but rather a class of compounds with weighted average αi and Kom ,i values. Mathematically, two products are the minimum needed to fit the observed behavior in smog-chamber experiments—a one-product model is clearly insufficient. On the other hand,

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inclusion of a third or additional product does not improve the degree of fit noticeably, so they may be regarded superfluous. However, inclusion of a more detailed chemical mechanism for αpinene oxidation may give rise to parameters for a set of six product species based on experimental smog-chamber data [Barthelmie and Pryor, 1999]. An important difference is that in this case the actual properties of identified reaction products were considered, and not the weighted average properties of virtual products. Albeit this more physical approach would be desirable for all biogenic VOCs, for the lack of detailed knowledge of the reaction mechanism and secondary reactions only the simplified concept is generally feasible. Alternatively, Kom,i values can be estimated directly using a thermodynamic approach based on the observed proportionality between activation energy and logpL0,i values [Kamens et al., 1999]. Smog-chamber experiments carried out by Odum et al. [1996] confirmed that fractional aerosol yields are clearly additive, as predicted by Eq. 3.26. That is, the four parameters obtained from measurements with a single compound can be used in a mixture of compounds and the overall aerosol yield can be reproduced by such calculations. Such parameterization, however, is only meaningful when it represents equilibrium conditions for the first-generation products of the parent compound [Bowman et al., 1997]. When gas-aerosol transport is slow or when the condensable species are secondary products, this approach may fail. For this reason, coefficients are usually estimated near the end of smog chamber experiments where all parent compounds has reacted, conversion to secondary products is complete and equilibrium between the gas and aerosol phases has been attained. Such parameterization is extremely useful as a simplified approach for a complex array of semi-volatile species for which speciation and experimental determination of key parameters are not feasible. Provided that there is some pre-existing organic aerosol (M0 > 0), which is typical for most part of the troposphere, semi-volatile organic reaction products begin to partition into the aerosol phase upon formation. For the case M0 = 0, it follows from the absorption equilibrium partitioning theory that a threshold exists for aerosol formation [Bowman et al., 1997]. It can be derived by rearranging Equation 3.10:

Ai K i , p = Fi ∑ k Fk and summing over all semi-volatile species i

(3.27)

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121

Fi =1 ∑ k Fk

(3.28)

∑ Fi K i , p = ∑ i

i

This equation defines a threshold for gas-phase mass concentration to allow aerosol formation. As long as ∑AiKp,i < 1, this threshold is not reached. It means that until sufficient parent compounds have reacted to produce the required gas-phase concentrations, no aerosol formation will be possible. Once this threshold has been surpassed, the gas and aerosol phases will maintain equilibrium and keep ∑AiKp,i unity. Concerning the mechanism of partitioning in this case, it is likely that partitioning occurs initially via adsorption until a sufficient organic layer develops to allow absorption to take over [Pankow, 1994]. 3.3.3.4 Aerosol Yield Functions for Various ROG in Smog-chamber Experiments a) Biogenic VOC Daylight smog chamber experiments were conducted with biogenic VOC in the presence of NOx, propene, and ammonium-sulfate seed aerosol [Hoffmann et al., 1997]. The VOC were present in mixing ratios between 20 and 150 ppb, whereas the mixing ratio of NOx varied in the range of 113–240 ppb. The secondary organic aerosol yield for α-pinene is shown in Figure 3.32 as a function of organic aerosol mass M0. Similar smog chamber experiments with biogenic VOC and ozone (150–330 ppb) were also conducted in the dark. Since OH radical scavenger such as cyclohexane was not used to prevent the reaction of alkenes with OH radicals generated in the O3 reaction mechanism, the results reflected the effects of simultaneous O3 and OH reaction, though ozonolysis was thought to be the predominant reaction pathway (typically 80 % of VOC reacted with ozone). Analogously to the daylight experiments the fractional aerosol yield (Y) can be plotted against the organic aerosol mass concentration, as shown in Figure 3.33.

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Figure 3.32. Secondary organic aerosol yield for α-pinene as a function of organic aerosol mass M0 in daylight smog chamber experiments (36 °C< T < 51 °C). Values used to generate the two product model line are 0.03 and 0.326 for α1 and α2, and 0.171 and 0.004 µg m−3 for Kom, 1 and Kom, 2, respectively (after Hoffmann et al. [1997]).

Figure 3.33. Secondary organic aerosol yield for α-pinene as a function of organic aerosol mass M0 in dark smog chamber experiments (after Hoffmann et al. [1997]).

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Careful comparison of the results of the daylight and dark experiments revealed that ozonolysis typically produced higher aerosol yield for αpinene, but with less pronounced dependence on the organic aerosol mass concentration. This type of partitioning behavior can be expected when oxidation products of very low volatility are produced, which partition almost entirely into the aerosol phase. Curve fitting revealed that the hypothetical products possessed absorption equilibrium constants of 0.2 and 0.005 m3 µg−1, and stoichiometric factors of 0.12 and 0.19, respectively. The former non-volatile product might be pinonic acid. When dark ozonolysis experiments were performed at substantially lower temperature (16.1 °C instead of 47.7 °C), a fourfold increase in the aerosol yield was observed, in qualitative agreement with the gas-particle partitioning model. This strong temperature dependence may primarily be caused by some more volatile species, such as pinonic aldehyde, whose equilibrium vapor pressure decreases markedly with temperature. It should be noted that the reaction of α-pinene with ozone was shown to lead to significant organic aerosol formation even at low mixing ratio for the monoterpene (20 ppb) [Kamens, 1981; 1982]. Smog chamber experiments with β-pinene showed that aerosol formation took place already above mixing ratio as low as 20 ppb, and the fractional aerosol yield varied between 0.1 and 8 % [Pandis et al., 1991]. The yield was found to exhibit a marked dependence on the initial β-pinene to NOx concentration ratio, with a maximum in the range between 10 and 15. The presence of a =CH2 group in a cyclic monoterpene structure profoundly affected aerosol yield, as in the case of limonene, β-pinene and sabinene. Sabinene, which has an exocyclic methylene group and a secondary threecarbon ring, is likely to undergo cleavage of both rings upon oxidation, thereby producing open-chain compounds and lower aerosol yield than βpinene [Yu et al., 1999b]. The position of a double bond internal to the ring structure may prevent carbon atoms from being lost from the molecule, as reflected in the higher yields of limonene, α-pinene, ∆3-carene and α- and γ-terpinenes than those of ocimene, myrcene, linalool and terpinolene [Hoffmann et al., 1997]. For the bicyclic olefins, the number of carbon atoms forming the secondary ring also seems to have an effect on the secondary aerosol yield. ∆3-carene has higher yield than α-pinene, which can be attributed to the somewhat higher boiling point of the former (168–169 °C as compared to 155–156 °C) due to the fact that these compounds were shown to retain their secondary ring structures upon oxidation [Yu et al., 1999b]. Terpinene-4-ol was found to have the lowest yield, possibly due to the presence of an –OH group instead of a double bond [Hoffmann et al., 1997].

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As a first approximation for modeling purposes, the aerosol yield from oxidation reactions of α-pinene is about half from β-pinene, which, in turn, is about 75 % of that from limonene [Kanakidou et al., 2000]. In smog chamber experiments, sesquiterpenes were found to exhibit the highest yield of secondary organic aerosol, as a consequence of the high carbon number of the parent compounds [Griffin et al., 1999a]. Aerosol yields for the ozonolysis of biogenic volatile compounds were found to be less dependent on organic mass concentration as compared to sunlight-irradiated photooxidation [Griffin et al., 1999a]. This clearly indicates that ozonolysis produces less volatile secondary products. The exceptions are β-pinene and sabinene, which show significantly lower aerosol yields under these conditions. Reactions with NO3 radicals induce significant secondary aerosol formation for β-pinene, ∆3-carene and sabinene, but not for α-pinene for which the formation of the highly volatile pinonaldehyde is expected [Wängberg et al., 1997]. This finding raises the possibility of significant secondary aerosol formation in forested areas during nighttime when NO3 may accumulate. In smog chamber experiments very low fractional aerosol yield (<0.8 %) was observed for isoprene, the most abundant biogenic VOC over the mixing ratio range of 0.07–6 ppm [Pandis et al., 1991]. The onset of condensation was found at an isoprene mixing ratio of about 250 ppb, and a steep increase in the aerosol volume yield normalized to the concentration of isoprene above the mixing ratio of 1.2 ppm. It should be noted that the maximum mixing ratio of isoprene ever observed outside a forest canopy was only 30 ppb [Ferman, 1981]. The best fit was obtained assuming three condensable products with saturation vapor concentrations of 1.2, 1.3 and 6.2 ppb at 31 °C. Since all of these products were highly volatile and their saturation vapor pressures were not expected to be even approached under ambient conditions, it was concluded that the aerosol yield for isoprene is practically zero. This view has become universally adopted to the extent that isoprene photooxidation has not been considered in atmospheric models as a potential source for secondary organic aerosol formation. It should be pointed out that nowadays this view is challenged by some studies suggesting new gas-phase photooxidation mechanisms, heterogeneous or multiphase pathways5 .

5

For details see sub-chapter 3.3.6.

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125

b) Anthropogenic VOC The virtual semi-volatile photooxidation products of m-xylene are expected to be second-generation products, S1 and S2, whose parameters can be derived as best fit to experimental smog-chamber data [Bowman et al., 1997]. The products can be generated with stoichiometric coefficients, α1 and α2 of 0.03 and 0.14, respectively. The absorptive partitioning coefficients, Kom,1 and Kom,2 were determined to be 0.06 m3 µg−1 and 0.002 m3 µg−1, respectively. In a model simulation initially the gas-phase concentrations, A1 and A2 increased, then aerosol formation started, first dominated by the lower volatility product, S1, and later by the more volatile but also more abundant S2. This finding implies that for the low volatility S1 partitioning into the aerosol phase occurs quite readily, but the more volatile S2 remains in the gas phase as long as an aerosol phase, composed of initially S1, builds up to allow significant partitioning. This is a general scheme which applies to secondary organic aerosol formation from most VOC. A series of sunlight-irradiated smog-chamber experiments were conducted with whole gasoline vapor, in which organic aerosol yields were determined as a function of ∆M0 for 17 aromatic species representing the predominant mass fraction of aromatic compounds in the gasoline [Odum et al., 1997b]. The results are shown in Figure 3.34 indicating that there is little variation in the yields for different isomers. When a mixture of five compounds was introduced into the smogchamber, the resulting yield was found to be remarkably close to that calculated from the contribution of the individual species by Eq. 3.28. This finding quantitatively supports the à priori hypothesis that the total amount of secondary organic aerosol is controlled by the aromatic components of the fuel. This is in accordance with the knowledge that secondary organic aerosol forms only from hydrocarbons containing at least seven carbon atoms [Grosjean and Seinfeld, 1989].

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Figure 3.34. Secondary organic aerosol yields as a function of organic aerosol mass concentration (∆Mo) for 17 individual aromatic species. Each data point represents an individual experiment. Curves are fit to the data using a two-product model in conjunction with Eq. 3.28 by minimizing the weighted squared residuals. Curve 1 is fit with the values 0.038, 0.042, 0.167, and 0.0014 for α1, Kom,1, α2, and Kom,2, respectively. The corresponding values are 0.071, 0.053, 0.138, and 0.0019 for curve 2; 0.083, 0.093, 0.22, and 0.0010 for curve 3; and 0.05, 0.054, 0.136, and 0.002 for curve 4 (after Odum et al. [1997b]).

3.3.3.5 New Particle Formation A fundamental question is whether the photooxidation products of any given VOC could generate new particles by homogeneous nucleation or they simply condense on pre-existing aerosol particles. The answer to this question is particularly important in understanding the formation of natural background aerosol which is a prerequisite of assessing current anthropogenic contribution to the tropospheric aerosol burden. One key parameter in determining nucleation in a VOC-oxidant system is the saturation vapor pressure of the photooxidation products formed. If the saturation vapor pressure is increased, the particle number concentration decreases much more rapidly than mass concentration, because condensation occurs more readily than nucleation. As it approaches 18 ppbv, nucleation is practically inhibited and very few particles are generated [Andrews et al., 1997]. Nucleation rate is highly sensitive to surface tension, as it depends exponentially on the σ3 in the classical theory of homogeneous nucleation. The critical nucleation rate, at which nucleation is considered to be

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significant, is generally accepted to be 1 particle cm−3 s−1 [Pruppacher and Klett, 1978]. According to the simulations of Andrews et al. [1997], this limit requires an organic saturation ratio of higher than 16, which practically excludes the possibility of homogeneous nucleation of any single organic species. a) Smog-chamber observations Environmental chamber experiments were performed to study new particle formation in low-level cyclopentene-, cyclohexene- and α-pineneozone reactions. When 50 ppb of cyclopentene and 100 ppb of ozone were injected into the chamber with very low (<10 cm−3, r >1.5 nm) initial particle number concentration, ultrafine particle formation was observed already after 3 minutes [Gao et al., 2001]. The particle number concentration then increased steadily up to a maximum of 2600 cm−3. Similar results were obtained with other alkenes suggesting that these compounds, even at mixing ratios of only tens of ppbv—which are actually much closer to ambient concentrations than the high concentrations used in previous smog chamber experiments—can readily form new particles in their reactions with ozone. The main reaction product in the case of cyclopentene oxidation was glutaric acid, followed by succinic acid. Dicarboxylic acids constituted about 15 % of the total aerosol mass formed, which was consistent with earlier findings [Hatakeyama et al., 1987]. However, the molar yields of total diacids were significantly lower in these low-level experiments (∼0.2 % for cyclopentene, ∼0.3 % for cyclohexene products, as against ∼0.6 % and ∼6 %, respectively). In contrast, in the α-pinene-ozone system high carbon yields (∼4 %) were observed. When 10 ppbv of SO2 was also introduced into the chamber, the observed nucleation rates became much higher. The concentration of SO2 was so low that it could not affect ozone, yet it had a dramatic effect on new particle formation. The classical binary nucleation theory was applied for the glutaric acid-water system, since glutaric acid has the lowest volatility among the identified reaction products [Gao et al., 2001]. The calculations revealed that the glutaric acid mixing ratio should reach as high as 800 ppb to produce the observed nucleation rate. This clearly indicated that no homogeneous nucleation can occur with the glutaric acid-water only. Other possible aerosol components, such as dialdehydes and ω-oxocarboxylic acids have even higher saturation vapor pressures and are less likely to participate in homogeneous nucleation.

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The most viable explanation—also supported by model calculations— for the observations was that first H2SO4-H2O nuclei form (SO2 could be present in trace amounts even in the ostensibly pure organic experiments), then the gas-phase organic products produced in the ozone reactions partition into these nuclei, initiating water nucleation and rapid increase of the surface area of the nuclei. It was presumed that such a scheme can be plausible in the marine boundary layer, where SO2, NH3, O3 and background hydrocarbons are present together with a relatively low number concentration of preexisting particles. The dicarboxylic acids can either be secondary species and their precursors, the unsaturated hydrocarbons derive from the ocean surface [Bonsang et al., 1991]. Homogeneous nucleation of secondary organic aerosol was shown to be controlled by stabilized Criegee intermediates formed by the ozonolysis of biogenic terpenes [Bonn and Mortgat, 2002]. The Criegee intermediate can further react with carbonyls to form secondary ozonides, or with water vapor, the latter being the dominant reaction pathway. Nucleation is more effective in exocyclic reactions (such as the external double bond of β-pinene) than in endocyclic ones (α-pinene), and its rate was found to anticorrelate with water vapor concentration [Boy and Kulmala, 2002]. The nucleation thresholds for the most abundant monoterpenes, i.e. αpinene, β-pinene, sabinene, limonene and ∆3-carene even for dry conditions are around 2 ppbv. It was calculated that monoterpene-ozone reactions had to occur for 6 days (α-pinene) to 170 days (β-pinene) to initiate homogeneous nucleation, which is highly unlikely in the atmosphere [Bonn and Moortgat, 2003]. Consequently monoterpene reaction products cannot initiate nucleation in the atmosphere. However, sesquiterpenes react with ozone at a rate about a 100 times higher than monoterpenes, and therefore their reactions are much less affected adversely by water vapor. Laboratory experiments with 25 pptv βcaryophyllene and 50 ppbv ozone showed very intense nucleation events with particle numbers exceeding 160,000 cm−3 even under dry conditions [Bonn et al., 2002]. This translates into a nucleation threshold of only 2.5 pptv, requiring only 3 min to be exceeded under conditions typical of a boreal forest site [Bonn and Moortgat, 2003]. This is a likely occurrence therefore ozonolysis of sesquiterpenes could be a plausible source of new particles whose formation has been evidenced by atmospheric observations [Mäkela et al., 1997]. Furthermore, recent evidences suggest that the contribution of sesquiterpenes to the mass concentration of secondary organic aerosol could be far larger than previously estimated [Bonn and Moortgat, 2003].

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3.3.4 Atmospheric Observations of Secondary Organic Aerosol The gas-particle partitioning of monoterpenes photo-oxidation products were studied in a coniferous forest in central Greece [Kavouras et al., 1999a]. A special aerosol collection system was used to remove oxidants during sampling and to control and evaluate losses of semi-volatile compounds. A clear diurnal pattern of ambient monoterpene concentrations was found with maxima at night and minima during the day. At the same time a maximum in Aitken nuclei concentration was observed in the morning, coinciding with a drop in the mixing ratio of terpenes, and a secondary one late in the afternoon. Among the photooxidation products two isomers of pinonic acids (cisand trans-) and norpinonic acids, pinic acid, pinonaldehyde and nopinone were identified both in the gas and aerosol phase. The presence of the latter two compounds in the aerosol phase can be explained only by the condensation onto pre-existing organic particles, since their vapor pressure are too high (e.g. 5.1 Pa at 298 K for pinonaldehyde) to even approach supersaturation [Hallquist et al., 1997]. Pinonic and pinic acids were found to reach their maximum gas phase concentrations during the period of intensive photochemistry, contrary to norpinonic acid which did not follow the same trend. Therefore norpinonic acid could not be the photooxidation product of α-pinene itself, but rather a product of further photooxidation with a time constant of several hours. The results also implied a competitive mechanism for the formation of pinic and pinonic acids from α-pinene. As for the particulate concentrations of these compounds, they were found in the range of 1.0–25.7 ng m−3, 0.4–4.4 ng m−3, and 0.2–5.4 ng m−3 for pinonic, pinic, and norpinonic acid, respectively. The fitting of the two-product partitioning model developed by Odum et al. [1996] to the experimental data yielded Kom values in the range of 0.22–0.30 m3 µg−1 (below 20 °C) and 0.02–0.05 m3 µg−1 for pinonic acid and pinonaldehyde, respectively. These compounds were present overwhelmingly in particles with diameter <0.5 µm, representing up to 26 % of their estimated total mass. The size-resolved chemical composition of aerosol was studied in a Eucalyptus globulus forest in Portugal [Kavouras et al., 1998]. In the size fraction below 0.5 µm cis-pinonic acid was found to be the most abundant, followed by trans-pinonic acid, pinonaldehyde and nopinone. Based on the observed diurnal changes in trans-/cis-pinonic acid concentration ratios, interconversion of the isomers by ultraviolet radiation was hypothesized, though the drop of nighttime concentrations of both isomers implied other chemical sinks as well. The diurnal variation of particulate cis- and transpinonic acid concentration followed closely that of the Aitken nuclei. On the

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contrary, the variations in the concentrations of pinonaldehyde and nopinone were inverse, which can be explained by their photolysis and/or reactions with OH radicals. Given the very high vapor pressure of these compounds, it is likely that their gas-particle partitioning is also governed by temperature variations between day and night. The possible importance of isoprene photooxidation in SOA formation has been raised very recently by the identification of two diastereoisomers of a polyol, 2-methylthreitol and 2-methylerythritol in forest aerosol, as shown in Figure 3.35 [Claeys et al., 2004].

Figure 3.35. GC-MS total ion current chromatogram obtained for a trimethylsilylated extract of the fine size fraction of a Hi-Vol sample collected during the LBA-CLAIRE 2001 campaign (25 to 27 July, day sampling only). Mal, malic acid; Lev, levoglucosan; Arab, arabitol; Glu, glucose (2 peaks); Man, mannitol; C16, palmitic acid; C18:1, oleic acid; C18, stearic acid. Compounds 1 and 2 correspond to the previously unobserved 2-methyltetrols, i.e., 2-methylthreitol and 2-methylerythritol, respectively, arising from the photooxidation of isoprene (after Claeys et al. [2004]).

The facts that these compounds have a C5 isoprene skeleton and an estimated very low vapor pressure (<1.6 × 10−5 torr), and that they are

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enriched in the fine size fraction, they are not known to be present in composite leaves, and they occur as a mixture of diastereoisomers typical of non-enzymatic formation processes, unanimously suggest that these compounds are of secondary origin. 3.3.5 Modeling of SOA Formation 3.3.5.1 Regional Estimates of SOA Contribution to Organic Aerosol The most important feature of SOA is its highly dynamic nature. The time-scale of both emissions and atmospheric processes typically ranges from minutes to hours. Therefore to provide reliable estimates for the contribution of SOA to organic aerosol time-resolved OC and EC data are needed. To produce a large body of high-quality data is a demanding task in terms of sampling, analysis, and data evaluation. Therefore it is not surprising that the methods for estimation of the contribution of SOA to organic aerosol mass were first applied, tested and perfected in the specific region of the South Coast Air Basin of California (metropolitan Los Angeles), which has the world’s most extensive gas and particulate matter database against which to evaluate the performance of atmospheric models. The lack of direct chemical analysis method for the determination of either primary or secondary organic aerosol led to the development of different indirect approaches. Nowadays there are basically four methods to estimate the relative mass contributions of primary and secondary organic aerosol. 1) The diurnal variations of the OC/EC ratio can be used to estimate the range of secondary organic aerosol formation. In this concept, secondary organic aerosol should manifest itself in the increase in OC/EC ratio, particularly during the summer peak photochemical smog season, showing an increasing trend as the pollution plume moves with the prevailing wind [Turpin and Huntzicker, 1991]. The method assumes all EC to be a conservative tracer for primary combustion-generated OC emissions. Thus primary OC (OCprim) can be expressed as the sum of primary biogenic OC (OCbio) and combustion-generated primary OC as follows [Lim and Turpin, 2002]:

OC PRI = Z + (OC EC )PRI × EC

(3.29)

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where (OC/EC)pri is an estimate of the primary OC/EC ratio. Consequently, secondary OC is inferred from the difference between measured particulate OC and OCprim. It should be noted that with the exception of two studies the contribution of OCbio is usually neglected [Turpin and Huntzicker, 1995; Lim and Turpin, 2002]. A critical step of this method is the estimation of a characteristic primary OC/EC ratio. For this several approaches have been reported in the literature: a) (OC/EC)pri can be estimated as weighted average of primary emissions. Using this approach, Gray [1986] found a value of about 3.2 for TC/EC ratio, identical to the weighted average of highway emission profiles including exhaust of gasoline-powered and diesel vehicles, tire dust and brake dust. b) At locations under the dominant influence of primary source emissions and at times when minimum photochemical activity is expected, the measured average OC/EC ratio can be regarded as a good approximation of (OC/EC)pri. For example, using an in-situ carbon analyzer during the Southern California Air Quality Study in 1987, an (OC/EC)pri range of 1.52.9 was inferred [Turpin and Huntzicker, 1991]. However, it should be kept in mind that even if conditions appear to be such that no SOA formation is expected, there is a possibility of SOA produced previously may also be sampled. Correcting for such likely occurrences, a best estimate of 2.4 (range 2.05–2.75) was given for the primary OC/EC ratio in downtown Los Angeles [Strader et al., 1999]. c) Alternatively, the minimum OC/EC ratio during the entire study period at urban locations may be viewed as an estimate for (OC/EC)pri. Data in the lowest 5–10% by OC/EC ratio in Atlanta constrained (OC/EC)pri to be about 1.8, and set a reasonable upper limit of about 2.1 [Lim and Turpin, 2002]. On an annual basis, between 27 % and 38 % of the organic carbon was found to be of secondary origin at downwind locations such as Rubidoux (60 miles from downtown Los Angeles) [Gray, 1986]. In summer periods in Atlanta the mean SOA contribution was 46 % of the measured OC, but at times as much as 88 % of the 1-h average OC concentrations [Lim and Turpin, 2002]. The SOA concentration showed distinctive bimodal diurnal variations with peaks in the early afternoon and before dawn. The latter peak was attributed to changes in temperature and relative humidity that favor partitioning of semi-volatile oxidation products into the particle phase. In addition, downward mixing of aged particles from aloft could be held responsible for the elevated nighttime SOA concentrations.

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In the Los Angeles area, the estimated contribution of SOA to measured OC in winter was found to be around 20 % on a daily average basis, but for transient periods during the early afternoon it could reach as much as 55 % [Strader et al., 1999]. Sometimes relatively high contributions were observed at night. 2) Another method relies on a Lagrangian trajectory model (SOAM) that simulates the formation, transport, and deposition of SOA. The SOAM model included the emissions of the precursors, their atmospheric dispersion and photochemical reactions, as well as the formation, dispersion, and removal of SOA in the Los Angeles area [Pandis et al., 1992]. Inherent in the model was the assumption that gas-phase reaction products condense upon exceeding their saturation vapor pressure. The model predicted SOA contribution of 15–22 % for inland locations in summer, and only 5–8 % for the coastal areas. The model was sensitive to reactive organic gas emissions, aerosol yields and partitioning of condensable species between the gas and particle phases. On average, 10–40 % of the total particulate organic carbon was found to be of secondary origin, generated by the photooxidation of aromatics other than toluene, toluene, terpenes, alkanes and olefins in relative proportions of 46 %, 19 %, 16 %, 15 % and 4 %, respectively, on a daily average basis. When the model was modified to allow for the condensation of semivolatile species below their saturation vapor pressure, an upper limit of about 15 µg m−3 was obtained for SOA mass concentration in winter [Strader et al., 1999]. This was about a factor of 2 higher than that obtained by SOAM, and compared well with estimates based on the OC/EC ratio. The evolution of predicted SOA concentration at an inland location is shown in Figure 3.36. The high values in winter can be explained by the stable meteorological conditions that allow the accumulation of precursors and the low temperatures favoring partitioning of semi-volatile compounds into the aerosol phase. 3) The third approach for the estimation of the contribution of SOA relies on the detailed chemical speciation of organic aerosol, the determination of tracer compounds for all major primary emissions, and the use of source-receptor models. This method allows the source apportionment of primary OC, as described in details in Chapter 2.2. Then the maximum SOA concentrations can be calculated by subtraction from measured OC concentrations. This approach placed an upper limit of 31 % and 15–18 % on the annual average SOA contribution (for the year of 1982) at Rubidoux and downtown Los Angeles, respectively [Schauer et al., 1996]. However, the time-resolution of the method does not allow diurnal variations to be followed.

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Figure 3.36. Predicted SOA concentration evolution for air parcels arriving in Fresno on a typical winter day during clear-sky conditions (after Strader et al. [1999]).

It is notable that in regions with substantial anthropogenic influence, the annual average contribution of SOA estimated by the different methods falls within a fairly narrow range. These values, however, might be deceiving as they conceal huge variations on a seasonal as well as diurnal time-scale. Furthermore, they might suggest that SOA is not particularly important on a larger, say continental or hemispheric scale. In fact, the situation could be just the opposite. What is said to be an upper limit for SOA contribution to OC in any relatively confined area under strong anthropogenic influence is in fact more like a lower limit when viewed from a global perspective. This assumption have recently been supported by a few recent studies on the estimated contribution of SOA to organic aerosol in rural regions outside the South Coast Air Basin, which were based on measured OC/EC ratios. In Portugal, SOA contribution to fine organic aerosol was estimated to be 45 % and 68–78 % for winter and summer, respectively [Castro et al., 1999]. In a national park in Texas, between July and October, the calculated contribution varied between 42 and 98 % [Brown et al., 2002]. It should be noted that in the latter study the conclusions were strongly supported by the analysis of molecular tracers, too. By incorporating SOA formation from biogenic VOC into an EMEP Lagrangian oxidant model for northern Europe, α-pinene was found to be the largest single contributor to SOA formation (23.5 %) on an annual basis, albeit somewhat less than its mass emission share (35 %) would suggest

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[Andersson-Sköld and Simpson, 2001]. Limonene (21.3 %) and caryophyllene (12.4 %) followed suit, despite their substantially lower shares in emission (10 and 2 %, respectively) [Lindfors et al., 2000]. The overall contribution of biogenic secondary organic aerosol to OC was calculated to be in the range of 2–50 %, increasing from the coastal sites of southern Norway and Denmark inland and further north. In general, biogenic VOC contributed far more to SOA formation than anthropogenic VOC, even in the region of northern Europe. Recently the size distribution has been assessed by the modeling of the dynamic behavior of SOA and its interactions with primary organic particles and inorganic aerosol [Koo et al., 2003]. The model predicted size distributions are shown in Figure 3.37.

Figure 3.37. SOA size distributions predicted by the equilibrium model with the modifed (solid line with squares) and old (solid line with crosses) weighting schemes and dynamic model (solid line with triangles) at 17:00 PST (after Koo et al. [2003]).

On the basis of these studies it can be hypothesized that on a global scale, SOA likely predominates the atmospheric effects of fine organic aerosol. On a mass basis, primary OC may be more important, but most of its mass is confined to source regions. On the other hand, SOA formation has a considerable time-lag as compared to emission of primary aerosol (at least with respect to the residence time of accumulation mode aerosol). Therefore most if its mass forms outside the boundaries of the immediate source regions, both vertically and horizontally. In addition, biogenic sources of SOA precursors are understood to be predominant, which spread over much larger areas than major anthropogenic sources. It follows that the spatial distribution of SOA should be much more homogeneous than that of primary

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organic aerosol. Consequently in the major part of the troposphere SOA should dominate organic aerosol mass and number concentrations. This statement, however, strongly contrasts the results of a recent modeling study, which has found very low mass contribution of SOA over the entire boundary layer, and somewhat higher, but still minority contributions towards the upper troposphere (see Figure 3.38) [Chung and Seinfeld, 2002]. It should be recalled that this GCM model has been based on primary emissions from fossil fuel combustion and biomass burning, and SOA formation from biogenic hydrocarbons which followed the traditional scheme of SOA partitioning (for details see sub-chapter 3.3.5.2). 3.3.5.2 Global Model Estimates of Annual SOA Formation Globally vegetation is considered to be the largest contributor to the organic fraction of fine particulate matter, as a combination of primary emission and secondary aerosol formation [Seinfeld and Pandis, 1998]. Andreae and Crutzen [1997] estimated global SOA formation in the order of 30–270 Tg yr−1, comparable to the secondary sulfate aerosol. Griffin et al. [1999c] estimated the annual global amount of SOA formed from the oxidation of biogenic hydrocarbons based on experimental data on the aerosol forming potential of 14 biogenic compounds [Hoffmann et al., 1997; Griffin et al., 1999c], and the two-product model described by Odum et al. [1996]. Since the low mixing ratio of NO3 in remote areas, smog chamber data were refined by removing the estimated contribution of NO3 to the aerosol yield of four bicyclic alkenes, α-pinene, β-pinene, ∆3-carene and sabinene. The emission inventory of monoterpenes and other reactive volatile organic compounds was taken from Guenther et al. [1995] and Müller [1992], and were broken down to relative contributions of individual species as given in Table 3.5. The OH and O3 concentration fields were simulated in a global model [Müller and Brasseur, 1995]. Although the aerosol yield is expected to decrease with increasing temperature, no temperature correction of the smog chamber data (310 K) was applied. The global amount of secondary organic aerosol formed by the oxidation volatile biogenic compounds was estimated to be 13–24 Tg yr−1, admittedly a lowerbound estimate since partitioning into primary organic aerosol and anthropogenic secondary organic aerosol formation were not taken into account. With a residence time of one week, the predicted atmospheric burden of secondary organic aerosol equaled to 0.36 Tg, of the same order of magnitude as that of combustion derived primary carbonaceous aerosol as suggested by Liousse et al. [1996]. In a

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recent study isoprene photooxidation products 2-methyltetrols alone have been estimated to form in amounts of 2 Tg annually [Claeys et al., 2004]. Kanakidou et al. [2000] performed an interesting 3-D modeling exercise to follow the impact of human activities on biogenic secondary aerosol formation due to the increasing level of tropospheric oxidants since the pre-industrial times. They only considered secondary organic aerosol formation by the ozonolysis of α- and β-pinene, thus admittedly underestimating aerosol yields. To achieve a representation of other reactive compounds, effective rate constants were adopted for these two species from kinetic studies. The reactions of the monoterpenes with OH and NO3 radicals were taken into account only as a sink, affecting solely gas-phase chemistry. They also considered two cases separately: first, when partitioning of semi-volatile species was allowed only on secondary organic aerosol generated by biogenic VOC oxidation, and second, when they also allowed for partitioning on all preexisting organic aerosol, including all primary and anthropogenic secondary organic aerosol. The oxidation products were assumed to have the same affinity for these phases as for biogenic secondary aerosol. As sinks, a dry deposition velocity of 0.1 cm s−1 was adopted from a model study [Liousse et al., 1996], whereas for lack of available data the efficiency of wet deposition was assumed to be the same as that for the highly soluble sulfate aerosol. Emission of organic aerosol from fossil fuel (28.5 Tg yr−1) and from tropical biomass burning (44.6 Tg yr−1) given by Liousse et al. [1996] were treated as primary emissions and were distributed according to the work of Crutzen and Zimmermann [1991]. Global emissions of VOC by vegetation, estimated by Guenther et al. [1995] to be a total of 1150 TgC yr−1, were taken, broken down to monoterpenes (127 TgC yr−1) and the most reactive and aerosol-producing part of other reactive volatile organic compounds (83 TgC yr−1) according to Griffin et al. [1999c]. For the pre-industrial scenario, fossil fuel and biomass burning organic aerosol emissions were reduced to 10 % of their present value, but the emission of monoterpenes was kept constant in spite of the reported decrease of the Earth’s forested areas (by 20 % between 1850 and 1980) [Houghton, 1992]. For the present-day scenario, the model was able to reproduce the observed organic aerosol mass concentrations within a factor of 2–4 in most continental rural and forested areas. It also showed that secondary organic aerosol from ozonolysis of biogenic compounds contributed about 40–60 % to the total organic aerosol in the tropics and continental regions at northern mid-latitudes. A major uncertainty of the model came from the temperature dependence of the aerosol yield in particular of the partitioning coefficient which was shown to negatively correlate with temperature. Since the smogchamber experiments were conducted at temperatures higher than 40 °C,

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their parameterization may result a substantial underestimation of secondary organic aerosol formation. With this in mind, the present production rate of secondary organic aerosol from the ozonolysis of biogenic volatile organic compounds were calculated to be in the range of 61–79 Tg yr−1, allowing for the partitioning of semi-volatile species onto primary and anthropogenic secondary organic aerosol. If only self-nucleation and partitioning into biogenic secondary organic aerosol were allowed, the values were reduced by a factor of about three. This estimate was actually reasonably close to that given by Griffin et al. [1999c], who applied the same restrictions but a different model using fixed oxidant fields and considering the contribution of OH reactions to aerosol formation. The resulting global burden of biogenic secondary organic aerosol up to 100 hPa was found to be 1.2–1.6 Tg, which, when compared to the organic aerosol burden from fossil fuel burning of 0.09 Tg [Cooke et al., 1999], or the burden of SO42− (1.14−2.26 Tg) [Langner and Rodhe, 1991], highlights the great importance of biogenic secondary aerosol in the troposphere. Running the model for the pre-industrial scenario revealed that biogenic SOA concentrations were much lower at those times when relatively little aerosol formation occurred over the forested areas on the Northern Hemisphere before human activity injected vast amounts of NOx and hydrocarbons into the atmosphere, leading to significant increase in the levels of oxidants, in particular ozone. In addition, human activity also released primary organic aerosol into which semi-volatile species can partition. The maximum absolute increase was calculated for the tropics where both O3 and primary aerosol concentrations were increased by the extensive biomass burning. For the preindustrial period, a global production estimate of 16.5–27.9 −1 Tg yr was obtained from the model, which was about 3–4 times lower than estimates for the present. It follows that ∼75 % of current biogenic SOA production may be induced by human activity. Despite the much lower formation rate, biogenic SOA constituted a major fraction of organic aerosol in preindustrial times, because of the little primary organic aerosol emission. SOA may be a major component of aerosol number concentration and possibly also of cloud condensation nuclei (CCN) concentration in the preindustrial atmosphere. Today it may contribute ∼50 % to the total organic aerosol burden, whereas, though substantially weaker than nowadays, may have contributed ∼80 %. Andreae [1995] calculated a global mean aerosol burden and concentration, based on estimates of global emissions and lifetimes. The aerosol was assumed to be the combination of two lognormal, fine and coarse particle size distribution. In his classification, organic material formed

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by the secondary processes from VOC had a total column burden of 4.5 mg m−2, corresponding to a total mass concentration of 0.38 µg m−3 for an average tropospheric height of 12 km (2 Tg globally). There was another class of aerosol in his scheme, biomass burning aerosol, of which a considerable fraction is again carbonaceous. Its total column burden and total mass concentration were estimated to be 3.9 mg m−2 and 0.32 µg m−3, respectively. A recent model study on the global distribution of carbonaceous aerosol included a specific module for SOA formation [Chung and Seinfeld, 2002]. Biogenic volatile organic compounds were grouped into 5 classes according to their experimentally measured aerosol yields. The 14 individual oxidation products formed in the reactions by O3 and OH were allowed to partition into the entire organic phase. An equilibrium partitioning model was applied, and the temperature dependence of the partitioning coefficient was accounted for using the Clausius-Clapeyron equation. The predicted global SOA production was 11.2 Tg yr−1, with 91 % due to O3 and OH oxidation. Somewhat surprisingly, this estimate was in fact lower than the low bound of a previous model prediction (13–24 Tg yr−1) [Griffin et al., 1999c]. Unlike previous models, the model by Chung and Seinfeld also predicts the share of SOA in organic aerosol which gives ample opportunity to be tested against available observations. The predicted annual zonal average SOA percentage of OC is shown in Figure 3.38. This model distribution suggests that near the surface, OC is dominated by primary organic carbon, the share of SOA remains below 10 %. These predictions are in severe conflict with various estimates of SOA for rural and remote aerosol, which suggest more significant contribution of SOA over most continental areas (25–70 %, for details see Chapter 4). The increased SOA contribution predicted by the model in the upper troposphere has not been confirmed experimentally as yet, but the physics behind it is quite evident. The inclusion of acidic aerosol into the same global model has yielded an upperbound estimate for SOA formation of about 80 Tg yr−1 [Gao et al., private communication]. The significant increase in predicted SOA formation (by a factor of 7) may also help bring SOA relative contribution closer to available observations, though it has not been provided explicitly in the paper. This discrepancy between model results and observations can be resolved if we assume that other mechanisms of SOA formation are also at play. These mechanisms may include heterogeneous reactions of carbonyl compounds and dienes on acidic particles, or multiphase oligomerization reactions involving semi-volatile biomass burning products [Jang et al., 2002; Limbeck et al., 2003; Gelencsér et al., 2003]. All of these reactions

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can drive vast amounts of volatile or semi-volatile organic species irreversibly into the aerosol phase. In the traditional model of SOA formation, these species would only be allowed to partition into the aerosol phase to the state of equilibrium, which represents negligible contribution to the aerosol mass. For detailed discussions of these reactions, refer to subchapter 3.3.2.

Figure 3.38. Predicted zonal annual average of the SOA global percentage of total organic aerosol (SOA + primary OA) (after Chung and Seinfeld [2002]).

In any case, there is an urgent need to reconcile atmospheric models with field observations not only with respect to mass concentrations of organic carbon, but also to the distribution of derived parameters such as relative contribution of SOA to organic aerosol, which is readily available from the most comprehensive global models.

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3.3.6 Secondary Aerosol Formation in Heterogeneous and Multiphase Processes It is evident that the classical theory of SOA formation is unable to answer several fundamental questions and to explain the atmospheric or even smog-chamber observations, in terms of organic chemistry. The most critical issue is that a major fraction of SOA remains unidentified, even when a single compound reacts in controlled smog-chamber experiments. In other terms, even if we claim that we can understand the fundamental principles of SOA formation, chemically we can only follow the initial steps of the processes then rapidly loose sight of what is happening. The current and most widely accepted theory of SOA formation tries to circumvent this problem by putting quantitative SOA formation on a physical basis. The most important parameter, the aerosol yield that is used for atmospheric modeling, has so far received very little support from chemistry. It is more than likely that there are other processes at work that can provide answers to at least some of the questions. These are chemical reactions that take place on or in the aerosol phase, or in hydrometeors. These reactions do not only help follow chemistry along the course of the traditional theory of SOA formation, but may also provide new routes for SOA formation that have never been considered in global SOA modeling. It is possible that the inclusion of such new routes in atmospheric models will reshape our understanding on SOA formation and chemistry, and thus on organic aerosol in general. In the context of the reactions of condensed species, it is important to distinguish between heterogeneous and multiphase reactions [Ravishankara, 1997]. Heterogeneous reactions are constrained to the surface of aerosol particles, as diffusion in and out of the bulk is assumed to be negligible; on the other hand, multiphase reactions take place in the bulk of hydrometeors where diffusion is of importance. Although it is now accepted that in the troposphere a significant fraction of nss-sulfate is formed in multiphase reactions in hydrometeors, heterogeneous and multiphase organic chemistry is still in its infancy. This might be surprising in the light of the fact that there is no particular reason for excluding organic species from such reactions: many of them can readily dissolve in droplets or adsorb on solid particles, and most of them are quite reactive. Nevertheless we know very little about the aqueous-phase chemistry of organic compounds in the atmosphere, in spite of the fact that the possibility of VOC oxidation in cloud and fog droplets was first raised more than 20 years ago [Graedel and Goldberg, 1983]. Aqueous-phase reactions of OH in cloud and fog droplets were proposed as pathways for the oxidation of aldehydes to their corresponding carboxylic acids [Chameides,

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1984; Graedel et al., 1986; Cho and Carmichael, 1986; Pandis and Seinfeld, 1989]. Aqueous phase reactions with OH may also be involved in the loss of polycyclic aromatic hydrocarbons (PAHs) associated with wood smoke particles exposed to sunlight [Kamens et al., 1988]. It was demonstrated that atmospheric waters contain chromophores (light-absorbing substances and/or functional groups) that can absorb solar ultraviolet radiation. This then initiates aqueous-phase photochemical reactions that lead to the formation of oxidants and free radicals aloft. A major product of such reactions is hydrogen-peroxide, which accounts for 97 % of detectable peroxides in cloud water [Faust, 1994]. Midday equinox normalized H2O2 photoformation rates can be as high as 8.5 µM h−1. The chromophores responsible for this aqueous-phase photoformation were not identified, but photoformation rates correlated most strongly with normalized fluorescence, absorbance per unit path length and dissolved organic carbon in cloud water [Anastasio et al., 1994]. Organic chromophores may undergo photochemical reaction to yield HO2 and H2O2. The photolysis of Fe(III) with oxalate and other polycarboxylates such as malonate or citrate was also proposed as likely candidates [Faust, 1994]. These species may also be formed by the reduction of O2 by the hydrated electron that is photo-ionized from aromatic compounds. The aqueous-phase OH photoformation rates in cloud water were of the same order of magnitude as estimated rates of direct scavenging of OH from the gas phase or the aqueous decomposition of O3 scavenged from the gas phase [Faust, 1993]. Photoreduction of Fe(III) species into Fe(II) is a very rapid process which is responsible for the substantial share of Fe(II) in cloud and fog droplets or aerosol particles. Alternatively, a dynamic steady-state is maintained by the rapid photoreduction of Fe(III) and the rapid re-oxidation of Fe(II) by photooxidants. This rapid aqueous-phase photochemical redox cycling of Fe(III)/Fe(II), and its coupling to the multiphase chemistry of the atmosphere, bears conceptual similarities to the classical NO2/NO photochemical oxidant cycles in the gas phase [Faust, 1994]. Although these reactions were most extensively studied in cloud and fog water, they undoubtedly also occur in other tropospheric hydrometeors such as hydrated aerosol, rain or dew, even on wetted foliage. It might well be that analogous reactions occur in stratospheric particles, which might influence stratospheric chemistry. The first indications for the importance of heterogeneous organic reactions stemmed from the endeavor to reconcile smog-chamber observations with partitioning theory. Specifically, to explain how highly volatile compounds end up in the aerosol phase, heterogeneous reactions were invoked [Jang and Kamens, 2001]. Such reactions are expected to result in a drastic reduction in the equilibrium vapor pressure of the SOA products.

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The possibility of heterogeneous reactions was reinforced by thermal desorption particle beam mass spectrometric studies on SOA which implied the formation of large peroxyhemiacetals in the reaction between αacyloxyalkyl hydroperoxide and formaldehyde [Tobias and Ziemann, 1999]. In fact, heterogeneous loss of formaldehyde on sulfuric acid aerosol in the upper troposphere/lower stratosphere was reported earlier [Iraci and Tolbert, 1997]. The hypothesis that heterogeneous or multiphase reactions can lead to the formation of humic-like substances (HULIS) as major SOA components was put forward by Gelencsér et al. [2002]. Very recently, laboratory evidence has been presented on acidcatalyzed carbonyl chemistry on aerosol particles [Jang et al., 2002]. This chemistry includes various acid-catalyzed reactions, such as hydration, hemiacetal and acetal formation, aldol condensation, and polymerization in the aerosol phase, as shown in Figure 3.39.

Figure 3.39. Acid-catalyzed heterogeneous reaction mechanisms of atmospheric carbonyls (after Jang et al. [2002]).

For example, it is known that equilibrium between an aldehyde and its hydrate favors the hydrate form, which may further react with carbonyls to yield dimers, trimers, and polymers. Hemiacetals are often unstable in

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themselves, but their further reactions in the presence of acids lead to the much more stable hemiacetals. These reactions are essentially interactions between the organic and inorganic components of tropospheric aerosol. Furthermore, such processes may exert a driving force for the partitioning of highly volatile species into the aerosol phase, thereby increasing the aerosol yield considerably. However, it is important to point out that these reactions, even polymerization were thought to be reversible. That is, during sampling and analysis the reaction products return to their original precursors, rendering the products undetected. This also means that while they can explain anomalous gas-particle partitioning of volatile species upon SOA formation, they cannot account for the formation of refractory organic components which can be observed in atmospheric aerosol. It has been shown very recently that semi-volatile photooxidation products of 1,3,5-trimethylbenzene tend to polymerize in the aerosol phase upon aging for more than 20 hours, yielding regularly spaced ion masses up to about 900 Dalton in laser desorption/ionisation mass spectrometry measurements [Kalberer et al., 2004]. In contrast to the mechanism proposed by Jang et al. [2002], in this case organic acids formed in gas-phase oxidation reactions seem to be sufficient to catalyze acetal polymerization reactions. It should be noted, however, that as the gas-phase formation of dimers of cis-pinic acid observed by Hoffmann et al. [1998] was challenged as being a potential artifact in the mass spectrometer, caution must be exercised to exclude such a possibility. Further experiments and independent methods should be called in to ascertain the presence of higher oligomers in the aerosol phase in the case of such reactions. The first laboratory evidence for the irreversible formation of HULIS in heterogeneous reactions has been presented recently for the case of dienes like isoprene in the presence of sulfuric acid [Limbeck et al., 2003]. The reactions led to colored polymeric products. The humic-like character of the products was evidenced by UV-spectrometry, thermal analysis, and FTIR diffuse reflectance spectroscopy. These reactions are part of a conceptually new mechanism of SOA formation, assigning secondary aerosol formation potential to isoprene, a biogenic species which is emitted by far the largest quantities on a global scale. This mechanism assumes that isoprene, whose SOA formation has been thought to be negligible [Pandis et al., 1991], is readily taken up by highly acidic atmospheric sulfate clusters and processed to humic-like polymers (HULIS). It should be noted that these experiments were performed in bulk, and no smog-chamber studies have confirmed their results as yet. If proven, however, this hypothesis can present a breakthrough in our understanding of global SOA formation.

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The irreversible formation of HULIS in atmospheric multiphase reactions has also been evidenced recently in the laboratory [Gelencsér et al., 2003]. The precursors studied were aromatic hydroxy acids, which are emitted in vast quantities by the pyrolysis of lignin during biomass burning. It was shown that even a single representative compound can react with OH radicals yielding colored products under typical conditions prevalent in cloud water. The reactions take place within a few hours in solution, which implies that the process does have atmospheric significance. The first step can be the formation of phenoxy radicals which are then stabilized by delocalization of the unpaired electron over the aromatic ring and its relevant substituents. This step may be followed by radical dimerization and oligomerization to furnish higher molecular weight products, as shown tentatively in Figure 3.40.

Figure 3.40. a) Formation of a phenoxy radical and possible ‘resonance-stabilized’ delocalized structures. b) Coupling derivatives of the phenoxy radicals and formation of a highly colored (V) radical (after Gelencsér et al. [2003]).

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It was also shown that the presence of such products in cloud water may enhance absorption of solar radiation at a wavelength of λ = 0.475 µm by a factor of 6. A follow-up study on the molecular weights of the reaction products has revealed a continuous distribution well below 1000 Dalton, which confirms the high degree of similarity to atmospheric HULIS [Hoffer et al., 2004]. In addition, this study has proven conclusively that the process is oligomerization rather than polymerization. In the same study, the results implied that HULIS consist of condensed and partially oxidized (e.g. quinone-like) phenolic structures crosslinked with short-chain aliphatic bridges which form by the oxidative cleavage of the phenolic ring. A recent smog-chamber study on the effect of acidic seed on α-pinene and ozonolysis has proven conclusively that acidity promotes SOA formation and increases aerosol yield by up to 40 % [Iinuma et al., 2004]. An even more important finding of this study is that SOA primarily consists of higher molecular weight species in the MW range of 300–1200 Da, and the distribution shifts towards larger MW with increasing seed acidity. Based on the chemical structures of the larger molecules polymerization and aldol condensation in the aerosol phase have been suggested as possible mechanisms for their formation. Although in this paper the term HULIS is not used, the apparent similarity of the species to HULIS (or vice versa) in their molecular weight distribution may entitle us to use these terms interchangeably. It should be mentioned that as regards the sources of HULIS MayolBracero et al. [2002] presented conclusive evidence that biomass burning is a major source of HULIS, however, without an indication of its possible formation mechanism. Based on the references cited in that paper, a primary origin is postulated here. Based on the recent findings presented in Chapter 4 as well as in the discussion so far, a scheme has been constructed showing all the available theories on atmospheric HULIS formation, as shown in Figure 3.41. Although experimental and/or observational evidences back up all the theories shown here, but there has been only one conclusive direct study, i.e. the observed oligomer formation in smog chambers. It is possible that to certain extent all mechanisms could be operative—this would explain the ubiquitous nature and abundance of HULIS in continental fine aerosol. What is known for certainty is that biomass burning is a source of HULIS— however its secondary origin still awaits experimental confirmation.

Secondary Organic Aerosol

Figure 3.41. Possible atmospheric pathways for HULIS formation.

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Chapter 4 ORGANIC CHEMISTRY OF AEROSOL In the field of carbonaceous aerosol there are generally two basic approaches for organic aerosol characterization. The first approach involves detailed organic speciation, primarily by means of gas chromatography-mass spectrometry (GC-MS). Very often the goal of such speciation studies is to obtain information only on specific target compounds, let they be potentially toxic compounds, important source tracers, or both. Most of these studies do not—and cannot—pay attention to the concentration and physical-chemical properties of organic aerosol as a whole. The other approach is based on the directly measurable concentration of total and organic carbon in aerosol. These quantities—after accounting for the mass of other elements in organic species—can be readily incorporated into aerosol chemical mass closures. A close companion of these direct measurements is the characterization of the physico-chemical properties of the bulk organic (carbonaceous) aerosol. These studies can supply information more relevant in the atmospheric effects of carbonaceous aerosol, but without regard to traditional organic speciation. Despite earlier endeavors to bring organic aerosol speciation to completion sometime in the future, there now seems to be a large and insurmountable gap between the two approaches. In other terms, complete organic speciation in atmospheric aerosol is theoretically impossible due to the complexity and nature of aerosol-phase organic species. Nevertheless, there has been a remarkable progress in organic speciation studies for the last 10 years. This has been largely achieved by a “change of paradigm”, that is the renewed search for water-soluble and polar organic compounds which are not directly accessible by GC-MS. The determination of such compounds has helped cover a substantial fraction of particulate organic carbon under certain conditions, though still the majority of the organic carbon mass has remained unaccounted for. Most recently, the introduction of the concept of “humic-like substances (HULIS)”, or “organic macromolecules”, or “polycarboxylic acids”, has been claimed to resolve most of the unexplained mass of particulate organic carbon. However, strictly speaking, these organic compound classes cannot be speciated, their presence can only be inferred from various bulk aerosol measurements. Although by now there is 149

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conclusive evidence that such compounds do exist in the aerosol phase in close analogy with soil or aquatic organic matter, their quantitative determination is still subject to great conceptual and experimental uncertainties. For the above reasons, this chapter dealing with the chemistry of organic aerosol will be divided into two parts: the first sub-chapter will address organic aerosol as a bulk, without regard to organic speciation. The quantitative information here is more element-specific rather than compound-specific, but will provide a general understanding on the importance and characteristics of organic aerosol in various environments, in comparison with those of the major inorganic aerosol species. Perhaps surprisingly, information on organic species obtained by individual particle analyses is also included in this sub-chapter. This is because at the level of today’s technology, such instrumental techniques are generally incapable of detailed organic speciation. Instead, they may provide invaluable information on the state of mixing of organic species, which is a key parameter in fundamental atmospheric processes. There follows the physical and chemical characterization of the bulk organic aerosol with various methods, some of which are also used for organic speciation. Even some tentative source apportionment of organic aerosol is feasible at this bulk level, as will be shown for the aerosol carbon isotopic composition. The second sub-chapter will be devoted to the complex issue of aerosol speciation. As the issue is truly multidimensional, there seems to be no single salutary approach to its description. Organic aerosol speciation could be presented by compound classes or by fundamental source types, but after all, it is difficult to bridge the orders of magnitude differences between the mass concentrations of various compounds and aerosol types. For the structure of this sub-chapter, I have selected a “hybrid approach” which makes use of both organizational principles. The most ubiquitous and abundant compound classes for which ample data are available will be introduced separately. Generally these compounds cannot be attributed to single sources, though their relative concentrations may carry some indications for their possible origin. In the second part, the compound-specific structure will be switched to a source-specific one. Here the major organic aerosol sources will be presented with a view to the emission of specific organic tracers. Albeit these tracer compounds generally do not contribute significantly to the organic aerosol mass, being specific, they convey important source information when detected in atmospheric aerosol. The concept of the source apportionment by organic tracers will also be introduced and illustrated with some examples. Water-soluble organic species and some of their properties will be treated separately at the end of the second sub-chapter. This is because water-

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solubility is a key property in atmospheric processes, which also governs important sinks for aerosol species, such as cloud or precipitation scavenging. This brings us to the organic chemistry of cloud water and precipitation, which will be briefly addressed here.

4.1 Chemical Properties of Bulk Organic Aerosol 4.1.1 OC Mass Concentrations The determination of carbon as an element in aerosol on carbon-free substrate is straightforward and can be performed with a high degree of accuracy. However, the translation of the measured data into concentrations of particulate organic carbon is subject to large uncertainties or biases. These uncertainties and biases can be grouped by two fundamental questions. The first is to what extent the carbon that is measured was in particulate form in the atmosphere. This leads to the issue of sampling artifacts, which—unless properly accounted for—may severely affect a large fraction of organic compounds of relatively high volatility. This issue is discussed in details in sub-chapter 2.1. The second question relates to the uncertainty of differentiating between organic and elemental carbon, which is treated separately in sub-chapter 2.2. Since this uncertainty is usually more critical for elemental than for organic carbon, most biases accompanying particulate organic carbon concentrations are in fact related to sampling artifacts. There is an immense body of data on OC mass concentration in atmospheric aerosol in various environments and in different periods of the year, most of them from surface measurements. Among them, however, there are very few data on the mass size distribution of organic carbon in atmospheric aerosol. A comprehensive account of these data available would fill in volumes and would certainly dominate this book. Instead, a compilation in Table 4.1 attempts to illustrate the very large spatial and temporal variability of OC mass concentrations in the troposphere. In fact, apart from urban sites, the mass concentration of OC is not as variable as other aerosol properties. The concentration values at rural and clean marine sites are largely within an order of magnitude, which is— considering the relatively short residence time of aerosol particles in the troposphere—could be an indication of significant natural sources of large geographical extent and of the sizable contribution of secondary organic aerosol formed in the atmosphere. There are, however, a few issues which should be mentioned in this context. First, observational data may be loaded with considerable sampling artifacts. These are most often adsorption (positive) artifacts, which are particularly pronounced at clean sites and/or for short sampling times (see detailed discussion in sub-chapter 2.1). Unless properly corrected for during

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sampling, this artifact results in a significant positive bias in measured concentrations. The magnitude of this bias can be assessed by comparing available measurements with and without artifact correction. Table 4.1. Observed OC mass concentrations Location

Longitude

Latitude

Time period

OC concentration (µg/m3)

Reference

79°W

39°N

August

2.190

Japar et al. [1986]

8.4°W

40.3°N

August

3.506

84.7°W

45.5°N

December– April

1.960

19.5°E

46.9°N

July–August

7.300

79°W

40°N

August

2.510

June– September 1990

2.55 (0.1–7.37)

Castro et al. [1999] Cadle and Dasch [1988] Molnár et al. [1999] Japar et al. [1986] Malm and Gebhart [1996]

Rural Allegheny Mountain, Pennsylvania Anadia, Portugal Cheboygan County, Michigan K-puszta, Hungary Laurel Hill, Pennsylvania Tahoma Woods, WA, USA Marine, remote Sargasso Sea

50°W

30°N

July

0.440

Puerto Rico

62.5°W

18.2°N

March

0.500

Ecuador

77.3°W

2°S

Annual

0.510

Peru

76°W

10°S

0.160

Cape Grim

144.7°E

40.7°S

March– April Annual

Jungfraujoch, Switzerland Atlantic Ocean

7°59’E

46°33’N

7–15°W

29–41°N

June 1995– July 1997 June–July

0.32 (0.005– 0.94) 0.800

0.230

Chesselet et al. [1981] Novakov and Penner [1993] Andreae et al. [1984] Cachier et al. [1986] Andreae [1983] Lavanchy et al. [1999a] Novakov et al. [2000]

For example, with careful artifact correction OC concentrations of ∼1.1 µg m were determined in the clean marine boundary layer [Mader et al., 2002], which are at about the same as reported previously for similar sites. There was another recent study which used the same methodology for sampling and analysis (with artifact correction) over the Indian Ocean, −3

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probing polluted air masses from the Northern Hemisphere as well as pristine air masses from the Southern Hemisphere [Neusüss et al., 2002]. While the differences between EC mass concentrations were about two orders of magnitude OC mass concentrations varied only by a factor of 5–10. These findings imply that sampling artifacts may not contribute significantly to measured mass concentrations at remote sites. Another potential problem at pristine locations—especially for shortterm measurements—can be contamination during sample handling, transport, and analysis. Although in practice blank samples are used to detect and correct for such effects, they cannot be fully excluded as total organic carbon (TOC) analysis provides no information on the chemical nature of organic carbon in aerosol. Contamination may generally cause a positive bias in low-level carbon measurements. In addition, most OC measurements are not resolved by particle size, and the reported concentration data refer to the total suspended particulates. Albeit it is known from a few size-resolved OC measurements that most of the OC is associated with fine aerosol, at times the contribution of the coarse mode can be significant. As it will be shown below and discussed in details in sub-chapter 3.2, organic surfactants may become associated with sea salt particles upon bubble-bursting during sea salt particle formation, causing a second coarse mode in the OC mass size distribution [Neusüss et al., 2002]. It means that similar total OC mass concentrations may conceal important differences in OC size distributions and consequently in their atmospheric effects. There is another reason why OC mass concentrations at high-latitude remote sites are not as low as might be expected. The same applies to concentrations at higher altitudes and in the free troposphere, as will be addressed below. The low temperatures and high relative humidities favor partitioning of semi-volatile organic species or even organic gases into the aerosol phase, thereby increasing the bulk OC mass in aerosol. Generally this volatile fraction of the particles is prone to severe losses upon airborne sampling unless proper precautions are taken. However, recently there have been indications that heterogeneous and multiphase reactions irreversibly convert the volatile species into higher molecular weight SOA components. For detailed discussion see sub-chapter 3.3. OC at higher altitudes—vertical profiles of OC mass concentrations For most fundamental atmospheric phenomena affected by the presence of organic aerosol, it is crucial to know how organic aerosol is distributed vertically in the troposphere. In other terms, it is the scale height of OC that is of interest, especially when related to those of other key aerosol species. Furthermore, the assessment of the contribution of organic aerosol to

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the mass of the tropospheric background aerosol would be important from a global perspective. Unfortunately, very little data are available to provide reliable answers to these fundamental questions. The existing data are clearly out of proportion either with surface measurements, or with the observed contribution of OC to aerosol mass. The main reason for this is that the role of organic aerosol in cloud condensation had been deemed insignificant until the mid-1990s when compared to those of the major inorganic species such as sulfate. Additional reasons can be the generally poor time-resolution of aerosol sampling for organic measurements (due to the limited sensitivity of analytical techniques), as well as the fact that OC measurements are loaded with large sampling artifacts under the conditions of airborne sampling. Airborne measurements on the East Coast of the United States revealed the ratio of TC/PM5 as a function of altitude, as shown in Figure 4.1 [Novakov et al., 1997]. To correct for possible sampling artifacts, the dualfilter strategy was applied for aerosol sampling.

Figure 4.1. Altitude dependence of the carbonaceous aerosol mass fraction. The line shown is derived from robust locally weighted regression smoothing analysis (know as LOWESS) (after Novakov et al. [1997b]).

The mean value was found to be 0.5 ± 0.33 which showed the predominance of carbonaceous matter in the aerosol mass balance, especially at lower PM5 concentrations (<15 µg m−3). In the continental boundary layer this ratio was found to be significantly lower (0.1–0.4) [Malm et al., 1994]. Therefore these observations can be interpreted that carbonaceous particles play a substantially larger role aloft than that inferred from ground

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measurements. This statement also implies that ground-based measurements can significantly underestimate the column organic aerosol mass budget. The submicron OC mass concentration was measured in the free troposphere on Tenerife in summer 1997 [Putaud et al., 2000]. Although the uncertainty of the determinations was high (± 74 %), the average background concentration of OC was found to be 0.20 ± 0.10 µgC m−3. It is worthy of note that this average value is virtually identical to that determined for the background marine boundary layer (see Table 4.1). On a mass basis, OC was clearly the predominant single component (43 ± 20 %). At the Jungfraujoch in summer 1998, the mean PM2.5 OC mass concentration was found to be 1.05 µgC m−3 in the free troposphere, though admittedly the site was periodically influenced by aerosol from the planetary boundary layer [Krivácsy et al., 2001a]. For this reason this value is likely an overestimate for free tropospheric background OC concentrations. Although only semi-quantitatively and on a relative scale, in-situ measurements of individual aerosol particles in upper tropospheric aerosol particles larger than 0.2 µm confirmed the predominance of organic matter over sulfate [Murphy et al., 1998]. These observations pointed out the importance of organic compounds in governing the optical and cloud nucleating properties of tropospheric aerosol at high altitudes. 4.1.2 Global Modeling of OC Concentration Distributions To date there have been a very few general circulation model (GCM) simulations on the global distribution of organic aerosol that comprise all major types of organic aerosol, including primary organic aerosol emitted by fossil fuel combustion, biomass burning and vegetation, as well as SOA formed by the photooxidation of anthropogenic and natural VOC. Such atmospheric models usually reflect the consolidated views on atmospheric species, ranging from emission inventories through the chemistry of their transformation to dry and wet deposition processes. The reliability of these models has been successfully demonstrated for the climatology of some aerosol species such as sulfate or BC, and for the prediction of their atmospheric effects. However, the application of such simulations for organic aerosol is a real challenge. The most critical point is that the majority of bulk organic matter in aerosol has not been identified, and those which have been exhibit a wide range of physical and chemical properties. The semi-volatile nature of many organic compounds complicates the issue, making both observations and simulations more difficult and much less reliable. In general, there are large gaps in our understanding of organic aerosol in virtually every field that needs to be incorporated into general circulation models. The reader is referred to the individual chapters of this book for more details.

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Given the complexity of such an enterprise it is not surprising that there have been only three studies on modeling global distribution of organic aerosol which include all major sources. The first model study of this kind was based on detailed emission inventories for organic particles from various biomass burning sources, but an OC budget scaled to BC emissions for fossil fuel burning, and a constant aerosol yield (5 %) attributed to α- and β-pinene oxidation [Liousse et al., 1996]. In fact, the emission of primary organic particles from plants, soil, and sea surface was not considered, so strictly speaking the global budget of OC was not complete. The model used (Grantour) was a Lagrangian model of transport, transformation, and removal, and was based on the wind and precipitation fields of another model (CCM1). The simulated average global distributions of the mass concentrations of organic matter1 in the months January and July are shown in Figure 4.2.

Figure 4.2. Global distribution of the surface concentrations of organic matter when all sources are included (in ng/m3) (a) for January, (b) for July (after Liousse et al. [1996]). 1

A conversion factor of 1.3 was used for OM/OC.

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The spatial distribution of the sources over the continents causes high concentrations of OC to spread widely over and off the continents in both hemispheres. However, the total abundance of OC is higher by a factor of 10 in the Northern Hemisphere than in the Southern Hemisphere. This is mainly due to the predominance of fossil fuel and natural emissions in the Northern Hemisphere, since biomass burning is distributed nearly evenly between the two hemispheres in the tropics. The vertical distribution of the model-predicted annual zonal average mass concentrations of organic matter is shown in Figure 4.3. Although it is difficult to directly compare simulated concentrations with observations, the authors found a reasonable agreement between them for some rural and remote sites in both hemispheres.

Figure 4.3. Predicted annual average and zonal average concentrations of organic matter (in ng/m3) (after Liousse et al. [1996]).

The global distribution of organic aerosol mass concentrations was also simulated in a relatively recent model, which—by relying on the IPCC inventory for biomass burning and fossil fuel OC emissions, and fixed aerosol yield (5 %) for monoterpenes—was much less developed than that by Chung and Seinfeld [Koch et al., 2001]. The model agreed with observations only within about an order of magnitude. In addition, it did not support increasing OC contributions (OC/S ratios) with altitude over most of the

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troposphere, a likely occurrence that the few available measurements do suggest. However, to date this model has been the only one which predicts future scenarios for the OC budget of the troposphere based on the IPCC’s future inventories. The projections are given in Table 4.2 for the years 2030 and 2100. Table 4.2. Organic carbon and sulfate budgets (after Koch et al. [2001]). 2000 Organic Carbon Sources, Tg OM/yr Industrial emissions Terpene emissions Sinks, Tg OM/yr Dry deposition Wet deposition Burden, Tg OM Lifetime, days

2030

2100

81.48 8.37

108.72 8.38

189.73 8.38

−29.75 −60.10 0.95 3.86

−38.42 −78.42 1.22 3.80

−65.16 −132.95 2.16 3.97

2.01 14.79 43.26

3.26 24.40 57.22

1.75 14.42 41.69

−10.75 −49.31 0.85 5.15

−15.17 −69.70 1.28 5.50

−10.18 −47.69 0.80 5.03

Sulfate Sources, Tg S/yr Industrial emissions Gas phase Aqueous phase Sinks, Tg S/yr Dry deposition Wet deposition Burden, Tg S Lifetime, days

These projections, which may be viewed as first estimates for future changes of the organic aerosol burden, suggest an increase in the global burden perfectly in proportion with the increase in industrial emissions. The secondary aerosol term they consider in the model is rather low and is not affected by changing oxidizing potential of the atmosphere [see Kanakidou et al., 2000]. Furthermore the mixing state, solubility, and bulk properties of organic aerosol are assumed not to change in the future as reflected in the nearly constant lifetime of the particles. In the light of the most recent findings in the field, none of these assumptions seem to be tenable. It should be admitted, however, that parameterization of current knowledge into atmospheric models—which is not yet feasible—would enormously increase the complexity of the system, possibly far beyond the level that even the most advanced GCM could cope with.

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4.1.3 Size Distribution of Organic Aerosol Unfortunately, there have been very few studies in the literature that explicitly address the mass size distribution of organic carbon. In determining the size distribution of carbon mass concentrations, Hoffman and Duce [1977] found that over 80 % of carbon mass were in particles with diameters of <2 µm. The general view has long been that the fine mode carbonaceous aerosol is of anthropogenic origin, whereas the organic carbon on the large (>2 µm) particles is released from natural sources [Chesselet et al., 1981]. In rural aerosol the size distribution of TC was found to be unimodal in the accumulation mode, except in summer when non-negligible carbon concentrations were measured in the smallest size particles (cutoff 30 nm) [Temesi et al., 2003]. This secondary peak, however, did not appear in the size distribution of WSOC, indicating that most WSOC likely formed in cloud-processes. Unfortunately, in this study the carbon concentrations were not determined in the coarse mode, as an upper cutoff of 1.70 µm was used. However, previous TC size distribution studies at the same site showed that less than 5 % of the TC mass was above this size limit [Temesi et al., 2001]. The mass size distribution of TC was also studied at the high-alpine site Jungfraujoch in weekly impactor samples collected during summer [Krivácsy et al., 2001b]. It should be noted that the distributions included that of BC which was found to be significant relative to OC in the samples. Two characteristic types of distributions obtained are shown in Figure 4.4.

Figure 4.4. Size distribution of carbon at the Jungfraujoch. Sampling periods: (A) 15–21 July 1998; (B) 21–28 July 1998 (after Krivácsy et al. [2001b].

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Distribution A shows a bimodal distribution within the accumulation mode, the first and second maxima corresponding to the condensation and droplet mode, respectively. This likely reflects the substantial influence of the planetary boundary layer, which periodically affects the site during summer. In contrast, the distribution B exhibits a single maximum within the accumulation mode, and another small one in the coarse mode. This type of distribution is typical of aged aerosol, with non-negligible contribution of (possibly) Saharan dust. Due to the long sampling times and the lack of control to shut off sampling during the influence of planetary boundary layer aerosol, possibly neither distribution represent free tropospheric aerosol. Nevertheless the overwhelming majority (90–95 %) of TC is present with particles with d <2.5 µm. The importance of carbonaceous species in the accumulation mode at the high-alpine site Mt. Sonnblick (3104 m a.s.l.) was also confirmed by TC size distribution measurements with a Berner-type low-pressure cascade impactor [Hitzenberger et al., 1999]. Simultaneous measurements of major inorganic ions revealed a relative enrichment of carbon species in particles below 0.5 µm, as shown in Figure 4.5.

Figure 4.5. Mass size distribution of major ions (NH4+, SO42−and NO3−) and total carbon (TC) obtained for the impactor samples in Mt. Sonnblick (after Hitzenberger et al. [1999]).

In aerosol collected over the Indian Ocean during the INDOEX campaign, the mass size distribution of OC exhibited a bimodal distribution with another maximum in the diameter range of 2–4.1 µm, corresponding to the sea salt mode [Neusüss et al., 2002].

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4.1.4 Organic Aerosol in Mass Closure Studies As it was pointed out in the introduction, the organic (carbonaceous) and inorganic components of the aerosol are so intimately connected that the chemical mass closure of atmospheric aerosol merits some discussion. Obtaining mass closures is a prerequisite for accurate assessment of the contribution of major aerosol species to various atmospheric effects. Since there is absolutely no hope that the mass balance can ever be closed at the level of individual organic species (see discussion in the introductory part of this chapter), such closure is generally based on the measured OC concentrations, which are relatively straightforward to obtain. Since, however, organic species also contain other elements besides carbon, a mass conversion factor should be applied to obtain a quantity (organic mass) that fits directly into an aerosol mass balance. Accurate determination of such a factor would refer back to the complete identification and quantification of individual organic species, or would require complete preparative-scale separation of the organic component. None of these options are available, so this factor must be estimated from chemical information, or inferred from measurements. Although the value is constrained from the higher end by the overdetermined set of parameters (i.e. the calculated total mass of all species should not be higher than that determined gravimetrically), its estimation still introduces considerable uncertainty into mass closure studies. Some studies bypassed this problem by putting organic carbon (OC) mass concentrations directly into the balance, and assigning mass of other elements in the organic compounds into the fraction n.d. (not determined). The atmospheric effects of aerosol are then normalized to organic (or elemental) carbon mass concentrations (and not to mass concentrations of organic matter). This approach has the great advantage that it remains on the ground of measured parameters which are deemed more reliable than estimates (though they may not be always so). On the other hand, this approach is not so “elegant” in the sense that it leaves a much larger mass fraction unaccounted for, and makes estimation of other important aerosol components (e.g. water) much more uncertain. An excellent review on fine particles in the global troposphere was presented by Heintzenberg [1989]. The average chemical compositions of the particles was reported for urban, non-urban (rural), and remote (oceanic) environments, as shown in Figure 4.6.

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Figure 4.6. The average mass composition of aerosol at (a) urban, (b) rural, and (c) remote locations (after Heintzenberg [1989]).

Note that in the balance the mass of carbon is given, and not that of organic matter. Nevertheless the importance of organic compounds is obvious for the continental sites, and to a lesser extent also for remote sites. In another frequently cited study on fine aerosol chemical mass closure at coastal, rural, and polluted sites the organic matter was divided into water soluble (WSOC) and water-insoluble (WINSOC) fractions [Zappoli et al., 1999]. The mass of WSOC and WINSOC was estimated from carbon measurements using the conversion factors of 1.4 and 1.2, respectively. The average mass balances for the three sites are shown in Figure 4.7. While for the coastal site the total calculated mass exceeded that measured gravimetrically, the non-determined fraction was significant at the other two sites, in particular at the rural site. It is remarkable how well the mass balance for the rural site agrees with that reported for non-urban sites by Heintzenberg [1989], if the appropriate conversions are accounted for (see Figure 4.6/b).

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Figure 4.7. Average mass balance of the aerosol at the three sampling sites. (a) Aspvreten; (b) K-puszta; (c) San Pietro Capofiume. EC—elemental carbon; WINSOC—water insoluble organic compounds; WINSIC—water insoluble inorganic carbon; WSOC—water soluble organic compounds; other ions, ND—not determined (after Zappoli et al. [1999]).

The chemical mass closure of submicron free tropospheric and marine boundary layer aerosol was studied at Tenerife during ACE-2 [Putaud et al., 2000]. For the estimation of the organic mass, the conversion factor of 1.7 ± 0.3 was used, which encompassed most values reported in the literature. It should be noted that this value is substantially higher than those used in Figure 4.7 by Zappoli et al. [1999]. The mean compositions of submicron aerosol for four different conditions are shown in Figure 4.8.

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Figure 4.8. The mean composition of submicron aerosol in various environments. MBL— marine boundary layer; FT—free troposphere (after Putaud et al. [2000]).

It was speculated that the background free tropospheric aerosol results from the dilution of continental boundary layer aerosol injected into the free troposphere by convective transport through precipitating clouds. In any case, organic carbon was relatively more abundant in the free troposphere than at any site in the continental boundary layer. The chemistry of organic aerosol in the free troposphere, however, remains unknown, as to date no such chemical characterization study has been available. The in-situ measurement of individual aerosol particles (above diameters of 0.2 µm) at altitudes between 5 to 19 km by laser ionization-time-of-flight mass spectrometry as reported by Murphy et al. [1998] established the significance of organic compounds. The relative contribution of organic aerosol was found to be less significant in the marine boundary layer, but there are possibly much larger differences in the chemical composition and origin of organic matter when compared to those in the continental boundary layer or in the free troposphere.

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Mass conversion ratio between organic carbon and organic aerosol One of the most important uncertainty factors in mass closure calculations, the conversion of the mass of organic carbon to the mass of organic aerosol merits some discussions. In fine aerosol collected at urban and rural (receptor) locations in the Los Angeles area, the organic mass to organic carbon mass conversion ratio was estimated to be between 1.2 and 1.4 [Gray, 1986]. Using these values, the mass balance of fine aerosol was closed within ± 11 % of the gravimetrically determined mass. These values were long used for mass balance calculations. Recently, however, these values have been challenged by several studies, especially for use for aerosol in non-urban areas. It was hypothesized that larger contributions of bioaerosol particles as well as the highly oxidized biogenic and anthropogenic SOA in rural areas may lead to much higher values [Turpin and Lim, 2001]. The ratio of 1.9–2.3 was suggested to be more accurate for aged aerosol, and 2.2–2.6 was thought to better represent aerosol heavily impacted by woodsmoke. These latter values can be rationalized by the abundance in levoglucosan in wood smoke, which have an organic mass to carbon mass ratio of 2.25. Even the ratio suggested for urban aerosol increased slightly (1.6 ± 0.2). Whereas the principles behind higher ratios have gained widespread acceptance, most authors have refrained from using too high values. The “consensus value” has become 1.7 ± 0.3 [Putaud et al., 2000; Gao et al., 2003]. An experimental study that isolated a considerable fraction of WSOC in rural aerosol for direct determination of this ratio inferred a value of 1.9– 2.0, which is at the higher end of the consensus range [Kiss et al., 2002]. A recent systematic approach based on FTIR measurements, which broke down individual organic compounds to functional groups and carbon chains, attempted to detect most of the organic mass that remains unresolved by speciation techniques [Russell, 2003]. The uncertainty of the organic mass measurement was declared to be about 25 %. By analyzing aerosol samples collected on stretched Teflon filters during ACE-Asia by FTIR, and calibrating with standards of alkane, alkene, aromatic, alcohol, carbonyl, organonitrogen and organosulfur compounds for functional groups, the frequency distributions of OM/OC ratios shown in Figure 4.9 were obtained for fine organic aerosol. The estimated OM/OC ratios span a range of 1.1–2.2, with most values fall between 1.2 and 1.6 for all samples considered. Average values fall slightly below 1.4, indicating that there is a predominance of unsaturated alkene and aromatic (including polycyclic) structures, and photochemical oxidation of hydrocarbons did not occur to a great extent during atmospheric transport for the bulk of organic matter.

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Figure 4.9. Frequency distribution of OM-to-OC ratios for submicron aerosol samples measured by FTIR spectroscopy (i) on the NCAR C130 during PELTI (black), (ii) on the C130 during ACE-Asia (gray), and (iii) on the R/V Ronald H. Brown during ACE-Asia (white) (after Russell [2003]).

These low values found specifically for samples from ACE-Asia, would hardly overturn the consensus that seems to be reached in the literature. It is highly probable that a single universal value does not exist for this ratio: for urban and polluted aerosol a value between 1.4 and 1.6 may be realistic2, whereas for non-urban aerosol the range between 1.7 and 2.0 may better approximates reality. It is difficult to believe that the observed high share of WSOC in organic aerosol over most of the troposphere (see subchapter 4.3) could sustain such a low abundance of polar functional groups as would be inferred from an average ratio as low as 1.4. 4.1.5 State of Mixing of Organic Aerosol as Inferred from Individual Particle Analyses As it has already been pointed out, pure organic aerosol particles are unlikely to exist in the atmosphere, not even in the vicinity of aerosol sources. Therefore in order to correctly attribute atmospheric effects to these aerosol components, knowledge of the state of mixing of aerosol is required. Until now, most chemical information on organic aerosol came from bulk aerosol measurements, which cannot reveal how the organic species are associated with other aerosol constituents. Electron microscopy has long been used for the characterization of individual aerosol particles. Generally, however, the conditions used in 2

Note that the average ratio found by Russell fell to the lower limit of the consensus range.

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electron microscopy are highly unfavorable for the observations of organic species. In particular, the high vacuum of the electron microscopy could induce significant losses of semi-volatile organic compounds. Furthermore, those compounds which persist in the vacuum are often damaged and volatilized off by the focused electron beam under the electron microscope. On top of these, the technique had long lacked the potential of obtaining element-specific chemical information. For these reasons organic species or particulates were not favorite targets in electron microscopic studies. The only exceptions were bioaerosol particles whose morphologies and properties have long been studied by electron microscopy. Nevertheless, even the earliest electron microscopic observations of aerosol particles could reveal some interesting phenomena which were attributed to the presence of organic constituents. Husar and Shu [1975] concluded that the haze aerosol particles of marine origin in California were most likely represented by an elastic organic coating over a volatile core. As the volatile interior evaporated in the electron microscope, a wrinkled coating remained which was likely made up of macromolecular surfactants. If two particles with such a coating coagulated, the lower surface tension resulting from the surfactants would allow it to divide again into two particles to minimize surface free energy. The size of the two particles would be exactly those needed to accommodate all surfactant molecules as Langmuir films on the surface, with no surplus or deficiency [Tervahattu et al., 2002]. Recent studies suggested that organic secondary particles may become associated with secondary inorganic materials [Murphy and Thomson, 1997a,b; Noble and Prather, 1996]. Russell et al. [2002] characterized marine boundary layer particles using soft X-ray spectromicroscopy at atmospheric pressure. This technique allows the mapping of organic functional groups in individual particles with diameters between 0.2 and 10 µm. The measurements at atmospheric pressure retain semi-volatile compounds, including the species forming organic films, which are typically lost in the high vacuum of electron microscopy. The maps showed clear boundaries between ketone-like compounds and carboxylic acids in the particles, the latter preferentially occurred at the particle surface, possibly due to their more surface-active character. The state of mixing of aerosol constituents was inferred in in-situ measurement of individual aerosol particles at altitudes of 5–19 km by laser ionization time-of-flight mass spectrometry [Murphy et al., 1998]. Whereas the negative ion spectra of stratospheric aerosol particles were simple, with sulfate being the predominant component, those of the upper tropospheric aerosol contained many organic peaks as well, indicating internal mixing between the main aerosol components. Although quantitative information could not be derived from the mass spectra observed, it was clear that the tropopausa was associated with a distinct change in aerosol chemical composition.

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Individual marine aerosol particles greater than 0.16 µm were observed by laser mass spectrometry [Middlebrook et al., 1998]. During clean marine conditions, as much as 53 % of the particles which had detectable chlorine signal also had detectable organic peaks. This ratio was even higher (62 %) under polluted conditions. The organic compounds were always found internally mixed with sea salt particles. They also found that the relative contribution of sodium sulfate above a certain threshold was larger in particles showing the presence of organic peaks. This fact, together with the great abundance of internally mixed organic/sea salt particles, indicated similar mechanism for the incorporation of both species upon aging. Below that threshold, the organic content tended to positively correlate with iodine and phosphate. The presence of iodine in organic particles was consistent with the source of the organic material being the iodine-enriched surface film on seawater. Similarly, the co-existence of phosphate with organic matter suggested the presence of low-volatility organophosphates which are known surfactants. Although the method did not allow quantitative determination of the aerosol species, sensitivity considerations would suggest that organic matter may make up about 10 % of the dry marine aerosol mass. Tervahattu et al. [2002] observed irregularly shaped transparent particles under the scanning electron microscope which changed shape continuously under the electron beam, enlarging in one side and shrinking in another. EDX analyses showed that the main elements in these particles were carbon, oxygen, sulfur and sodium. They assumed that they observed aged sea salt aerosol particles, covered with a lipid membrane, and gaseous sulfur oxides or water vapor expanding inside the lipid membrane when heated by the electron beam. This assumption was in harmony with the inverted micelle model of marine aerosol, where the hydrophilic part of the lipid molecule is bound to the polar core, with the hydrophobic chain facing outward. Upon aging, the hydrophobic film oxidizes by OH radicals and hydrophilic sites with alcohol, aldehyde, ketone, and carboxylic functional groups are formed [Ellison et al., 1999]. 4.1.6 Bulk Chemical Characterization of OC 4.1.6.1 Solubility Classification The first organic speciation studies in atmospheric aerosol in the 1950s were targeted to the potentially carcinogenic polycyclic aromatic hydrocarbons in urban particulates [Cholak et al., 1955]. Hence the amount of total organic carbon was not particularly interesting, and it is not surprising that use of nonpolar solvents, such as cyclohexane,

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dichloromethane or benzene were preferred. Nonetheless, even some of the earliest studies suggested that significant amount of water-soluble organic compounds could form upon irradiating automobile exhaust [Renzetti and Doyle, 1959]. It was already in the early 70s when total carbon was first measured in atmospheric aerosol [Mueller et al., 1972], and was classified into broad solubility classes, including the water-soluble fraction [Grosjean, 1975]. Systematic solubility studies of particulate carbon started in the early 1980s, into broad polarity classes with different organic solvents [Daisey et al., 1982; Gundel and Novakov, 1984], or quantification of the water-soluble fraction of organic carbon [Cadle et al., 1982; Mueller et al., 1982]. The concept that organic material could be associated with either hygroscopic or hydrophobic particles is consistent with the enormous diversity of their chemistry. Polar and nonpolar solvents tend to extract distinct, but intersecting subsets of organic compounds from atmospheric aerosol. Quantitatively polar solvents (water, methanol, or acetone) were more efficient, on average by 5–33 %, than nonpolar solvents such as benzene [Grosjean, 1975]. It also means that significant amounts of water-soluble species cannot be recovered in nonpolar solvents. The most effective extraction media were mixtures of polar and nonpolar solvents which may extract as much as 75 % of total carbon [Cadle et al., 1982]. The combination of solvents of different polarity can be so effective that it can also remove a measurable fraction of BC, thus reducing attenuation through the filter [Gundel and Novakov, 1984]. Nonpolar solvents, such as dichloromethane, generally extract significantly more organic carbon from urban than from rural aerosol (e.g. 41–44 % and 26 %, respectively) [Alves et al., 2002]. This is not surprising since in urban aerosol there are an abundance of neutral organic compounds, whose elutable fraction carries the distinctive imprint of vehicular exhaust. Besides, there are partially oxidized products, mainly acidic compounds, whose share increases markedly when moving away from sources. The major components of the eluted extract were aliphatic and acidic compounds, and alcohols, with minor concentrations of PAHs, aldehydes, and ketones. The extraction of the backup filter in cyclohexane removed about 70 % of the total carbon, a small fraction of which (∼5 %) was recovered in the extract after evaporation to dryness [Gundel and Novakov, 1984]. Approximately one-fifths to two-third of total carbon was found to be soluble in water. Mueller et al. [1982] found most WSOC/TC ratios to vary between 0.35 and 0.6, with average values close to 0.5. Jacobson et al. [2000] suggested that water extraction should be performed twice, once with large excess of water and second with the amount typically available in clouds. The two methods would likely provide different information. Novakov et al. [1997a] found that more than 70 % of organic carbon mass was water soluble in marine aerosol samples collected in the trade winds at

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Puerto Rico. It should be noted that the authors applied large excess of water for extraction, and the number above also includes adsorbed gaseous phase species which were retained on the front filter. In rural aerosol WSOC were found to make up 20–70 % of the total fine aerosol carbon [Temesi et al., 2003]. If rural PM10 aerosol was extracted successively with water after extraction with a nonpolar solvent such as dichloromethane, the extracted organic carbon increased from 26–38 % to 41–70 % [Carvalho et al., 2003]. At a forest site, the same combination recovered 31–78 % of OC, though the dichloromethane-extracted fraction was highly variable (12–52 %). These results—which also included part of the coarse fraction of aerosol—pointed out the importance of water-solubility among the organic species of the aerosol. Estimations for the concentrations of primary and secondary organic aerosol can be derived from the measured concentrations of non-polar and polar carbon fractions, respectively [Appel et al., 1976]. Aged aerosol particles collected in semi-urban locations had roughly three times more polar than nonpolar carbon, with at least one-third of the TC was insoluble. In contrast, the polar-to-nonpolar carbon ratio in samples collected at close proximity of sources was ∼1:1. In vehicular exhaust, it dropped to 0.6:1 [Gundel and Novakov, 1984]. Since a significant fraction of polar compounds were associated with combustion source samples as well, they may not be entirely of secondary origin, as it was previously supposed. Sequential n-hexane and 2:1 benzene/isopropanol extractions were applied to characterize fine and total particles in summer in the vicinity of the Grand Canyon National Park [Mazurek et al., 1997]. In this study the solvent-soluble compounds that can be quantified by GC-FID were referred to as elutable organics, operationally divided into acidic and neutral fractions. About half of the elutable fraction was composed of highly polar organic compounds, which would have been eluded analysis without prior derivatization. Their carbon number distribution was approximated by the elution characteristics of n-alkane homologues. For example, the range allocated by the retention times of the n-C15 and n-C20 alkanes was expected to comprise monoterpenoid, sesquiterpenoid and some diterpenoid species. A typical distribution of acidic and neutral elutable organics is shown in Figure 4.10.

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Figure 4.10. Distributions and ambient mass concentrations of acid and neutral elutable organics obtained from GC-FID analyses of Hopi Point fine particles, August 1989 monthly composites. Each bar shows the quantity of all organics that elute between the elution points of the Cn and Cn+1 n-alkanes, including the Cn alkane itself, where the alkane carbon number (n) is stated below each bar (after Mazurek et al. [1997]).

This distribution can be compared to that obtained in Los Angeles and at a receptor site in Rubidoux 60 miles inland, showing substantial compositional differences and suggesting disparate origins. Based on an FTIR spectrocopic study combined with solvent rinses on marine aerosol, a clear distinction was established between the solubilities of organic aerosol in the boundary layer and the free troposphere [Maria et al., 2002]. Below about 1800 m, an average of 35 % of OC was hydrophobic, whereas above that most samples had no hydrophobic fraction, and 60–90 % of OC was actually water-soluble. In the marine boundary layer, the observed solubility behavior was consistent with a model mixture of 30 % nnonadecane, 20 % citric acid, and 50 % oxalic acid, by mass. On the contrary, in the free troposphere the corresponding composition was 0–10 % n-nonadecane, 40 % adipic acid, and 50–60 % oxalic acid. 4.1.6.2 Bulk Chemical Characterization of Organic Aerosol The very first information on the bulk carbonaceous aerosol was actually a by-product of the analytical determination of aerosol carbon by evolved gas analysis [Novakov, 1981]. In this technique, the aerosol filter sample is heated at a predetermined rate in an oxidizing atmosphere. The evolved gases are flushed through an oxidizing catalyst (MnO2 or Cu2O) held at high temperature to ensure complete oxidation to CO2, which is then detected by a non-dispersive infrared detector. By recording the detector signal as a function of temperature, a thermogram can be obtained. After

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careful calibration the integrated area under the thermogram gives an accurate measure of total carbon on one hand, and the features of the thermograms shed some light on the thermal stability, volatilization and oxidation properties of the bulk organic species as well as black carbon. The very first application of this technique for atmospheric aerosol collected at an urban and a receptor site in the Los Angeles Air Basin during the period of intense photochemical smog revealed distinctive features between the thermograms from the two sites, and even substantial diurnal variations for the receptor site [Ellis and Novakov, 1982]. The principal features of the thermograms were the appearance of poorly resolved peaks (marked with α, β, γ, δ), corresponding to volatilization and/or oxidation of carbon species of increasing thermal/oxidation stability, respectively [Ellis et al., 1984]. The peak α observed between 100 and 230 °C was attributed to the volatilization of low molecular weight organics, and peak δ above 450 °C to BC. After normalization to the source thermograms the excess peaks β and γ—which were more prominent in the samples from the receptor site—were suggested to be a first-order measure of secondary organic carbon. These secondary species—solely on the basis of their thermal/optical properties—were hypothesized to be high molecular weight polymeric material. This was a brilliant idea at that time—and far too premature to gain widespread recognition. It took another 15 years for the first field observations to emerge and prove that this is actually the case [Havers et al., 1998; Zappoli et al., 1999]. Nowadays we know a lot more about the nature and properties of these high-molecular weight substances. This basic method was refined in subsequent studies by taking into account sampling artifacts and the varying combustion temperatures of BC, and led to the development of an integrated approach for the determination of OC/BC, which is discussed in details in sub-chapter 2.2. Typical carbon thermograms of suburban, rural and trade wind aerosol are shown in Figure 4.11 [Novakov et al., 1997a]. Care has to be taken when interpreting these thermograms since substantial amount of carbon was recovered on the backup quartz filter (for details, see Figure 2.1 and related discussion in sub-chapter 2.1). The thermogram of a marine submicron aerosol samples (d <0.6 µm) were confined to temperature below about 300 °C and showed no discernible peaks of high-temperature peaks indicative of BC, contrary to those of rural or suburban aerosol samples. The virtual absence of BC particles indicated that oceanic background organic aerosol mainly resulted from sources other than combustion. The observations regarding the substantial concentrations

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of polar gaseous species suggested that some of them might be precursors of organic aerosol.

Figure 4.11. Typical carbon thermograms of (a) suburban aerosol sample (collected June 3–4, 1995, in Berkeley, California, aerodynamic diameter (Dp) <2.5 µm, (b) rural aerosol particles (collected October 22–23, 1993, at Point Reyes, California, (Dp) <2.5 µm, during a stagnant air period), and (c) trade wind aerosol sample (collected February 15–16, 1995, at Cape San Juan, Puerto Rico, (Dp) <0.6 µm (after Novakov et al. [1997a]).

Combination of this thermal technique with water extraction revealed additional features of the bulk organic matter. For example, in trade wind

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fine aerosol the amount of carbon was reduced to less than one third after water extraction, indicating the water-soluble nature of aerosol carbon [Novakov et al., 1997a]. Somewhat less, about half of the carbon was found to be water-soluble in rural fine aerosol collected in summer, and a similarly large fraction of the refractory aerosol component appeared to be soluble in water [Gelencsér et al., 2000a]. The thermal properties of this refractory carbon differed markedly from those of the coarse aerosol, but seemed to coincide with those of a reference humic acid on pre-baked quartz filters. Fourier Transform Infrared Spectroscopy (FTIR) of aerosol samples may be used to quantify organic matter by functional groups, thus providing a compromise between bulk OC measurements (such as evolved gas analysis, EGA or thermo-optical analysis) and speciation by gas chromatography-mass spectrometry [Russell, 2003]. FTIR analysis of sizeresolved aerosol collected in the Smoky Mountain in summer was coupled with successive solvent rinses to determine the polarity of the organic component of the aerosol [Blando et al., 1998]. The FTIR spectrum of the aerosol fraction in the size range of 0.5–1.0 µm is shown in Figure 4.12, together with the spectra acquired after rinsing the sample successively with hexane, acetone, and water.

Figure 4.12. FTIR spectra of 0.5–1.0 µm diameter particles collected on August 21–22, 1995, at Look Rock, Smoky Mountains, TN. Shown are the original spectrum (a), the spectrum after removing nonpolar organics with hexane (b), the spectrum after subsequent removal of polar organics in acetone (c), and the spectrum after subsequent removal of inorganic salts in water (d). Several absorbances are labeled with functional group assignments and with absorbance bands in wavenumbers (after Blando et al. [1998]).

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The characteristic absorption bands are marked on the spectra. While hexane caused little change in the spectrum, rinsing in acetone removed several large peaks, indicating that a large fraction of the aerosol was polar. In particular, the carbonyl (at 1730 cm−1) and hydroxyl (at 1630, 3450 cm−1) bands were effectively removed, together with peaks at approximately 875, 1045, 1225 and 2490 cm−1. These peaks were attributed to organosulfur compounds, e.g. alkyl sulfonic acids and sulfonates which are formed in the aqueous phase from SO2-aldehyde reactions [Olson and Hoffmann, 1989]. As a likely candidate, camphorsulfonic acid was postulated which can be formed in the reaction of biogenic camphor with aqueous acidic sulfate. The predominance of carbonyl and organosulfur absorbances underscored the significance of secondary formation mechanisms. The bulk characterization of organic carbon, in particular of its watersoluble fraction, revealed that almost all observed properties resembled closely those of natural humic substances. The first observations of this kind were related to urban particulate matter, the techniques used were UV-VIS spectrophotometry and proton nuclear magnetic resonance spectrometry (HNMR) [Havers et al., 1998]. These authors were the first to introduce the term “humic-like substances”, HULIS, which has become widely used in the literature, being the most popular among all the synonyms that have ever been suggested. At about the same time, there appeared a comprehensive study on the bulk properties of water-soluble organic matter in aerosol from a polluted, rural and marine environment [Zappoli et al., 1999]. Its major finding was that WSOC made up a significant fraction of fine aerosol carbon, and its detailed analytical characterization revealed a stunning resemblance to a reference humic acid. In spite of the high degree of similarity, the authors meticulously refrained from using the term “humic-like” for this major class of compounds. Instead, they termed this fraction “macromolecular”, though none of their analytical methods yielded direct evidence that these species were indeed of high molecular weight. Their hesitation was understandable since the presence of such compounds in fine aerosol, which predominantly form in secondary processes, was baffling to the scientific community. Terrestrial humic and fulvic acids were shown to be important constituents of aeolian particulates, which predominantly occurred in the coarse size range [Simoneit, 1977]. Their formation mechanism was well-understood, and for decades, these substances were out of the focus of aerosol research. Albeit baffled by their own findings, Zappoli et al. [1999] felt compelled to come out with a plausible explanation for the origin of these “alien” species. They suggested biomass burning to be the most likely source, and postulated direct condensation of high-molecular weight burning products as a possible formation mechanism. Their hypothesis assigned

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primary anthropogenic origin to this compound class, and seemed plausible to account for its observed abundance in the fine particle size range. It should be noted that in biomass burning aerosol, there is no clear boundary between aerosol components of primary and secondary origin, since equilibrium changes rapidly in a cooling plume. However, the general term “secondary organic aerosol” is tacitly restricted to cases when gasparticle transition is preceded (or followed) by chemical transformation. In the light of this concept, direct condensation of high molecular weight combustion products is predominantly a primary (physical) process.3 These pioneering works stimulated further studies on the occurrence and properties of HULIS in atmospheric aerosol. New analytical techniques were called in to exploit the properties of the bulk WSOC, and compare them to those of natural humic matter. It was shown that the electrochemical properties and metal-complexing ability of polluted fog water mirrored those of a reference humic acid [Gelencsér et al., 2000a]. Other studies applied various separation methods to characterize such compounds, primarily on the basis of their acid-base properties. These efforts made it clear that a traditional separation into individual organic species is not possible: the outcome was at best a few poorly resolved, broad peaks, and/or an unresolved “hump” on the chromatogram or electropherogram. Capillary zone electrophoresis of polluted fog water and aqueous extract of rural fine aerosol showed a broad distribution of electrophoretic mobilities, which suggested a similar trend in the charge-to-size ratios [Krivácsy et al., 2001a]. The observed pH-dependence of the electrophoretic behavior implied that most acidic groups had dissociation constant inferior to 10−5. In other terms, these polycarboxylic acids were found to be weaker acids than acetic acid. The water-soluble organic fraction of the fine aerosol collected at a polluted site was separated into the generic classes of neutral/basic compounds, mono- and dicarboxylic acids, and polycarboxylic acids (with at least 3 negative charges per molecule), by preparative ion-exchange chromatography [Decesari et al., 2000]. The chemical structure of these broad compound classes were revealed by 1HNMR spectrometry. The spectra of the neutral/basic compounds suggested the presence of mainly hydroxylated/alkoxylated aliphatic species, with indications for the presence of polyols. Mono- and dicarboxylic acids appeared to be predominantly aliphatic carboxylic acids and hydroxy carboxylic acids, whereas 3

The formation mechanism is not explicitly stated in the original work. I take the courage to classify it as primary because classification of available hypotheses so requires; furthermore, I myself was one of the co-authors of that paper.

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polycarboxylic acids had a more pronounced unsaturated character, with an aromatic core bearing aliphatic chains with −COOH, −CH2OH, −COCH3 or −CH3 terminal groups. The observed features of the polyacidic compounds closely resembled those of terrestrial or aquatic humic matter. In general, the polycarboxylic acids were the most abundant class of WSOC throughout the year, except in summer, when mono- and diacids were predominant [Decesari et al., 2001]. It is of interest to modeling of atmospheric radiative transfer that a major portion of the WSOC was UVabsorbing, and the specific UV-absorptivity was highest for the polycarboxylic acids. An alternative preparative-scale separation method was developed specifically for the isolation of HULIS from aqueous extracts of aerosol [Varga et al., 2001]. The method was carefully optimized to isolate the fraction of WSOC that retained the key spectral properties of reference humic and fulvic acids. The separation is based on molecular interactions with the non-dissociated species, and is therefore conceptually different from the method based on the separation of ions. However, polycarboxylic acids can be separated both in their ionic and molecular forms. Therefore both methods are believed to target broadly the same generic class of compounds. It also implies that the terms “polycarboxylic acids” and “HULIS” possibly refer largely to the same fraction of organic aerosol. Since, however, there has been no intercomparison between these methods, this statement is merely based on an assumption derived from the observed chemical properties of the isolated compounds, and the fundamental principles of the separation. There is, however, an important difference between the two methods. The method by Varga et al. [2001] allows the isolation of HULIS from inorganic compounds too, which is not possible in method by Decesari et al. [2001]. This fact facilitates analytical determinations which otherwise would not be feasible in the presence of interfering inorganic species. For example, the elemental composition of HULIS isolated from rural fine aerosol was found to be remarkably constant throughout the year, corresponding to an average molar ratio of C:H:N:O of 24:34:1:14 [Kiss et al., 2002]. This finding may imply that either the carbonaceous particles share common origin (most likely combustion sources), or that the processes in the atmosphere producing or modifying the particulate organic matter result in a similar net composition throughout the year4. Although the former approach has long been thought to be the case, nowadays there has been mounting evidence that the latter hypothesis is more tenable. An important step towards the understanding of the origin of watersoluble organic compounds in rural aerosol was the determination of their molecular weight distribution by various techniques, including ultrafiltration, liquid chromatography-atmospheric pressure ionization mass spectrometry, 4

This statement was first put forward by Ketseridis in 1976.

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and vapor pressure osmometry [Kiss et al., 2003]. The most interesting finding of this study was that virtually all WSOC passed through an ultrafiltration membrane having a 500 Da nominal molecular weight cut-off. The isolated HULIS—which made up of more than half of WSOC by mass—was further characterized to determine their ion mass distribution. Since the ionization technique was optimized to minimize the occurrence of multiple charges, it can be regarded as a molecular weight distribution of HULIS. A continuous molecular weight distribution was found between about 100 and 500 Dalton, with maxima in the range of 200–300 Da. These conclusions were confirmed by vapor pressure osmometry, which provided direct estimates for the average MW of HULIS as shown in Figure 4.13.

Figure 4.13. Corrected MW estimates (Mn) of isolated HULIS as obtained by vapor pressure osmometry (after Kiss et al. [2003]).

The average molecular weight was markedly lower than those found for reference aquatic humic and fulvic acids under the same conditions. These observations are among the very few that pointed to important differences between HULIS and natural humic substances, and implied distinct mechanisms of their formation. In fact, very few analytical techniques are capable of providing chemical information directly on the carbonaceous component of the bulk aerosol collected on filter substrates or impactor plates. For the determination of elemental concentrations, quite a few techniques (e.g. nuclear methods) are available which do not require any sample preparation step. On the other hand, for organic aerosol the only known method is pyrolysis-gas

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chromatography-mass spectrometry which yields information on organic chemical structures directly from aerosol particles. Taking into account the fact that most organic aerosol species contain functional groups (e.g. carboxylates) which yield non-specific thermal decomposition products (e.g. carbon dioxide) upon conventional analytical pyrolysis, in its very first application in aerosol chemistry a derivatization technique was introduced [Gelencsér et al., 2000b]. The thermally assisted hydrolysis-methylation allowed labile functional groups to be converted into their respective esters, thus preventing decarboxylation upon pyrolysis and yielding more specific pyrolysis products.5 The analysis of rural fine aerosol by this method revealed overall structural similarities to those of natural humic substances. The predominant pyrolysis degradation products both in the aerosol and terrestrial humic acids were n-alkanoic acids, α,ω-dicarboxylic acids (in the carbon number range of C4–C9), and benzenedicarboxylic acids. The apparent structural similarity to terrestrial humic substances made the authors suggest the term “atmospheric humic matter” in place of HULIS. The rationale behind this suggestion was that HULIS in aerosol were thought to be chemically indistinguishable from the wide variety of natural humic substances present in other reservoirs. To what extent this idea can be justified, it is currently not known. There have been no systematic studies aimed at classifying HULIS into the broad class of natural humic substances. However, the suggested nomenclature did not gain recognition at all, and even its proponents returned to using HULIS in their subsequent works. The major objection to using the word “humic” without the restricting clause was probably the lack of any feasible relation between HULIS and natural humic substances. Using the same method similar compounds were found in urban aerosol, except that there higher substituted lignin pyrolysis products were also observed [Subbalakshmi et al., 2000]. The pyrograms of biomass burning aerosol from Brazil, however, revealed some differences with respect to those of rural fine aerosol [Blazsó et al., 2003]. Most importantly, in biomass burning aerosol there were several higher substituted aromatic compounds which were absent from rural aerosol. These species—which are typical lignin degradation products—were also shown to be present in the pyrogram of soil humic and fulvic acids [Martín et al., 1994, 1995]. This finding made Gelencsér et al. reconsider their previous results on HULIS in rural fine aerosol [Gelencsér et al., 2002]. Their conclusion was that in spite of all apparent similarities, these differences unambiguously prove the disparate origin of HULIS, namely their atmospheric formation in

5

This can be done by applying a few hundred µl of the solution 10 % tetramethyl-ammonium hydroxide in water onto a filter spot then allowing it to dry at room temperature.

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heterogeneous or multiphase processes. Their hypothesis on HULIS formation is detailed in sub-chapter 3.3. 4.1.7 Bulk OC/BC Source Apportionment 4.1.7.1 Radiocarbon (14C) Measurements The 14C isotope (or rather its ratio to 12C), which is produced naturally in the atmosphere, is a geophysical clock widely used in many fields of geosciences. In the field of aerosol chemistry, however, the 14C/12C ratio is not used for dating, but for fundamental apportionment of modern and fossil fuel carbon. Apart from the experimental difficulties and the related uncertainties, this simple division alone cannot provide apportionment of traditional biogenic and anthropogenic sources. In other terms, the presence of modern carbon does not necessarily imply biogenic contribution, since it can equally result from biomass burning or cooking operations. The magnitude of these effects must be estimated independently using molecular tracer measurements. Instrumental developments have only recently made it possible to determine 14C in microgram quantities of aerosol. Formerly radiocarbon determination required grams of carbon, which is now reduced to about 100 µg—an amount which is within the scope of high volume aerosol sampling even in the most pristine environment. Nevertheless the sample preparation involves many steps and may be prone to severe artifacts. The preparation of filter-collected carbonaceous material for 14C accelerator mass spectrometry involves the combustion of the filters at high temperature (e.g. 900 °C for 1 hour) in sealed pre-cleaned quartz tubes containing CuO and Ag wire. All carbon is converted to CO2, which is then reduced over granular Zn to carbon-monoxide at 400 °C and further reduced over Fe powder to graphitic carbon at 600 °C. The 14C results are reported as percentage of modern carbon (pMC) expressed as 14

pMC ( PM 2.5 ) =

C 14

13

0.95 * C

C PM 2.5 13

C SRM 4990b

* 100

(4.1)

where SRM4990b is oxalic acid for radiocarbon dating, and the denominator is defined as modern carbon. When applied to aerosol measurements this simple concept is complicated by various factors. The most important is that the 14C content of

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the biomass varies as a consequence of the combined effects of three factors: a) solar modulation of cosmic rays, b) dilution of atmospheric radiocarbon with fossil fuel generated CO2 (Suess effect), and c) artificial release of radiocarbon (primarily from nuclear tests). Among these, the perturbation caused by nuclear tests predominates, causing about 30 % excess in radiocarbon of living vegetation [Currie, 1982]. For example, it may happen that pMC can be well in excess of 100 % (e.g. 127 %) when logs about 50 years old are burned, which had grown at the times when additional 14C was injected into the atmosphere from nuclear tests. In contrast, the pMC for biogenic emission is on average 107. Fossil fuels do not contain modern carbon, consequently their pMC value is essentially 0. 14 C measurements were made on fine particles collected at an urban site more than two decades ago [Dzubay et al., 1982]. The mean value of the estimated combined contribution of modern carbon was found to be 40 ± 14 %. When PM2.5 aerosol samples pMC values are depicted as a function of EC/TC ratio, a traditional indicator of primary fossil fuel burning emissions, a pronounced anticorrelation can be observed between the two variables, as shown in Figure 4.14 [Lemire et al., 2002].

Figure 4.14. Percentage modern carbon, based on the total carbon mass, versus EC/TC ratio. Error bars represent the standard uncertainty (u) in the EC/TC ratio, and the uncertainty in pMC is less than the symbol size (after Lemire et al. [2002]).

As it can be seen, a substantial fraction of the aerosol carbon at an urban/suburban site and a forested rural site in Southeast Texas was composed of modern carbon. The contribution of modern carbon varied between 27–73 % and 44–77 % at the two sites, respectively. Gaffney et al. [1984] measured radiocarbon in aerosol samples collected in Alaska and found that more than 30 % of TC was biogenic by

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origin. This could be from either biogenic secondary aerosol or biomass burning. The 13C isotopic data were consistent with predominantly C3-type plant sources. They concluded that EC in the Artic haze was predominantly from fossil fuel combustion, as it was also indicated by the high correlation between EC and aerosol sulfur. 4.1.7.2 Stable Carbon Isotope Measurements The isotopic composition of atmospheric CO2 is primarily determined by the atmosphere-ocean exchange, but is also influenced by photosynthesis and human activities. Due to equilibrium isotope fractionation, atmospheric CO2 is depleted in 13C as compared to marine carbonates. The present mean δ13C of atmospheric CO2 is about −7.8 ‰ relative to the PDB standard. This depletion is further enhanced by kinetic isotope fractionation (preferential 12 C uptake) during photosynthesis. Plants using the Calvin-Benson cycle (C3 plants), have δ13C-values in the range of −23 ... −30 ‰ (vs. Pee Dee belemnite). Virtually all trees, most shrubs and mid-latitude and boreal grasses and sedges belong to the class of C3 plants. The other major type is the C4 plants which utilize the Hatch-Slack (or dicarboxylic acid) cycle, and have δ13C-values in the range of −10 ... −16 ‰. These plants comprise warm-season grasses and sedges, thus they can be found predominantly in tropical savannas, temperate grasslands and semideserts [Cerling et al., 1993].6 The marine biosphere exhibits a geographical variability of its isotopic composition: for temperate and low latitude regions, the δ13C range is −22 ‰ to −18 ‰, except for upwelling areas where planktonic tissues can be more enriched in 13C (−20 ‰ to −14 ‰). Fossil fuels, such as petroleum, have δ13C-values close to the mean of C3 plants (26 ± 2 ‰) [Sackett, 1989]. As a result of the differences in δ13C, stable carbon isotope analysis— either compound-specific or in bulk—can provide information on the origin of the vegetation-derived organic aerosol fraction. The sources of this aerosol fraction, however, could be manifold. It may be emitted directly or formed in atmospheric reactions of VOC emitted by biomass burning. Alternatively, plants can directly emit primary particles and VOC which are then converted to biogenic SOA by atmospheric processes. Therefore on the basis of stable carbon isotopic composition apportionment of biogenic and anthropogenic

6

In fact, there is a third type, the CAM (Crassulacean Acid Metabolism) plants which use both carbon fixation pathways so their δ13C-values are intermediate. These include many succulents such as cacti, which contribute very little to the biomass in any habitat.

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sources is not possible. Nevertheless over the oceans distinction between marine and land aerosol sources can be made with caution. Combustion processes seem to slightly modify stable carbon isotopic composition. Combustion experiments showed that measurable isotopic fractionation did occur during combustion of solid fuels, which limits any simplistic application of δ13C to within-class (e.g. C3) apportionment [Currie et al., 1999]. Burning of C3 plants under controlled conditions was found to shift δ13C by +0.5 ‰ compared to source vegetation, whereas burning of C4 plants caused an opposite change by −3.5 ‰ [Turekian et al., 1998]. Fortunately the isotopic signature of the aerosol appears not to be modified upon atmospheric transport [Cachier et al., 1986]. The organic carbon concentration and isotopic composition in sizesegregated atmospheric aerosol particles collected at remote marine locations were determined [Chesselet et al., 1981]. The total OC concentrations were found to be in the range of 0.3 to 1.2 µg m−3, with 80 % present in the fine particles (d <1 µm). The δ13C value for fine particles was −26 ‰ ± 2 ‰, strongly supporting continental origin, most likely through atmospheric formation from gaseous precursors originating from the continent. The OC in the coarse fraction, on the other hand, appeared to have direct marine origin (δ13C = −21 ± 2 ‰), and implied an enrichment of organic carbon by a factor of 100 over bulk seawater. In general, oceanic aerosol samples in the Northern Hemisphere typically exhibited very low δ13C values (−26 ‰) suggesting the predominance of combustion emission sources, whereas those in the Southern Hemisphere had much higher δ13C values (−23 ‰ to −17 ‰), indicating either marine or continental natural sources [Cachier et al., 1986].

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4.2 Major Compound Classes and Atmospheric Tracer Compounds Organic speciation studies targeted at potentially toxic, carcinogenic or mutagenic organic aerosol components or important source tracers have a much longer history and are incomparably more numerous than studies on bulk organic aerosol. From the viewpoint of this book, however, these topics have very little—if any—relevance provided that we disregard global issues such as the effect aerosol-borne organic pollutants on remote ecosystems. Since pollution itself is a global issue which is a subject of many excellent books and reviews in the field of environmental (analytical) chemistry, I decided not to bring it into the focus of this book. To be honest, the neglect of this aspect helps maintain the integrity of my approach, which would have otherwise been seriously damaged by the imbalance between the studies on organic speciation and bulk characterization of organic aerosol. Therefore in this sub-chapter only those speciation studies are presented which have some relevance in source apportionment of bulk organic aerosol. Given the vast number of publications even with this restriction in mind, this sub-chapter will be much more illustrative rather than comprehensive. It will only give a flavor of what has been achieved in this field, and how these studies can assist in source apportionment of organic aerosol, sometimes in combination with other approaches. 4.2.1 Major Classes of Organic Tracers 4.2.1.1 Lipid Fraction of Organic Aerosol a) Alkanes The n-alkanes—though being minor components of organic aerosol— are probably one of the most frequently characterized compound classes in atmospheric aerosol. Due to their apolar character and volatility they are readily amenable to gas chromatographic analysis. In addition, identification of the most abundant n-alkanes is straightforward since they are present in the form of homologue series. The sources of n-alkanes are manifold. They can be directly released by vegetation or as part of soil detritus, volatilized by burning processes, or may result from the sea surface microlayer by bubble bursting. Their common feature is only that all of these sources release n-alkanes as primary

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product, i.e. these compounds cannot form in atmospheric reactions. This does not mean, however, that n-alkanes are released solely in aerosol particles. Since most of their representatives are semi-volatile, compounds released in the gas phase may also partition into the aerosol phase, and vice versa, compounds released in particulate form can equilibrate into the gas phase. This also means that the sampling of aerosol for n-alkane analysis is prone to substantial sampling artifacts, either positive or negative. Due to their high volatility, homologues below about n-C20 cannot be quantified accurately in the aerosol phase. N-alkanes are typically manifested in atmospheric aerosol in the form of a distinctive homologue series, ranging from about n-C15 to n-C33. This range is constrained by the high volatility of the compounds from the lower end, and by the limited capability of gas chromatography to separate lowvolatility species from the higher end. It should be noted that recent development of high-temperature gas chromatography has allowed this upper limit to be extended considerably. Long-chain n-alkanes are abundant constituents of lipids found in terrestrial higher plants epicuticular waxes [Eglinton and Hamilton, 1967]. These biosynthetic n-alkanes (>n-C26) can be formed by elongation of the most abundant C16 and C18 fatty acids in plant waxes, followed by decarboxylation. These waxes form a protective coating on leaves and stems in order to sustain the water balance in the plants under dry conditions. They are frequently present as microcrystals or hollow tubes with dimensions of about a micrometer or below [Hall and Donaldson, 1963]. Parts of these wax protrusions may get easily sloughed off by the effect of the wind and become airborne. Alternatively, decaying soil organic matter can be lifted during dust storms. In these primary particles long-chain n-alkanes are always accompanied by homologous series of fatty acids and n-alcohols. However, n-alkanes are the most refractory compounds among all lipid constituents therefore they are probably the best biomarkers for terrestrial vegetation waxes. In biomass burning the wax constituents are volatilized directly from the vegetation burned. These processes can release n-alkanes with carbon numbers in excess of 35, but these low-volatility homologues can only be determined by high-temperature gas chromatography-mass spectrometry [Elias et al., 1999]. A distinctive feature of wax-derived n-alkanes is the marked predominance of odd carbon-numbered homologues. This can be expressed as carbon preference index (CPI), which is given as [Kolattukudy, 1976]

CPI = 0.5 *

n

n −1

n

n +1

i

i −1

i

i +1

∑ ( X even ) / ∑ ( X odd ) + 0.5 * ∑ ( X even ) / ∑ ( X odd )

(4.2)

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where Xi denotes the mass concentration of homologue i. The CPI of nalkanes in vegetation waxes is high (>5) [Simoneit and Mazurek, 1982]. The high CPI values are preserved in the solvent-soluble lipids of aeolian particulates, indicating leaf-wax origin [Simoneit, 1977]. The highest concentration homologue is typically found at n-C29, n-C31, or n-C27. It should be noted that the intermediate high-molecular weight alkanes with even number of carbon atoms, C28H58, C30H62, and C32H66 may also originate from anthropogenic sources [Veltkamp et al., 1996]. The mass-weighted average chain length of n-alkanes may indicate taxonomic differences in vegetation, though it is also likely influenced by conditions such as aridity or growing season temperature [Alves et al., 2001; Gagosian and Peltzer, 1986]. Petroleum derived n-alkanes, on the other hand, have virtually no odd carbon number predominance (CPI∼1) above n-C23 [Simoneit, 1984]. This fact allows the estimation of the contribution of the wax input to n-alkane concentrations in atmospheric aerosol. This can be simply done by subtracting the average concentrations of the preceding and subsequent even carbon number homologues (assigning zero to the negative values) from those of the odd carbon number ones. N-alkanes without odd carbon number preference may also have marine sources, most likely bubble bursting from the sea surface microlayer [Marty and Saliot, 1982]. Homologues from this source, however, typically fall in the range of n-C20 to n-C25, with minor concentrations between n-C15 and n-C19. The different sources of n-alkanes as well the different volatilities of the homologues result in differences in the size distribution of these compounds in atmospheric aerosol. Typical mass median aerodynamic diameters of n-alkanes were 0.45 µm for urban aerosol, whereas the values for background marine and forest aerosol were 2.00 and 0.63 µm, respectively [Kavouras and Stephanou, 2002]. In the western equatorial Atlantic, under very stable meteorological conditions, about 60 % of the total n-alkane concentration was found to be associated with fine particles (<1 µm) [Marty and Saliot, 1982]. It should be noted, however, that failure to correct for adsorption artifacts during filter sampling may result in a severe bias in the mass size distribution of n-alkanes towards the fine particles. In the gas chromatogram of organic aerosol extracts containing the peaks of n-alkanes sometimes an envelope of unresolved complex hydrocarbons (branched and cyclic) appears which is called “hump”. It may either be a narrow one centered between C19–C20, as typically present in chromatograms of degraded algal detritus or of other microorganisms [Cox et al., 1982], or a broad hump centered around C27 which is often regarded as

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an approximate measure of the level of contamination from petroleum residues [Simoneit and Mazurek, 1982]. b) N-alka(e)noic acids (n-fatty acids) The n-alkanoic acids, also known as n-fatty acids typically range from 3). The distribution of the homologues is usually bimodal, with maxima at n-C16 (palmitic acid) and either at n-C22 or n-C24 [Simoneit and Mazurek, 1982]. This typical distribution of the homologues in aerosol usually results from the superposition of different distribution patterns characteristic of the major source types. The n-fatty acid homologues below C20, and possibly partly below C24, are derived directly from airborne microorganisms, or microorganisms colonizing the epicuticular wax on the plant surfaces. The microbial component, however, is also characterized by anteiso-C15, -C17, and -C19 fatty acids, and to a lesser extent, by iso-C14 and -C17 species [Cox et al., 1982]. The concentrations of these microbial tracers are typically higher in summer than in winter [Simoneit and Mazurek, 1982]. In addition seeds, plant organelles, leaf cells, chloroplast and pollen contain primarily palmitic, stearic, as well as mono-, di- and polyunsaturated fatty acids [Rogge et al., 1993a]. Microbial detritus could also arise from dust remobilization, as observed in aeolian particulates [Simoneit, 1977]. Additional anthropogenic sources of low molecular weight n-alkanoic acids (< n-C18) can be biomass burning, food preparation, vehicle exhaust, and tire wear debris. Among the carboxylic acid homologues hexadecanoic acid (palmitic acid) was found to to have the highest emission rate from biomass burning (14–21 mg kg−1 in pine and oak wood smoke, respectively), in accordance with the predominance of this compound in the biomass [Rogge et al., 1998]. Some of the n-alkanoic acids may partly be of secondary origin: for example, n-nonanoic acid can be a reaction product of oleic acid with ozone [Rogge et al., 1993b]. While the homologue series of n-fatty acids in aerosol usually exhibit a maximum at C22 or C24, those in plant wax maximize at C26, C28 or C30. This is because most of the n-fatty acids in epicuticular wax are bound as esters, so the presence of free fatty acids in aerosol possibly involves atmospheric decomposition (hydrolysis) of wax esters [Simoneit and Mazurek, 1982]. Therefore strictly speaking a considerable fraction of the lipid component can be of secondary origin, though in fact these compounds are never considered to be part of SOA. In case the maximum of this mode in aerosol is shifted to higher carbon numbers (C26 and C28), accompanied with pronounced even carbon number predominance (CPI 4–

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11), this might represent the unesterified fraction of plant wax, as in the case of background Artic aerosol [Li and Winchester, 1993]. The unsaturated fatty acids present in aerosol include the homologues of C14:1, C16:1 (palmitoleic) and C18:1 (oleic). In marine background aerosol even C20:1 acid could be detected, whereas forest aerosol showed the presence of pentadecenoic acid (C15:1) [Kavouras and Stephanou, 2002]. These compounds are indicators of recent biogenesis. The ratio of C18:0/ΣC18:1 was found to be variable in rural aerosol samples, the average being ∼13 in summer and ∼6 in winter [Simoneit and Mazurek, 1982]. This ratio can be used as an indicator for the age of aerosol, since the mono-unsaturated acid breaks down much faster than its saturated analog [Simoneit et al., 1988]. The high values at rural locations may reflect a combination of local biogenic emissions and aged aerosol being advected from urban areas [Brown et al., 2002]. C12–C18 fatty acid derivatives are abundant components of marine aerosol, as they together may make up 5–15 % by mass of the total organic material in the particles [Barger and Garrett, 1976]. It is likely that on alkaline sea salt particles the fatty acids are present as their sodium salts [Gagosian et al., 1982; Sicre et al., 1990; Seidl, 2000]. Generally palmitate is the most prominent individual lipid compound in marine aerosol. These species, which are well-known organic surfactants, possibly originate from the decomposition of marine phytoplankton, and are enriched in the ocean surface as a “biological oil slick”. They become airborne by the bubble bursting mechanism together with sea salt particles. The release of surface layer material by bubble-bursting was confirmed by the observed positive trend between organic material and iodine concentrations in freshly formed particles [Middlebrook et al., 1998]. As a result, these compounds are enriched in marine aerosol by a factor of 5–9 × 104 in relation to bulk seawater [Marty et al., 1979]. These lipids—due to their high surface activity—possibly constitute a surface film on sea salt particles, too. Time-of-flight secondary ion mass spectrometry of marine aerosol particles—which yield information on the outermost one or two monolayers—revealed the presence of palmitic acid and to a lesser extent other C14–C18 n-fatty acids as well as unsaturated C18 fatty acid [Tervahattu et al., 2002]. Their presence on the surface may have an important effect on the hygroscopic properties of sea salt particles. These effects are discussed in details in sub-chapter 5.2. The unsaturated fatty acids are also abundant in marine phytoplankton and are enriched in the surface microlayer slicks [Kawamura et al., 1996b]. They become part of sea salt aerosol just as their saturated counterparts, with the exception that the sodium salts of unsaturated fatty acids may undergo

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photooxidation in the atmosphere yielding sodium salts of dicarboxylic acids [Stephanou, 1992]. A recent compound specific radiocarbon and stable carbon isotope study shed light on the diverse origin of n-fatty acids in total suspended particulates [Matsumoto et al., 2001]. The samples were collected in summer, at a semi-urban site which was also influenced by marine air masses. The values of δ13C for C16–C18 fatty acids (−24.3 to −23.3 ‰) suggested that these species predominantly originated from marine organisms, in contrast with the fatty acids in the C24–C34 range (−27.9 ‰ to −30.3 ‰) which were derived from terrestrial C3 plants. The 14C ages of individual fatty acids were found to be rather diverse, ranging from modern carbon to ∼5900 years. The 14C ages of C16 fatty acid and the mixture of C18, C20 and C22 fatty acids were modern, confirming sea surface microlayer origin, though that of n-C16 fatty acid indicated contribution from the elevated atmospheric 14C level of atmospheric CO2 in the 1960s. The 14C ages of the mixtures of C24 and C26 fatty acids and C28, C30 and C32 fatty acids were 5900 and 270 years, respectively. c) N-alkanols (n-fatty alcohols) The n-fatty alcohols in rural aerosol covered a range of C10 to about C34 with a strong even carbon number predominance (CPI ∼ 4–20). The maximum concentration of the homologues was usually at C26, sometimes at C28 [Alves et al., 2001]. In another study, similar homologue distribution was observed in urban, marine and forest aerosol [Kavouras and Stephanou, 2002]. The homologues n-C20 derive from epicuticular waxes of higher terrestrial plants [Abas and Simoneit, 1996]. Unlike higher homologues of n-fatty acids, n-fatty alcohols from plant waxes usually preserve their characteristic fingerprint in the aerosol and constitute a major fraction of the lipid extract of the aerosol. As an example, these compounds can also be found in background Artic aerosol, while still retaining the distribution characteristics typical of higher plant waxes [Li and Winchester, 1993]. On the other hand, in combustion processes these wax constituents can be released less readily than, for example, the more refractory n-alkanes. As a result, n-alkanols are present at highly variable concentrations in biomass smoke due to their facile thermal degradation to olefins [Elias et al., 1999].

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Carbonaceous Aerosol

d) N-alkanals N-alkanals ranging from C9 to C23 were detected at an urban site, with maxima at C17 and C22. On the other hand, in background marine aerosol higher molecular weight n-alkanals were identified, with maxima at C26 and C28, exhibiting pronounced even carbon number predominance [Kavouras and Stephanou, 2002]. Mid-molecular weight n-alkanals are thought to be of primarily biogenic origin, though it is not clear whether they are emitted directly or formed in secondary reactions [Veltkamp et al., 1996]. e) N-alkan-2-ones The distribution of n-alkan-2-ones between C8–C38 showed an odd carbon number predominance in rural PM10 samples [Pio et al., 2001b]. The homologues below C20 are thought to derive from anthropogenic activities or microbial degradation, whereas longer chain n-alkan-2-ones are mainly from plant waxes. f) Long-chain wax esters An important class of high molecular weight compounds in continental aerosol was long chain wax esters, in the carbon number range of 38–58, with strong even carbon number predominance [Elias et al., 1999]. This class is comprised of palmitic acid esterified with the fatty alcohols ranging from C22 to C34, as well as minor amounts of stearic and eicosanoic acids esterified with the C32 and C34 alcohols. These compounds would normally elude analysis using conventional gas chromatographic techniques, but can be readily determined by high-temperature gas chromatography. The mass concentrations of these species were comparable to those of the n-alkanes. 4.2.1.2 Other High Molecular Weight Oxygenated Compounds of Primary Origin a) Cellulose Cellulose has recently been shown to be a major constituent of the insoluble fraction of organic aerosol, and a “macrotracer” for plant debris [Puxbaum and Tenze-Kunit, 2003]. Cellulose occur as linear polymers of 300–3000 D-glucose monomers connected by β-(1,4)-glucosidic bonds. The

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linear macromolecules form micelles by hydrogen bonding, and are organized into micro fibrils. This type of pure cellulose is termed as “free cellulose”, as they are readily accessible to saccharification. A considerable portion of these fibrils in leaves and woody tissues, however, are encapsulated by lignin, and in bark also by suberin. These materials also protect the encapsulated cellulose from saccharification, thus for the determination of total cellulose a delignification step must be performed. Since in leaves of terrestrial plants the mass fraction of total cellulose is about 50 %, the concentration of plant debris in aerosol is to be estimated by multiplying total cellulose content by a factor of 2 [ibid.]. Pollen also contains cellulose at concentrations of 3–7 % (m/m), and some bacteria, e.g. Acetobacter xylinum can even synthesize it. In a recent year-round study on urban aerosol, significant average mass concentrations (0.374 µg m−3) of total cellulose was observed, peaking in fall suggesting production from leaf litter [Puxbaum and Tenze-Kunit, 2003]. The secondary maximum in spring was attributed to seed production and to repulsion of plant materials. These concentrations translate into a relative contribution of plant debris of about 16 % to the mass of insoluble organic material. Somewhat surprisingly, “free cellulose” was found to make up 0.7 % of the total mass of fine aerosol, forming wettable, but insoluble particles. b) Humic acids and carbonaceous residue Simoneit and Mazurek [1982] found that humic acids constituted a major fraction of oceanic aerosol. Based on observed H/C ratio and δ13C values, a mixed origin from soil and lacustrine mud was inferred. The predominant fraction of the oceanic aerosol was residual carbon, appeared black and fine-grained. The H/C values for this fraction (H/C = 0.8–1.2) indicated a rather aliphatic character as compared to diesel soot (H/C = 0.4), and falling in the range of terrestrial protokerogens. The δ13C values of this fraction ranged from −22 to −24.4 ‰, which compared with the value of −25.3 ‰ for diesel soot. Simoneit and Mazurek [1982] assumed that residual carbon may originate from several sources, of which incomplete combustion of petroleum residues was likely predominant. 4.2.1.3 Major Compounds of Predominantly Secondary Origin a) Dicarboxylic acids Dicarboxylic acids, which are ubiquitous and abundant components of atmospheric aerosol, are generally considered to be of secondary origin. These species, however, may form in photochemical reactions from a

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Carbonaceous Aerosol

multitude of precursors, which can be of both natural and/or anthropogenic origin. Dicarboxylic acids can also form in the gas phase by the oxidation of volatile hydrocarbons and partition into the aerosol phase upon aging. Alternatively, unsaturated gaseous hydrocarbons may first be oxidized to aldehydes or ketones, which partition into the aerosol phase where they undergo further oxidation to yield formic, acetic or low molecular weight dicarboxylic acids. Within the class of dicarboxylic acids, low molecular weight species predominate. Among them, oxalic acid is the predominant compound in aerosol at virtually any locations. It is suggested to be formed by the photooxidation of aromatic hydrocarbons [Norton et al., 1983]. In an urban location, its concentration was found to exceed 100 ng m−3, and the majority of its mass was in the accumulation mode [Mészáros et al., 1997]. In the marine atmosphere, it is followed by malonic acid in abundance [Kawamura and Sakaguchi, 1999]. In haze sampled over the surface of the Indian Ocean, total dicarboxylic acids accounted for 2–7 % and 15 % of TC in the fine and coarse particles, respectively. Oxalic acid was the most abundant organic acid in both modes, followed by malonic and succinic acid in coarse particles, but malic acid and tartaric acid were similarly abundant in fine particles [Neusüss et al. 2002]. It should be noted that in addition to atmospheric formation, malonic and succinic acids were found to be predominant in the pine and oak smoke, respectively [Rogge et al., 1998]. In the case of pine wood combustion, the total class emission rates of these species were close to that of the n-alkanoic acids. In mid-latitude continental sites, the second most abundant dicarboxylic acid is typically succinic acid. Another major identified species are adipic acid (C6), fumaric acid (unsaturated C4 diacid) and malic acid (hydroxy C4 diacid). Small concentrations of 4-oxopimelic acid (keto C7 diacid), oxomalonic acid (keto C3 diacid) and maleic acid (unsaturated C4 diacid) can also be detected. The chemical structures of these species are shown in Figure 4.15. Albeit in marine aerosol both cis- and trans- configurations of unsaturated C4 diacid are detected, its trans configuration, fumaric acid was more abundant than the cis configuration, maleic acid. The average maleic acid to fumaric acid concentration ratios were low (0.26), in contrast with continental aerosol for which this ratio was greater than unity (1.5). This is because photooxidation of benzene and toluene yields predominantly maleic acid [Kawamura and Ikushima, 1993]. In the marine atmosphere, isomerization of maleic acid to fumaric acid was hypothesized [Kawamura and Sakaguchi, 1999]. Furthermore, a relative enrichment of the C2–C3 diacids was observed in remote marine aerosol, which was interpreted by the

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photochemical formation of these acids, especially in the region of the tropical Pacific.

Figure 4.15. Chemical structures of dicarboxylic acids detected in the marine aerosol (after Kawamura and Sakaguchi [1999]).

The total concentrations of the C2–C10 dicarboxylic acids over the western North and equatorial Pacific were in the range of 10 to 248 ng m−3 (average 62 ng m−3) [Kawamura and Sakaguchi, 1999]. The observed size distributions for C2–C4 dicarboxylic acids in marine aerosol collected in the central to western North Pacific were found to be bimodal [Mochida et al., 2003]. Whereas off the coast of East Asia only a small fraction of oxalic acid (11–25 %) was present on coarse particles, in the remote ocean the relative share of the coarse mode increased considerably (29–42 %). This trend was even more significant for malonic and succinic acids. The maximum in the oxalic acid distribution corresponded largely to that of the sea salt surface distribution, suggesting uptake of gaseous species by the coarse particles. Oxalic acid, in particular, may form the insoluble calcium oxalate with sea salt particles, which in turn can enhance uptake of gaseous oxalic acid. An alternative explanation can be that dicarboxylic acids are formed on the surface of the sea salt particles in heterogeneous reactions.

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Carbonaceous Aerosol

In some cases, for example in semi-urban aerosol, in which the dicarboxylic acids (172 ng m−3) made up of about 11 % of polar compounds extracted in methanol, markedly different homologue distributions were observed [Yokouchi and Ambe, 1986]. The most abundant dicarboxylic acid was azelaic acid (C9), which forms in the ozonolysis of oleic or linoleic acids. This was in contrast to a smog episode in Los Angeles, in which succinic and glutaric acids were more important, while azelaic acid was only a minor component [Grosjean, 1978]. In aerosol associated with savanna fires in Africa, considerable shift was observed between the smoke aerosol and haze in terms of the concentrations of the dicarboxylic acids [Gao et al., 2003]. In the smoke aerosol, oxalate was the most abundant species, followed by glutarate (C5) and succinate (C4). Strikingly, C3 dicarboxylic acid, malonate, was not detected. In contrast, in the haze aerosol again oxalate had the largest concentrations, but then malonate, succinate and glutarate followed. It seems that sequential decarboxylation and oxidation steps gradually reduce the chain length of dicarboxylic acids upon upward transport. In rural aerosol higher homologues of α,ω-dicarboxylic acids (C10–C24) can also be observed, with a strong even carbon number predominance and maxima at C12, C16 and C20 [Simoneit and Mazurek, 1982]. Sometimes only dicarboxylic acids with carbon number >C8 can be detected, as in rural aerosol in Finland [Alves et al., 2002]. Long-chain dicarboxylic acids are produced in soils by bacterial ω- and ω-1 oxidation of plant-derived fatty acids, and can be used as tracers for the atmospheric transport of dust aerosol [Kawamura and Gagosian, 1990b]. At Alert (Canadian Artic) a homologous series of α,ω-dicarboxylic acids was detected, characterized with a weak even-to-odd predominance, the relatively higher abundance of C9, C13, C16, C18, C20 and C22 homologues [Kawamura et al., 1996b]. The mean concentration of the total diacids was 1040 ± 170 pg m−3, and among them the concentration of higher diacids (C20–C26) was 271 pg m−3, peaking in late April. This implied that soil dust particles were transported to the Artic in spring, most likely from the Gobi desert region of China. There are also aromatic dicarboxylic acids which are ubiquitous in atmospheric aerosol. Among them, the isomers of benzenedicarboxylic acids (phtalic acids) are the most abundant. These species can either be produced by the photooxidation of naphtalene (a combustion-derived PAH), or from phenolic compounds present in sea surface slicks [Kawamura et al., 1996a]. The photochemical origin of dicarboxylic acids is supported by the fact that their concentrations showed maximum values during the intensified haze events following large-scale forest fires in Indonesia [Narukawa et al., 1999]. The total concentration of these compounds varied between 87 ng m−3 and

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3.8 µg m−3. The temporal variations in their concentrations followed closely those of TC and WSOC, with the exception of adipic acid, which showed maximum concentrations during weakened forest fire periods. This acid can be a secondary product of the photooxidation of cyclohexene [Hatakeyama et al., 1985]. It should also be noted that it is a precursor of 6,6-nylon, a synthetic polymer used for fiber material. Direct release from biomass burning was assumed to be the predominant source of all the other dicarboxylic acids. Dicarboxylic acids, once formed, might not remain unchanged in the aerosol phase. Although their very low equilibrium vapor pressure favors their partitioning overwhelmingly into the aerosol phase, some processes might cause selective losses. For example, the longer chain diacids (C8, C9) were shown to be more stable in the marine atmosphere than some of the shorted chain diacids (C5 and C6) [Kawamura et al., 1996b]. In addition, it was shown recently that several strains of bacteria and fungi consume C4–C9 dicarboxylic acids at pH >2 [Ariya et al., 2002]. The process is metabolical, and also produces lower molecular weight compounds such as acetamide, acetoacetic, butanoic, and propionic acids. The lifetime for the microbial biodegradation was estimated to be a few days, making the process potentially feasible in cloud and fog water as well as in aerosol particles, especially because other nutrients required by the bacteria and fungi are also available there. b) Other secondary species These species include a host of multiply substituted compounds, such as oxocarboxylic and ketocarboxylic acids. The major source of oxocarboxylic acids may be the photooxidation of cyclic olefins and unsaturated fatty acids. In rural aerosol in Finland, the only oxocarboxylic acid detected was glutaraldehydic (oxopentanoic) acid [Alves et al., 2002]. In urban aerosol in Los Angeles, the most abundant species was only tentatively identified as tetrahydro-2-furoic acid by its mass fragmentation pattern [Sakugawa and Kaplan, 1995]. However, the source or formation mechanism of this compound remains unknown. Among the ketocarboxylic acids detected in semi-urban particulates, 4-oxo-pentanoic acid and 4-oxo-pimelic acid were positively identified [Yokuchi and Ambe, 1986]. 4.2.2 Source-specific Organic Tracer Compounds Molecular tracers in aerosol are indicator compounds that can be traced down to individual sources. Such tracer molecules have definitive chemical structures which are correlatable either directly or indirectly via a set of

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diagenetic changes to their biogenic, geologic or synthetic sources [Simoneit and Mazurek, 1982]. Generally, the molecular tracers need to be characteristic of some sources, preferably not too many, and they also need to be largely unreactive within the time-scale of atmospheric transport. Such tracers need to be of primary origin with no contribution from secondary processes, and must have sufficiently low volatility not to suffer volatilization losses by gas-particle partitioning during atmospheric transport. Furthermore, the application of a tracer for source apportionment requires that it should be distinctive, its emission rate relative to total particle emission should be known, and each of its major sources be considered. Molecular tracer compounds are usually present at trace concentrations in atmospheric aerosol, so their contribution to the mass or atmospheric effects of aerosol particles is negligible. The only exception is levoglucosan, a tracer for biomass burning, whose mass concentration can be significant in biomass burning smoke. Nevertheless, these tracer compounds carry invaluable information regarding the relative contribution of various sources to particle mass at any given site. This information, however, can only be retrieved through carefully constructed atmospheric tracer models, as will be presented in sub-chapter 4.2.2. Universal combustion tracers such as BC or polycyclic aromatic hydrocarbons (PAHs) were and are still being used quite often in source apportionment studies. However, BC is not a single carbonaceous species therefore it is considered in bulk source apportionment in sub-chapter 3.3. On the other hand, in spite of their widespread use as combustion tracers, most PAHs have turned out not to meet most requirements for atmospheric tracers. In addition, being highly carcinogenic compounds, PAHs have been a subject of an immense body of scientific literature. Therefore their role as atmospheric tracers will only be addressed briefly before presenting the source-oriented approach. Polycyclic aromatic hydrocarbons (PAHs) The unsubstituted parent PAHs from naphtalane to coronene originate primarily from anthropogenic combustion sources, whereas natural burning of biomass (e.g. wildfires) emits alicyclic aromatic compounds [Simoneit, 1998]. The combustion-related PAHs (9 major nonalkylated compounds: fluoranthene, pyrene, benz[a]anthracene, chrysene, benzofluoranthenes, benzo[a]pyrene, indeno[1,2,3-cd]pyrene, and benzo[ghi]perylene) were found to be enriched in particles with d <1.5 µm, as opposed to petrogenic PAHs (phenanthrene, metyl- and dimethyl-phenanthrene) which were prevalent in particles with d >1.5 µm in urban atmosphere [Kavouras and Stephanou, 2002].

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While low-temperature burning processes (smoldering, pyrolysis) lead to decomposition products with alkyl groups attached to the aromatic core, high temperature combustion favors fused rings only, with the number of fused rings increasing with combustion temperature. The ultimate product of incomplete, high temperature combustion is graphite itself. The concentrations and sizes of PAH depends most strongly on the burning efficiency, and do not seem to be affected by the type of the biomass burned [Jenkins et al., 1996]. On the basis of these differences, very often diagnostic ratios of PAHs were set up, such as the phenenantrene/(phenantrene+anthracene) ratio representing the importance of petrogenic sources to biomass burning. Values higher than 0.70 are usually attributed to lubricant oils and fossil fuel combustion [Sicre et al., 1987]. The ratio of benzo(a)pyrene to (benzo(a)pyrene+crysene) is a measure of the contribution of vehicular emission. Its value is ∼0.16 in rural sites without vehicular traffic, and ∼0.33 in an urban environment [Gogou et al., 1996]. The ratio of indeno[1,2,3cd]pyrene to (indeno[1,2,3-cd]pyrene+benzo[ghi]perylene) is 0.18 and 0.37 for gasoline and diesel emissions, respectively, and 0.67 for wood burning [Kavouras et al., 1999c; Gogou et al., 1996]. However, PAHs with molecular weights below benzo[k]fluoranthene, with the possible exception of cyclopenta[cd]pyrene, were seen to suffer selective losses in the atmosphere either by volatilization or heterogeneous reactions [Schauer et al., 1996]. Therefore in source profiling only the four highest PAHs should be included, the indeno[1,2,3-cd]pyrene, indeno[1,2,3cd]fluoranthene, benzo[ghi]perylene, and coronene. For example, 14C analysis of individual PAHs revealed that benzo[ghi]perylene may be a good tracer for soot from fossil fuel burning [Currie et al., 1997]. Therefore many diagnostic ratios formerly in use for source apportionment may not be considered reliable. 4.2.2.1 Biomass Burning Tracers Biomass burning is a diverse category of anthropogenic activities that covers deforestation, slash-and-burn agricultural practice, pest and weed control, residential heating, waste disposal or food preparation. Globally, most of the mass of biomass burned is wood. Wood typically consists of various lignins (20–30 % dry weight), cellulose (40–50 %), hemicellulose (20–30 %) and extraneous compounds (extractives and ash together 4–10 %). Cellulose provides a supporting mesh which is reinforced by lignin polymers. The softwood genera, including pines (pinus), spruces (picea), larches (larix) and firs (pseudotsuga) have horizontal and vertical resin ducts and are prolific resin producers. During wood weathering the woody tissue undergoes photochemical degradation to yield organic acids, vanillin,

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syringylaldehyde and other water-soluble high-molecular weight compounds, mainly in the outmost layer [Rogge et al., 1998]. When the wood is heated, its constituents start to hydrolyze, oxidize, dehydrate and pyrolyze forming combustible volatiles, tarry substances and a highly reactive carbonaceous char. The onset of combustion marks the ignition temperature of volatiles and tarry substances. In the first flaming stage of the combustion the heat released fuels the gasification of wood substrate and the evaporation of capillary water first then the bound water stored in the cell walls. In this phase extractives such as resinous compounds and decomposition products of celluloses, hemicelluloses and lignin are vaporized then undergo partial or complete combustion in the flaming zone. If the flux of volatiles drops below a minimum level necessary to maintain flaming combustion, the smoldering phase starts which can be regarded as the gradual oxidation of the reactive char (solid phase combustion). In this phase, typically below 300 °C, additional volatile decomposition products evolve by volatilization or steam stripping. The first molecular tracer for wood combustion was retene, a fully aromatized PAH that results from the thermal alteration of resin acids present in conifer wood [Ramdahl, 1983]. It is no longer used as a tracer, but instead, there are various molecular tracers in use that carry the imprint of the type of wood burned, and of the stage of combustion. As a result of the processes during burning, biomass smoke aerosol contains a host of unaltered and thermally altered biomarker compounds from major vegetation taxa in the carbon number range of C8–C31. These compounds include phytosterols, lignans, lignin-derived phenolic species and diterpenoids from resins [Simoneit et al., 1993]. Certain of these substances are vaporized directly from the burning plant material then rapidly condense onto pre-existing particles without undergoing chemical transformation. This process is analogous to steam distillation/stripping and was documented for meat frying and grilling [Rogge et al., 1991]. Other species suffer pyrolysis and may be involved in dimerization, but still retain fingerprints of the vegetation taxa. Quantitatively, more than half of the mass of organic compounds identified in pine and oak wood smoke was assigned to specific lignin pyrolysis products, the rest was made up of resin acids (in the case of pine wood combustion), n-alkanoic acids, dicarboxylic acids, lignans, and phytosterols. Trace amounts of PAHs and oxy-PAHs were also detected in the smoke. It should be noted that even with a comprehensive analytical methodology only less than 25 % of the organic mass in the solvent extracted fraction of the particulates were identified [Rogge et al., 1998].

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a) Lignin-derivatives Lignin, the prevalent biopolymer in wood tissues is primarily derived from three aromatic alcohols, p-coumaryl, coniferyl and sinapyl alcohols, whose structures are shown in Figure 4.16.

Figure 4.16. Summary of the biochemical precursors for lignins and lignans, and of lignin burning products as tracers for biomass sources (after Simoneit [2002]).

The proportions of these biomonomers vary considerably among the major plant classes. Lignins in hardwood (angiosperms) are derived from sinapyl alcohol, those in softwood (gymnosperms) mainly from coniferyl alcohol with minor contribution from sinapyl alcohol, and grasses (Gramineae) contain the products from p-coumaryl alcohol. Upon burning of wood the thermal decomposition of lignin yields phenols, aldehydes, ketones, acids, and alcohols as breakdown products, generally with the retention of the original substituents (−OH, −OCH3) on the phenyl ring, as shown in Figure 4.17. Because monomers in the lignin are connected via the substituent para to the −OH group in the polymer, the lignin breakdown products differ only in that substituent. Guaiacol (2-(methylox)phenol) derivatives are typical of both hardwood and softwood smoke, whereas syringol (1,3dimethoxyphenol) derivatives are exclusive to hardwood combustion [Hawthorne et al., 1989]. Pine wood smoke contains primarily vanillin and vanillic acids, while oak wood smoke is enriched in syringaldehyde and syringic acid, beside a host of pyrolysis products from coniferyl and sinapyl lignin, respectively [Simoneit et al., 1993]. In grass smoke, phydroxybenzdehyde and p-hydroxybenzoic acid can be found, with minor amounts of other p-coumaryl-type lignin pyrolysis products. The individual vanillyl-, syringyl- or coumaryl-type compounds in themselves are not unique tracers of the type of biomass burned, but their relative proportions may carry useful information on possible sources.

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Carbonaceous Aerosol

Figure 4.17. Chemical structures of compounds formed in biomass burning. II—4-methyl-2-methoxyphenol, III—1-(4-methoxyphenyl)propane, IV—l-methyl-3,4dimethoxybenzene, VII—1-guaiacylpropane, X—guaiacylacetone, XII—1-(3,4-dihydroxy-5methoxyphenyl)ethanone, XIV—3,4-dimethoxyphenylacetone, XVII—1-(3,5-dimethoxy-4hydroxyphenyl)ethanone, XIX—1-syringylethanone, XX—1-(3,4,5-trimethoxyphenyl) ethanone, XXII—1-(3,4,5-trimethoxyphenyl)propanone, XXIII—3,4,5-trimethoxybenzoic acid, XXV—dianisyl, XXVI—divanillyl, XXVII—bis(3,4-dimethoxyphenyl)methane, XXVIII—divanillylmethane, XXIX—diveratryl, XXX—l,2-divanillylethane, XXXI— bis(guaiacylsyringy1), XXXVIII—disyringyl, XXXIX—tetrahydro-3,4-divanillylfuran, XL— tetrahydro-3-vanillyl-4-veratrylfuran, XLII—bis(3,4,5-trimethoxyphenyl)ethane, XLIII— tetrahydro-3,4-diveratrylfuran, XLIV—dihydro-3,4-diveratryl-2(3H)-furanone, XLV— dihydrovanillylsyringyl-2(3Zf)-furanone and isomers, XLVII—dihydro-3-(2’,3’,4’-trimethoxyphenyl)-4-veratryl-2(3H)-furanone and isomers (after Simoneit et al. [1993]).

However, the methoxyphenols appear not to have long atmospheric residence time as they probably degrade to other compounds such as cathecol, hydroxybenzoic and dihydroxybenzoic acids [Simoneit and Elias, 2001]. Therefore their application as tracers for wood burning should be largely confined to source areas. Apart from their release upon biomass

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201

burning, their emission into the atmosphere may also occur by exudation or volatilization of biodegraded detritus [Devi et al., 1997]. b) Lignans Many wood tissues contain lignans, i.e. dimers of p-coumaryl, coniferyl and sinapyl alcohols, which serve for toxins, supportive fillers or other purposes. These species are released unaltered in the smoke of softwood and to a lesser extent, in hardwood smoke. Major lignans in softwood smoke are matairesinol and shonanin, whereas hardwood smoke contains various isomers of dihydro-vanillyl-syringyl-2(3H)-furanone, dihydro-3(2`,3`,4`-trimethoxy-benzyl)-4-veratryl-2(3H)-furanone, and tetrahydro-3-vanillyl-4-veratrylfuran, as shown in Figure 4.17 [Simoneit et al., 1993]. It is also possible that lignin pyrolysis products undergo radical recombination to yield secondary dimers such as divanillyl (i.e. 1,2diguaiacylethane), divanillylmethane and divanillylethane in pine wood smoke and bisguaiacyl-syringyl and disyringyl in oak wood smoke, as shown in Figure 4.17 [Simoneit et al., 1993]. In grass smoke only dianisyl has been identified so far [Simoneit, 2002]. Since the lignans and secondary dimers generally retain the substituent pattern on the phenolic ring, they can complement lignin pyrolysis products in tracer applications [Rogge et al., 1998]. c) Diterpenoids Resin acids such as abietic or pimaric acids are biosynthesized mainly by conifers (gymnosperms) in temperate regions. Upon combustion of coniferous wood unaltered or partially altered diterpenoids are released which can be used as tracers [Standley and Simoneit, 1987]. The major marker compound emitted from the burning of conifer is dehydroabietic acid. Originally, the smoke also contains other resin acids, such as pimaric, isopimaric, sandaracopimaric and abietic acids, as well as rearranged and oxidized derivatives 8,15-pimaradien-18-oic acid and dehydroabietic acid. The chemical structures of these compounds are shown in Figure 4.17. Predominant ketones are abieta-8,11,13-trien-7-one which is oxidation product of abietane, a diterpenoid of higher plants, in particular conifers [Simoneit, 1989]. In fact, in pine wood combustion the unaltered resin acids such as abietic and sandaracopimaric acids showed the highest emission rates (42 and 47 mg kg−1 wood, respectively), and dehydroabietic acid was found to be the most abundant altered resin acid (37.2 mg kg−1) [Rogge et al., 1998].

202

Carbonaceous Aerosol

However, in the atmosphere all these resin acids are oxidized to dehydroabietic acid, which thus becomes the key tracer for coniferous wood burning, being considerably more abundant than retene, a high-temperature alteration product of resin diterpenoids [Simoneit and Elias, 2001]. Thus, dehydroabietic acid is a major component among the polar alcohol fractions in aerosol collected near coniferous forests. d) Cellulose tracers In woody tissues cellulose provides the supporting fibrous mesh (40– 50 % of dry mass), which is reinforced by lignin polymers. The biopolymers of wood also contain hemicellulose (20–30 % of dry mass). A cellulose molecule is a long-chain, linear polymer made up of 7000–12000 D-glucose monomers. The individual molecules form bundles (elementary fibrils) which are organized into larger parallel fiber structures [Simoneit et al., 1999]. In contrast, hemicellulose is a mixture of polysaccharides derived from glucose, mannose, galactose, xylose, arabinose, 4-o-methylglucuronic acid and galacturonic acid, made up of only 100–200 of these sugar monomers. Hemicellulose is also much less structured than cellulose molecules and their monomer composition varies widely among different wood species. Above the temperature of 300°C, cellulose decomposes by transglycosylation, fission and disproportionation reactions, yielding tarry anhydro sugars and volatile products, as shown in Figure 4.18. This mechanism gives rise to source specific molecular tracers, mainly the 1,6-anhydride of glucose called levoglucosan and the furanose isomer. It is emitted at an exceptionally high rate upon burning of wood (40–1200 mg kg−1 of wood) [Simoneit et al., 1999]. It also forms as pyrolysis product of lignocellulose and peat wood, and could result from the thermal alteration of carbohydrates and starch, but at much higher temperatures that are normally available during cooking and baking. In contrast, lignite and semibituminous coals do not produce levoglucosan upon combustion. The concentration of levoglucosan varies widely from 2.0 % for cowdung emission up to 15 % for pine smoke, as a result of the varying cellulose content of different biofuels [Sheesley et al., 2003]. Levoglucosan was shown to be resistant to atmospheric degradation [Simoneit and Elias, 2001]. Under specific conditions, such as in acidic cloud droplet, however, it was shown to be selectively removed by acid-catalyzed hydrolysis [Fraser and Lakshmanan, 2000]. Due to its exceptionally high concentrations and good atmospheric stability levoglucosan was suggested as a universal tracer for biomass burning.

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Figure 4.18. Schematic showing the major structures of the products from pyrolysis (T>300°C) of cellulose (after Simoneit et al. [1999]).

Two additional anhydrosugars, 1,4:3,6-dianhydro-α-d-glucopyranose and levoglucosenone have been found in aerosol associated with savanna fires in Africa [Gao et al., 2003]. Glucose and mannitol—also detected in smoke aerosol—are likely derived from the hemicellulose and pectin components of plants [Simoneit et al., 1999]. e) Sterols Sterols are produced by the marine and terrestrial biota in the range of C25–C30, with the predominance of C27 and C29, respectively. Phytosterols (sterols of higher plants) comprise C28 and C29 sterols, with β-sitosterol (C29) being the main constituent (Figure 4.19). The sterol distribution of algal (e.g. phytoplankton, diatoms) detritus differs from those of vascular plants, with a predominance of cholesterol (C27) (Figure 4.20) which is also the main representative of faunal sterols. Therefore β-sitosterol which was found in pine and oak wood smoke was a thermally unaltered compound and no indicators of thermal dehydration or pyrolysis of phytosterols (sterones, steradienes, etc.) were detected [Simoneit et al., 1993]. High emission rate of β-sitosterol (45.5 mg

204

Carbonaceous Aerosol

kg−1 wood) was determined in pine wood smoke, and this compound was also present in oak wood smoke. A product of the mild thermal degradation of β-sitosterol, stigma-4-en-3-one (Figure 4.19) was also detected in the wood smoke samples [Rogge et al., 1998].

Figure 4.19. Chemical structures of β-sitosterol and stigmast-4-en-3-one (after Rogge et al. [1998]).

Molecular tracers for dried cowdung burning are three stanols, 5βstigmastanol, coprostanol and cholestanol [Sheesley et al., 2003]. The chemical structures of these compounds are shown in Figure 4.20. Typically the concentration of 5β-stigmastanol in the aerosol resulting from burning of dried cowdung is much higher than that of coprostanol.

Figure 4.20. Structures of molecular markers for dried cowdung and jackfruit branch smoke (after Sheesley et al. [2003]).

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The β configuration of the C-5 proton of 5β-stigmastanol and coprostanol is characteristic of anaerobic microbial reduction of sitosterol and cholesterol in the digestive tracts of higher animals. The thermodynamically more stable product of the hydrogenation of cholesterol is cholestanol, in which the C-5 proton is in the α-configuration. This compound is formed preferentially during aerobic digestion of cholesterol and sitosterols in aquatic environments. Furthermore, cowdung is the only biofuel where octadecanoic acid is emitted in higher quantities than levoglucosan. f) PAH derivatives The dominant oxy-PAHs found in the smoke were polycyclic aromatic ketones (PAKs) and phenantranequinone, possibly formed by atmospheric photooxidation of parent PAHs [Korfmacher et al., 1980; Simoneit et al. 1991]. Polycyclic aromatic ketones such as 7H-benz[de]anthracene-7-one or 9-fluorenone, or quinones such as 2,6-di-tert-butyl-p-benzoquinone or 2ethyl-anthraquinone could either form by the oxidation of the parent PAHs or could also be decomposition products of lignin [Rogge et al., 1998]. 4.2.2.2 Fossil Fuel Combustion Tracers Most petroleum contains isoprenoid hydrocarbons (phytane and pristane). Pristane and phytane are diagenetic products of phytol and are not primary constituents of terrestrial organisms. Hopanes and steranes, molecular fossils which are present in crude petroleum, have long been used as tracers for motor vehicle exhaust [Simoneit, 1985]. What makes them good tracers for exhaust is that they are present in the lubricating oil and also in diesel fuel. For the same reason, however, their presence cannot differentiate between gasoline-powered and diesel vehicles. This calls for an additional tracer, which is typically soot (elemental carbon), known to be emitted mostly from diesel vehicles. Supplemental, but certainly less specific tracers for vehicle exhaust can be the highest molecular weight PAHs, coronene and benzo[ghi]perylene [Cass, 1998]. The chemical structures of the major organic tracers for fossil fuel combustion are shown in Figure 4.21.

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Carbonaceous Aerosol

Figure 4.21. The chemical structures of some useful organic molecular tracers for particulate air pollution sources (after Cass [1998]).

4.2.2.3 Molecular Markers for Terrestrial Vegetation Molecular markers for terrestrial vegetation are sesqui- and diterpenoids, which are predominant constituents of higher plant resins. The major diterpenoid compounds identified in aerosol are dehyroabietane and dehydroabietin. The triterpenones consist of α- and β-amyrone, which are oxidative products of α- and β-amyrin, major constituents of higher plant (Angiosperm) epicuticular waxes and gums [Abas et al., 1995]. The sesquiterpenoids include calamenene, 5, 6, 7, 8-tetrahydrocadalene and cadalene. The various isomers of cadinenes and cadinols are ubiquitous in essential oils of many higher plants, whereas the abietane-type diterpenoids

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are predominant components of resins in conifers [Simoneit and Mazurek, 1982]. A possible marker for secondary biogenic aerosol can be 6,10,14trimethylpentadecan-2-one. It results from the thermal alteration and oxidation of phytol from chlorophyll, which itself cannot be detected in rural aerosol [Simoneit and Mazurek, 1982; Simoneit et al., 1988; Alves et al., 2001]. This compound is ubiquitous in rural aerosol [Brown et al., 2002]. Alcohols such as pinene-2-ol, pinanediol, norpinanol, bornanol, longiborneol, are either directly emitted or formed by photooxidation of biogenic VOC [Pio et al., 2001a]. The contribution of vegetative detritus can be traced by the homologue series of high molecular weight n-alkanes (C27C34) with high odd carbon number preference. 4.2.2.4 Other Specific Source Markers One of the largest uncontrolled carbonaceous particle sources in cities is cooking. In particular, natural-gas fired charbroilers are notorious particle emitters. The dripping fat from the meat volatilizes on the hot surface and dispersed as a fine aerosol mist escaping through the exhaust hood. In most American cities these charbroiler plumes remained one of the last visible emissions in the 1990s [Cass, 1998]. Important molecular markers for these emissions are cholesterol and acyl monoglycerides (e.g. 1-palmitin) [Simoneit, 2002]. The dominant fatty acids emitted are nonanoic acid, hexadecanoic acid, and various C18 acids, such as octadecanoic, 9octadecenoic, 9,12-octadecadienoic, and 9,12,15-octadecatrienoic acids [Simoneit, 2002]. Cyclic ketones, such as 2(3H)-furanone may form by atmospheric oxidation of alkenals or emitted directly by meat cooking operations. In the latter source, it is formed by the lactonisation of hydroxyfatty acids found in triglycerides [Rogge et al., 1991]. Their concentrations in urban aerosol are typically below 1 ng m−3 [Alves et al., 2001]. Smoke from grilling of crustacean seafood contains markers of another biopolymer, chitin, particularly 1,6-anhydro-2-acetamido-2-deoxyglucose. Tracers for tobacco smoke are iso- and anteiso-alkanes whose concentrations in urban aerosol (0.28–0.36 µg m−3) were comparable to those estimated from the daily cigarette consumption in Los Angeles [Rogge et al., 1994]. This is because the leaf surface wax of the tobacco plant is enriched in iso-alkanes and anteiso-alkanes (anteiso-triacontane, anteiso-hentriacontane, anteiso-dotriacontane, see Figure 4.21). The best molecular tracer for tire dust is the styrene/butadiene copolymer that constitutes a large portion of most synthetic tire tread. Supplemental tracers include high molecular weight even carbon number nalkanes (e.g. C34, that is above the molecular weight range emitted by

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Carbonaceous Aerosol

petroleum sources) as well as benzothiazole (Figure 4.21) which is used as a vulcanization accelerator in the tire manufacturing process [Cass, 1998]. In the case of natural gas combustion, since there are no particulate species in the fuel, all carbonaceous components must be formed pyrogenically during combustion. The molecular tracer for emissions from natural gas-fired home appliances is benz[a]anthracene-7,12-dione, which has never been detected in vehicular exhaust [Cass, 1998]. 4.2.2.5 Source Apportionment Case Studies based on Atmospheric Organic Tracers Generally two approaches are used for source apportionment based on source emission estimates and measured ambient concentrations. Sourceoriented models use emission data and transport calculations to predict concentrations at specific receptor sites, and can be validated against atmospheric observations. Receptor-oriented models implicitly infer source contributions by calculating the best-fit linear combination of emission source chemical profiles needed to reconstruct measured profiles at the receptor sites. Mathematically, the concentration of tracer i at the receptor site k can be expressed as

cik =

m

∑f

ijk aij s jk

(4.3)

j =1

where sjk is the contribution of source j to the total particulate mass concentration at the receptor site k, aij is the relative concentration of tracer i in the total particle emission from source j, and fijk is the fractionation coefficient that corrects for selective losses (or gain) of tracer i during atmospheric transport. The fractionation coefficient should be near unity for a reliable atmospheric tracer. This concept is not affected by the presence of secondary organic aerosol as long as the compounds that can be formed in the atmosphere are not included in the calculations. If the number of chemical species included in the mass balance calculations exceeds the number of sources, the system of equations becomes overdetermined and the ordinary least-squares solution to the system is employed [Schauer et al., 1996]. Source profile collinearity might be a problem, so the chemical profiles for each source type must be different in a statistical sense. For example, in Los Angeles source profiles were not statistically different for hardwood and softwood combustion, or catalyst-

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equipped and non-catalyst gasoline powered vehicles, so they had to be combined into a single emission weighted average source profile [Schauer et al., 1996]. Among the urban carbonaceous particle sources, relatively few were identified as contributing to the majority of carbonaceous aerosol mass concentration. In Los Angeles, for example, among the 12 source types initially considered, emissions from industrial boilers burning distillate fuel oil, vehicular brake lining wear, and roofing tar pot effluents could not be detected in ambient aerosol. The nine sources that remained in the model accounted for ∼80 % of the primary organic carbon emissions according to the emission inventory [Schauer et al., 1996]. The source apportionment of fine particulate mass is detailed in Table 4.3. Table 4.3. Source apportionment of fine particulate mass concentration: 1982 annual average determined by chemical mass balance (avg ± std in µg m−3) (after Schauer et al. [1996]). Source Diesel exhaust

5.27±0.72

Downtown LA 11.6±1.19

Tire wear debris

0.29±0.11

0.22±0.09

0.25±0.09

a

Paved road dust

3.46±0.43

3.62±0.46

3.00±0.39

5.50±0.61

Vegetative detritus

0.33±0.10

0.24±0.12

0.38±0.11

0.18±0.08

0.047±0.02

0.040±0.019

0.034±0.016

0.029±0.008

0.18±0.03

0.26±0.045

0.20±0.028

0.19±0.032

2.41±0.46

1.74±0.34

2.03±0.39

1.94±0.35

1.63±0.20

2.12±0.23

1.44±0.16

0.34±0.05

2.70±0.43

1.85±0.31

2.65±0.41

0.54±0.10

1.46±0.66

1.16±0.76

b

b

1.94±0.44

5.9±0.60

6.6±0.65

5.9±0.60

5.8±0.51

2.1±0.27

3.0±0.54

1.9±0.29

10.4±1.2

2.6±0.34

3.0±0.37

2.3±0.23

5.1±0.59

28.3±1.5

35.5±1.9

25.3±1.4

37.3±1.8

28.2±1.9

32.5±2.8

24.5±2.0

42.1±3.3

Natural gas combustion aerosol Cigarette smoke Meat charbroiling and frying Catalyst and noncatalyst gasolinepowered vehicle exhaust Wood smoke Organics (other+secondary) Sulfate ion (secondary+background) Secondary nitrate ion Secondary ammonium ion Sum Measured a

Pasadena

West Los Angeles 4.36±0.64

1.03±0.71

Rubidoux 5.35±0.51

Not statistically different from zero with greater than 95% confidence, and therefore removed from CMB model. b Not statistically different from zero with greater than 95% confidence.

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Carbonaceous Aerosol

The model results showed that secondary aerosol can account for up to 31 % of total organic aerosol at the receptor site Rubidoux, whereas the upper limit for the other three sites was in the range of 15–18 %. These values compare well with previous estimates by Gray [1986], who estimated a secondary contribution of 27–38 % for Rubidoux based on measured OC/EC ratios. Pandis et al. [1992] estimated secondary aerosol formation during the 1987 summer smog episode as 5–8 % at coastal sites and 15–22 % farther inland. In general, source apportionment data based on organic tracers in rural and remote locations are extremely sparse. During the analysis of PM2.5 organic aerosol collected in the Big Bend National Park in Texas, the relative contributions of primary anthropogenic and natural sources were constrained by the evaluation of molecular tracers [Brown et al., 2002]. Since the concentrations of most organic tracers were at or below detection limit, only the maximum source contributions were established from the detection limits and available source profiles. The upper bounds for vehicular exhaust contribution were found to be between 0.5 % and 4.0 % (median 1.5 %) on the basis of the detection limit 17α(H)21β(H)-hopane. The high detection limit for the meat smoke tracer cholesterol gave unrealistically high maximum contributions (median 10.4 %). The only positively identified primary source was wood burning. Possible wood smoke contribution based on measured vanillin concentrations ranged from 1.3 % to 71 % with a mean of 13.1 % and a median of 5.5 %. Thus indirectly a very high contribution from SOA was inferred by subtraction, which was confirmed independently by bulk source apportionment based on measured OC/EC ratios (42–98 %) as well as by the detection of 6,10,14-trimethylpentadecane-2-one, a marker for biogenic SOA. Based on the concept of molecular tracers, Pio et al. [2001b] apportioned 49 % of total extractable lipids in PM10 at a rural location to vegetation waxes, 29 % to anthropogenic sources, dominated by vehicular exhaust, and the remaining 22 % to microbial sources.

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4.3 Chemistry of Water-Soluble Organic Compounds (WSOC) 4.3.1 Water-soluble Organic Compounds 4.3.1.1 Definition and Theoretical Considerations Saxena and Hildemann [1996] defined water-soluble organic compounds as those having solubilities >1 g/100 g water. Their compilation, which classifies major compound classes as insoluble according to this definition, is given in Table 4.4. Table 4.4. Water-insoluble organic compounds in atmospheric particles (after Saxena and Hildemann [1996]). Compound class

Solubilitya in water

C10–C34 n-alkanes

<0.002

C9–C30 n-alkanoic and n-alkenoic acids & their esters

≤0.026

C10–C35 n-alkanols

≤0.0037

C9, C14 aldehydes Dehydroabietic and other diterpenoid acids and retene

<0.1 ib

Phthalic and other aromatic polycarboxylic acids & their esters Polycyclic aromatic hydrocarbons (PAH)

0.003

Polycyclic aromatic ketones and quinones

N/Ab

Cholesterol and other steroids

0.26

1,2-dimethoxy-4-nitro-benzene and other aromatic Ncontaining compounds Lignans

N/Ab

Cellulose

≤0.7

0.03 ib

a

Solubility in g of solute/100g of water at temperature closest to 25 °C taken from values reported in 15 to 25 °C range. Solubilities for some compound classes are illustrated by the value for a specific compound in that class (e.g., naphthalene for PAH and anisic acid for lignans). Based upon available data for n-alkanes, n-alkanols and alkanoic acids, they assume that for a compound class, solubility in water decreases with increasing molecular weight. b Either solubility data not readily available (N/A) or listed as ‘insoluble’ (i). Based upon their molecular size and functional group composition, these compounds can safely be assumed to be water insoluble.

They also provided a comprehensive list of water-soluble organic compounds that might be present in atmospheric aerosol, based on their

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solubility and reported occurrence, if available. This list is far too extensive to be included here (the interested reader is referred to Table III of the original publication). Several compounds, including monocarboxylic acids and alcohols up to C4, diols and triols up to C7, or multifunctional acids such as lactic acid are completely miscible with water. The distinction between water-soluble and insoluble species is based on the solubility of the parent compounds: it should be noted that neutralization by alkaline ions can easily render a waterinsoluble compound water-soluble. Although not listed in the table, methanol dissolves a much wider range of organic species than water (e.g. C1–C17 nalkanoic acid and phtalic acid), consistent with the observations than methanol extracted more organic matter from aerosol than any other solvents [Grosjean, 1975]. The most frequently identified water-soluble species in atmospheric aerosol are low molecular weight dicarboxylic acids, with minor amounts ketoacids and dicarbonyls, yet they account for only 5–20 % of WSOC in urban aerosol. Therefore there is a large gap in our understanding regarding the major fraction of WSOC. The hypothesis that the remaining fraction would consist of water-insoluble compounds, which have small but finite solubility in water, can be tested. The amount of a solute dissolving in water depends not only on its solubility, but also on the amount of water used for extraction. If, however, we suppose that most of these compounds are quantified by extraction with a nonpolar solvent followed by GC-MS, this fraction is clearly insufficient to account for the unidentified portion. The second hypothesis, which seems to be more tenable, assumes that the unidentified fraction consists of water-soluble polar organic species. Since determination of highly polar compounds require targeted derivatization methods which are specific only to certain functional groups, it is not surprising that many polar species can elude analysis. The primary mechanism for the incorporation of water-soluble organic compounds into the aerosol phase is most likely absorption. When hydrated aerosol particles or hydrometeors are present, a water-soluble organic species can distribute between the gas and the aerosol (droplet) phase according to its air-water partition constant and the relative volume of the two phases. Saxena and Hildemann [1996] set a cutoff value of 106 M atm−1 for the effective air-water equilibrium constant for a compound to be considered as contributor to the water-soluble organic fraction of the aerosol phase. It is important to point out that absorption could take place at any partial pressure, so there is no need for significant gas-phase mixing ratio to build up, as in the case of homogeneous nucleation. Based on the combination of water solubility and effective air-water equilibrium, the following compound classes emerged as likely contributors to WSOC:

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• Dicarboxylic acids (C2 to C6) • Dicarbonyls (glyoxal) • Ketoacids (C2 to C5) • Multifunctional compounds (e.g., 1-malic acid) containing acidic, basic and carbonyl functional groups (C3 to C6) • Polyols (C2 to C7) • Hydroxyamines, amino acids (C2 to C6) and nitrophenol There are some discrepancies with atmospheric observations. For example, low molecular weight monocarboxylic acids such as formic or acetic acids, which are thought to be in the gas phase, were found in particulate phase in measurable concentrations [Krivácsy and Molnár, 1998]. The reason for such discrepancies can either be adsorption artifact in sampling, or modification of air-water equilibrium by aqueous phase reactions and/or the presence of an organic film. Nevertheless, the suggestions by Saxena and Hildemann [1996] were broadly consistent with observations and previous hypotheses. For example, as possible sources for polyols and amino acids, microbiological processes, direct emission from plants and degradation of organic litter were hypothesized [Wauters et al., 1979; Gorzelska et al., 1992]. Along this line, bioaerosol particles might also be potential sources for water-soluble organic compounds. Alternatively, polyols and other multifunctional compounds may be photooxidation products of anthropogenic or biogenic VOC. The direct emission of oxygenated volatile organic compounds by vegetation was estimated to contribute as much as 22.5 % to VOC global emission flux [Guenther et al., 1995], and the photooxidation of such unsaturated species was shown to yield bifunctional compounds such as hydroxy carbonyls [Grosjean, 1993]. The water-solubility of organic compounds is relevant in the hygroscopic properties of the particles. Insoluble compounds may form micelles or surface films to prevent condensation or evaporation of water to and from particles, whereas water-soluble species may absorb water. The expectation is that compounds having a finite solubility will deliquesce (like inorganic salts) and those being completely miscible will be hygroscopic (like sulfuric acid). Unfortunately, the properties of polar organic compounds are very sensitive to their structure: e.g. among the polyols, 2,5-hexanediol is completely miscible, 1,2-hexanediol is surface active, and 1,2,3-hexanetriol forms hydrophobic aggregates [Saxena and Hildemann, 1996]. This sensitivity makes any theoretical prediction of the hygroscopic behavior of ambient organic aerosol practically impossible. On the basis of group separation and HNMR measurements of the WSOC fraction, a representative mixture of individual compounds was suggested to simulate the physical and chemical properties of aerosol WSOC

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in model calculations [Fuzzi et al., 2001]. The composition of such a mixture is given in Table 4.5. Table 4.5. List of the organic compound classes chosen to represent the whole WSOC of the aerosol sample used. Specific examples of compounds identified in atmospheric aerosols are also reported (after Fuzzi et al. [2001]). Compound classes neutral compounds dialkyl ketone Polyol Polyphenol mono-/di-acids alkanedioic acid hydroxyalkanoic acid aromatic acid polycarboxylic acids

Examples pinonaldehyde (hydrated) levoglucosan catechol nonandioic acid 3-hydroxybutanoic acid 3-hydroxybenzoic acid fulvic acid

mass % in WSOC

References

16

1

9 2

2 2

14

3

15

4

3 41

5 6

References: 1,Yu et al. [1999]; 2, Simoneit et al. [1999]; 3, Rogge et al. [1993b]; 4, Souza et al. [1998]; 5, Yokouchi and Ambe [1986]; 6, Fuzzi et al. [2001].

The selection of the individual compounds should be regarded as a conceptual approach which would serve as a basis for further research on organic aerosol. 4.3.1.2 Atmospheric Observations of WSOC In the water-soluble extract of rural PM10 aerosol low-molecular weight hydroxy-dicarboxylic acids were found: malic acid (2hydroxybutanedioic acid) and 2-hydroxyglutaric acid were the most abundant [Carvalho et al., 2003]. Their total concentrations ranged between 10 and 90 ng m−3. Malic acid may either be produced by the hydroxylation of succinic acid [Kawamura and Ikushima, 1993], or as a by-product of the photooxidation of toluene or other monoaromatics [Hjorth et al., 2002].7 The photochemical formation of malic acid was supported by its mass size distribution, which peaked in the fine mode. Sugar alcohols containing carbon atoms and hydroxyl groups in a proportion of 1:1 were also found in rural PM10 aerosol. The most abundant compounds were mannitol and arabitol [Carvalho et al., 2003]. Polyols are 7

In the latter reactions, the main particulate-phase products were pyruvic and oxalic acids.

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common reserve materials in fungi (up to 15 % of dry weight), and play a role in osmoregulation. Whereas the mass size distribution of mannitol may support such bioaerosol sources at a rural site in Germany, the origin of its large abundance in the nucleation mode of forest aerosol is currently unknown [ibid.]. This is because secondary polyhydroxy compounds are assumed to also contain carbonyl groups [Pun et al., 1999]. Sugars such as glucose and sucrose were quite abundant components in the water-extract of in rural PM10 aerosol [Carvalho et al., 2003]. Their concentrations ranged between 28–180 and 16–213 ng m−3, respectively, at a rural site in Germany. Since the concentrations were more than an order of magnitude lower at a forest site (Hyytiälä) in Finland, their anthropogenic origin can be postulated. Biomass burning may be a major source, but polyhydroxycarbonyls could also result from reactions of alkenes, hydroxyalkenes, branched alkanes or branched alcohols with O3 or OH [Pun et al., 1999]. N-containing compounds Organic nitrogen compounds may represent an important component of water-soluble organic carbon, since most nitrogen-bearing functional groups are highly polar. Such compounds may include urea, organic nitrates, amino compounds, organic heterocycles, nitro-aromatics, humic substances and nitrogenated soot. Of them, amino compounds have been studied most extensively in the atmosphere. In a recent study, organic nitrogen and specifically amino compounds were determined in PM2.5 fine aerosol collected at a suburban site in California [Zhang et al., 2002]. Organic nitrogen typically represented ∼20 % of the water-soluble total nitrogen in PM2.5 through 9 months of the year. This ratio was not determined in summer due to the excessive and uncontrolled losses of inorganic N during sampling. Concentrations of water-soluble organic nitrogen and water-soluble free amino nitrogen were found in the range of 3.1 to 57.8 nmol N m−3 (median: 15.6 nmol N m−3) and 0.08 to 3.32 nmol N m−3 (median: 0.42 nmol N m−3), respectively. The latter class of compounds typically represents 2–4 % of the total organic N. However, water-soluble combined amino nitrogen compounds, which include proteins, peptides and other forms, were generally more important, being present at concentrations 4 times higher than the free compounds. It was speculated that the free species result from the hydrolysis of these more complex compounds in hydrometeors. Together, they accounted for as much as 8.5–71 % (median: 23 %) of the total organic N. In late fall and winter, typically higher water-soluble organic nitrogen concentrations were observed, superimposed on a relatively stable background concentration of combined amino nitrogen compounds. There seems to be a positive correlation between RH and total organic nitrogen concentrations, suggesting that these compounds favorably partition into

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aqueous aerosol or droplets. By assuming an average N-normalized molecular weight of 100 Da, the median mass of nitrogen-bearing organic compounds was calculated to be 1.6 µg m−3, representing as much as 18 % of total PM2.5 mass. Albeit this estimate is highly uncertain, given that Nnormalized molecular weight can range from 30 (for urea) up to ∼400 (soil humic substances), the mass contribution of organic N-species is likely to be significant. Consequently, organic N-species can play an important role in the control of fine particle hygroscopicity. In addition, those with basic functional groups may also affect the acid-neutralizing capacity of the particles, and compounds which possess a chromophoric N-functional group (e.g. nitro group) may also influence the light-absorbing properties of the fine particles. The finding of this study seems to contrast those which found nitrogen containing organic compounds, such as urea, nitro-phenols, metoxypyridine, isoquinoline, 1,2-dimethoxy-4-nitrobenzene, in atmospheric aerosol particles at very low concentrations [Cornell et al., 1998; Schauer et al., 1996]. Recently mass size distributions of WSOC have been evaluated in marine and continental air masses [Yu et al., 2004]. The mass size distributions are shown in Figure 4.22.

Figure 4.22. Size distributions of WSOC (upper row) and G1 (lower row) in air masses of marine (left), continental-NE (middle), and continental-E (right) origins (after Yu et al. [2004]).

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Regardless of their origin, WSOC in aerosol exhibited a bimodal size distribution, with a dominant fine mode and a minor coarse mode having mass mean AED of 0.7 ± 0.1 and 4.0 ± 0.3 µm, respectively. The mass in the fine mode ranged from two-thirds to four-fifths of that of the total WSOC. Both modes were further deconvoluted to low, medium, and high molecular weight polar compounds based on their thermal evolution features. While the low MW species had a bimodal distribution with a dominant coarse mode, the medium and high MW compounds exhibited a single peak in the droplet mode. This was interpreted as evidence that these latter species—which might also be humic-like substances—likely form during cloud-processing of aerosol. This finding would support the possible multiphase formation of HULIS [Gelencsér et al., 2003]. 4.3.2 Scavenging Efficiency for Organic Species Clouds drive a natural cycling process in which soluble material is accumulated from a large volume of air into a small volume of cloud water. Limbeck and Puxbaum [2000] sampled cloud water and interstitial aerosol at Mt. Sonnblick. They found octadecanoic and hexadecanoic acid, phenol-4methoxy-diisobutyl as well as oxalic acid in the interstitial aerosol phase, at concentrations in the order of 10 ng m−3. The composition of cloud water samples were different, as the highest average concentrations were found for dicarboxylic acids, in the order of oxalic, succinic, pyruvic and malonic acids (at concentrations of 174 ng ml−1, 85 ng ml−1, 50 ng ml−1 and 39 ng ml−1, respectively). Octadecanoic acid was also present in measurable concentration (21 ng ml−1), even aromatic compounds could be detected (benzoic acid-ethoxy-ethyl-ester, 9.6 ng ml−1). The calculated mass based scavenging efficiencies, (i.e. the concentration in cloud water per unit volume of air divided by total concentration), for four classes of organic compounds, alcohols, monocarboxylic acids, dicarboxylic acids and polar aromatic compounds, were 0.4, 0.4, 0.8 and 0.6, respectively. The scavenging efficiencies for individual dicarboxylic acids decreased from oxalic to azelaic acid, closely adhering to the trend of decreasing water solubilities. The observed nearly log-linear relationship between scavenging efficiencies and water solubility strongly indicated that more polar compounds had a tendency to deposit on more polar (hydrated) aerosol particles. This simple concept of scavenging efficiency broke down for semi-volatile species, which were also expected to be scavenged from the gas phase. In addition, the artifact resulting from their adsorption on the filter may lead to an overestimation of their particulate concentration, and consequently to an underestimation of their scavenging efficiency. The behavior of hexadecanoic and octadecanoic acids is rather unexpected, as these compounds are scavenged more efficiently than would

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follow from their solubilities. It is possible that the surface activity of these compounds may help their uptake by cloud droplets. This reasoning was strictly valid for externally mixed aerosol. In the case of internally mixed aerosol, however, nucleation scavenging efficiencies are largely governed by the main component. 4.3.3 Organic Species in Cloud/fog Water and Precipitation Likens et al. [1983] found TOC in precipitation collected at a suburban site to vary between 0.45 and 4.87 mg l−1, with a volume weighted mean of 2.37 mg l−1. At this site TOC concentrations were consistently higher throughout the summer month and declined steeply in autumn, remaining low in winter than increasing again in spring. At a rural site the TOC concentrations were found to be lower (0.10–3.23 mg l−1, mean 1.28 mg l−1), with somewhat less seasonal variations. The majority of organic carbon in precipitation was in dissolved form (DOC), particulate carbon (PC) accounted for 16 and 20 % of TOC at the rural and suburban sites, respectively, though the distinction between PC and DOC is ambiguous. The molecular weight distribution of DOC was dominated by compounds with molecular weight below 1000 Daltons, though taken together PC and high molecular weight DOC (>1000 Daltons) accounted for more than half of the TOC. Of the latter, about half was attributed to carbohydrate and tannin/lignin compounds, whereas only a small portion (<10 %) of macromolecular carbon was proteinaceous. The largest identified fraction of the low molecular weight compounds was the class of carboxylic acids, dominated by C1–C5 monocarboxylic acids as well as lactic, glycolic, caproic, succinic and citric acid. Kawamura and Kaplan [1986] found n-alkane concentrations in precipitation at a rural site to be between 1.6 and 5.1 µg dm−3, corresponding to 14.6–20.1 % of the total organic matter identified. In contrast, the unresolved complex mixture was observed to be a minor component in rural precipitation samples. The major PAHs identified were phenantrene, fluoranthrene and pyrene, the total concentrations of all PAHs were in the range of 27–80 ng dm−3. Most of the identified organic mass (62–69 %) was represented by n-fatty acids, found in the concentration range of 5–17 µg dm−3. Their molecular distribution showed a pronounced even carbon number predominance, and unsaturated fatty acids, oleic (C18:1), linoleic (C18:2) and palmitoleic (C16:1) were quite abundant, in some cases they were the major species. In urban precipitation, the abundance of C5–C11 monocarboxylic acids was in some cases prevalent, but never in rural precipitation. Among the aromatic carboxylic acids, benzoic acid, 3-methyl

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benzoic acid and 4-methyl benzoic acid were detected in urban and rural precipititation. These species were assumed to derive from motor exhaust and used engine oil. Phenol and its derivatives, 2- and 4-methylphenol, 2nitrophenol and 2-nitro-3-methylphenol were found to be major organic compounds in urban precipitation, but were also detected in rural precipitation at total concentrations of 120–480 ng dm−3. Bank and Castillo [1987] analyzed insoluble particles (>0.2 µm) in cloud water collected at the summit of Whiteface Mountain, New York (1524 a.s.l.). They identified black particles as being ubiquitous and major components in all samples, and observed the distinguishing infrared absorption bands for both protein and cellulose. Brown particles, which were also found in varying amounts, showed the imprint of decomposed proteins. They seemed to be more abundant in clouds with lower pH, indicating the importance of possible acidic amine hydrolysis. Furthermore, colorless, opaque fibers were identified as cellulose. Li and Winchester [1993] found water-soluble organic compounds in background Artic aerosol. The compounds can be classified into two groups in term of their observed abundances: formate, acetate, oxalate and benzoate had median concentrations above 10 ng m−3 in both the coarse and fine size fractions, whereas the concentrations of lactate, propionate, pyruvate and methanesulfonate were below 5 ng m−3. Together with other compounds that were not measured but presumed to be present, water-soluble organic compounds contribute about 10–16 % to the total water-soluble mass of Artic aerosol. In the snowpack, acetate was the most abundant, followed by lactate, oxalate and methanesulfonate. Interestingly, the total concentration of these species relative to that of nss-SO42− was much higher in the snow than in the aerosol (1.14 and 0.18, respectively). This enhanced scavenging of organic species can be explained by in-cloud production of organic acids, scavenging of gas-phase organic species or preferential scavenging of organic aerosol particles. The 14C/12C isotopic ratios were determined in the water-insoluble carbon in snow collected at Mt. Sonnblick, Austria (3106 m a.s.l.). The average concentration of the water-insoluble carbon was about 310 µg dm−3. The 14C content was found to be between 64 and 88 pMC, with a weighted mean of 73.7 ± 1.7 pMC, indicating that ∼64 % of the filterable carbon in high alpine snow was of biogenic origin. Gorzelska et al. [1992] measured primary amine compounds in rural precipitation. Low molecular weight aliphatic amines are stronger bases than ammonium (pKb 4 and 5, respectively), and they tend to form water-soluble salts with mineral and carboxylic acids. Concentrations of nitrogen in individual amino acids and amines ranged up to 2.4 µM. The most frequently occurring amino compounds were methyl amine, ethanol amine, glycine, glutamic acid and serine. Quantitatively, arginine, asparagine, glutamine, methyl amine, serine, and alanine predominated. There was a pronounced

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temporal trend in the concentrations of individual species: whereas amino acids such as arginine, asparagine and glutamine were most abundant in spring, aliphatic amines, especially methyl amine were more important in late fall and winter precipitation. Although microorganisms can contribute to a variety of amino compounds, the observed springtime maxima in their concentrations cannot be solely attributed to the microbial component of the precipitation. It was hypothesized that plants in their early stage of development could be an important direct source of free amino compounds. Another potential source can be the degradation of organic litter which is also favored by conditions during springtime. Alternatively, as a secondary mechanism, free amino compounds could result from the degradation of proteinaceous material in marine areas, as suggested by Mopper and Zika [1987]. Formaldehyde was also found in rural aerosol in Germany at concentrations which were only a minute fraction of gaseous concentrations [Klippel and Warneck, 1978]. Nevertheless the observed concentrations were about 1000 times higher than equilibrium concentration calculated from simple gas-liquid equilibrium between gas-phase formaldehyde and aqueous aerosol. They argued that formaldehyde in aqueous aerosol may be complexed with some soluble species. The distribution of formaldehyde between the gas and liquid phase is largely determined by the formation of hydroxyalkylsulfonate adduct (HMSA): RCHO(aq ) + S (IV ) ↔ RCH (OH )SO3−

(4.4)

Facchini et al. [1992] found that formaldehyde was almost entirely present in polluted fog water as HMSA (<10 % in free form), which also accounted for up to 90 % of the total S(IV) species present in fog water. The ratio of HMSA to free HCHO increased with increasing pH. Overall, HMSA formation caused the equilibrium concentration of HCHO to increase by a factor of 100 in fog water.

Chapter 5 PROPERTIES OF CARBONACEOUS AEROSOL AND THEIR ROLE IN THE GLOBAL ATMOSPHERE 5.1 Optical Properties of Carbonaceous Particles and their Role in Radiative Transfer in the Atmosphere This sub-chapter deals with the interaction of carbonaceous aerosol components with solar radiation and its atmospheric consequences. From the climate’s point of view, these consequences can be summarized as direct effects, or direct forcing when only the anthropogenic contribution is taken into account. Throughout the discussion—except for the sub-chapter dealing with light-absorbing organic compounds—organic chemistry of the aerosol will not be explicitly considered. Carbonaceous aerosol is divided into the two principal classes of organic aerosol (carbon) and BC, largely corresponding to two distinct types of interaction with solar radiation. Organic aerosol is thought to primarily scatter solar radiation, whereas BC is understood to be the principal light-absorbing component of tropospheric fine aerosol. This simple classification, however, does not improve significantly our understanding of the global radiative effects of carbonaceous aerosol. First, observational data on the concentrations of carbonaceous aerosol, especially of organic aerosol are rather sparse in space and time. Secondly, sampling of organic aerosol is often loaded with severe artifacts, whereas distinction between BC and OC is rather uncertain and method-dependent. The resulting uncertainties are carried directly into the calculations of optical effects of the aerosol. Furthermore, the treatment of the interactions with solar radiation also requires detailed knowledge of fundamental physical properties of carbonaceous aerosol such as size distribution, mass scattering and absorption efficiency, hygroscopic growth factor, etc., in which chemistry is implicitly involved. Experimental data for these properties are either barely available, or if they are, they differ by a large factor. In spite of any effort to adhere strictly to the definition of carbonaceous aerosol, its state of mixing with other aerosol constituents will be one of the key parameters governing its optical effects, primarily for BC. 221

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This sub-chapter is organized into three sub-chapters. The first subchapter deals with the wavelength-dependent optical properties of individual aerosol components, namely the mass scattering and absorption coefficients of organic and black carbon, as well as the dependent variable of single scattering albedo, a key parameter in radiative transfer calculations. These can be derived experimentally, by relating aerosol optical measurements to measured bulk chemical concentration. Alternatively, these can also be calculated from the complex refractive index and size of the particles by the Mie-code. The mass absorption efficiency is not constant even at a single wavelength, but depends on the state of mixing with other aerosol constituents, which can be accounted for in the calculations. In the second sub-chapter, these basic parameters are used together with concentration data to derive a key parameter for radiative transfer calculations, aerosol optical depth, which can also be retrieved from both ground-based and space-borne measurements. Albeit optical depth is an additive quantity which allows the scattering and absorption of carbonaceous aerosol to be treated separately, closure studies require the understanding of the contribution of all other aerosol components as well. In the third sub-chapter, the derived key parameters are fed into radiative transfer models to assess the role of carbonaceous aerosol in the attenuation of solar radiation. Carbonaceous aerosol represents by far the largest source of uncertainty in the assessment of direct forcing by anthropogenic aerosol. The main reasons for this are the paucity of data on the concentrations and size distributions of carbonaceous aerosol, in particular of organic aerosol, the generally poor understanding of the anthropogenic and natural aerosol sources, as well as the variability and uncertainty in the optical properties of carbonaceous aerosol particles. Nevertheless, there are a few model estimates for the direct forcing term associated with anthropogenic carbonaceous aerosol. Since most of these models include the sulfate aerosol whose direct effects are better understood, the estimates for carbonaceous aerosol are presented in comparison with those for sulfate. 5.1.1 Optical Properties of Individual Carbonaceous Aerosol Components In principle, the fundamental optical property of a homogeneous aerosol particle is its refractive index which—together with its size and shape—would allow its interaction with radiation to be characterized with well-established mathematical codes. However, carbonaceous aerosol particles or internally mixed particles with a carbonaceous component are generally far from being homogeneous. In addition, even BC particles which have relatively uniform chemical composition do not have a single material constant for refractive index. But apart from these limitations, in radiative transfer calculations it is

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more practical to use optical properties that are directly related to the mass concentration of carbonaceous aerosol species. This is because we can conveniently measure bulk concentration of aerosol carbon, but we have virtually no information on individual carbonaceous particles, if any. Such property is the wavelength-dependent mass extinction efficiency (αext) which relates the extinction coefficient (be) to the mass concentration of a given species therein (c) be = c × α ext (5.1) The unit of αext is m2 g−1, therefore it is sometimes referred to as specific extinction cross-section. The extinction coefficient has the favorable property of being additive, i.e. equal to the sum of light absorption (ba) and scattering (bs) coefficients:

be = b a + b s

(5.2)

It follows that a mass scattering (αscat) and mass absorption (αabs) efficiency can be defined analogously to αext above. Furthermore, if the aerosol is externally mixed, the extinctions caused by the individual aerosol species are additive, thus total extinction can be calculated knowing the mass concentrations and mass extinction efficiencies of each species. Consequently, αscat and αabs can be defined for any individual aerosol species which are assumed to make up individual aerosol particles. In case the aerosol is not externally mixed the above simple approach will not be strictly valid. But even then, statistical methods allow species-specific αscat and αabs to be retrieved from concurrently measured bs and ba values and mass concentrations of all major aerosol species. However, in this case these parameters will not have an exact physical meaning; instead they will refer to the properties of the individual species as if they were externally mixed. The mass extinction coefficient αext and in particular αscat depends markedly on relative humidity. With increasing RH, the size of the particle changes with the uptake of water, and so does its geometric cross-section which interacts with radiation. Therefore αext and αscat are usually defined for dry conditions, and their RH dependence is expressed in the form of a separate function as fRH(αscat). For αabs generally no dependence on RH is assumed in radiative transfer models, though some recent studies have suggested that this may not be necessarily true (see sub-chapter 5.2). The single scattering albedo, ω0, the ratio of the aerosol scattering coefficient to the extinction coefficient, is a critical parameter in radiative transfer calculations. It is a dimensionless quantity, and has a range between 0 and 1. ω0 is not an independent variable, as it depends on two independent

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variables, αscat and αabs. It follows from the above discussion that it also has a marked dependence on RH, ω0 increases with increasing RH. The use of mass extinction (scattering, absorption) efficiencies has another advantage, namely that it allows an important and measurable optical property of the atmosphere, the aerosol optical depth (τ) to be derived. For a given mass load m (column mass per surface area), τ can be calculated by

τ = α ext × m

(5.3)

A disadvantage of estimating aerosol optical depth from column mass loads and mass extinction efficiency the loss of information on the vertical distribution of aerosol, which is particularly crucial in the case of absorbing aerosol [Tegen et al., 1997]. With the assumption of externally mixed aerosol, τ has the feature of being additive with respect to the absorption and scattering of all individual aerosol species. Physically, the aerosol optical depth expresses the attenuation of radiation through a vertical atmospheric column:

Φ = Φ 0 × e −τ

(5.4)

where Φ and Φ0 is the radiant flux (W) at the surface and at the top of the atmosphere, respectively. In atmospheric radiative transfer modeling aerosol optical depth is one of the key parameters combining the concentrations and optical properties of all individual aerosol species. 5.1.1.1 Mass Scattering Coefficients for Organic Aerosol Particles Organic aerosol were long known to contribute predominantly to the scattering of incident solar radiation, its absorbing components have just recently been recognized (see Chapter 4). Therefore traditionally a complex refractive index for organic aerosol of m = 1.43 − 0.0035i was assumed (for the wavelength of 0.55 µm), a value close to that of sulfate aerosol [Liousse et al., 1996]. The mean mass scattering efficiency for organic particles (for the same wavelength) was assumed to be 4 m2 g−1 for low humidity conditions. This compares to the mean αscat of 5 m2 g−1 for sulfate. It should be noted that this parameter can either be given for organic carbon mass which is directly measured, or for organic mass, which is related to the former through the assumed mass conversion factor. A first approach to predict the dependence of αscat on RH was to assume similar hygroscopic growth as that for sulfate aerosol [Liousse et al., 1996]. This would increase the αscat of organic aerosol to 6.8 m2 g−1 for RH=80 %.

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Similarly, Shah et al. [1984] assumed αscat of 5 m2 g−1 for dry carbonaceous aerosol, identical to that of sulfate. A more rigorous approach was taken by McInnes et al. [1998], who measured the hygroscopic growth factor for aerosol scattering, fRH(bs), with a dual-nephelometer humidigraph on an island in Canada, a site which was influenced by both marine and polluted air masses. The hygroscopic growth factor for aerosol scattering means the ratio of the light scattering at a given relative humidity to the dry light scattering. This term is different from— albeit explicitly includes—the hygroscopic growth factor, which accounts for the effect of relative humidity on the geometric size of the particles. The hygroscopic growth factor for aerosol scattering is shown as a function of relative humidity in Figure 5.1.

Figure 5.1. Hygroscopic growth factor of aerosol scattering, fRH(bs), measured for the marine and anthropogenically influenced case. The greater scatter for the marine case is caused by much lower aerosol concentrations (after McInnes et al. [1998]).

The differences in fRH(bs) cannot be explained by the different particle size distributions alone, which would predict scattering hygroscopic growth factors for aerosol scattering of 2.8 and 6 at RH = 85 % for the marine and polluted aerosol assuming identical chemical composition [ibid.]. The large predicted value for anthropogenic aerosol resulted from the fact that its measured dry size distribution was dominated by small particles with less scattering efficiency (Rayleigh scattering). As shown in Figure 5.1, the measured fRH(bs)for the marine aerosol is actually very close to the predicted value (2.7), but it is only 1.8 for the anthropogenic aerosol. These results clearly showed that the differences in chemical composition play a much larger role than particle size distribution in determining the hygroscopic growth of the scattering efficiency of the aerosol. It can be anticipated that organic aerosol play a significant role in the polluted air masses, though the results do not refer explicitly to carbonaceous aerosol or to any other aerosol component. Since most of the organic aerosol is

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internally mixed with inorganic aerosol components, which would probably dominate water uptake, it seems reasonable to assume that the hygroscopic growth factor of aerosol scattering is governed by the main inorganic components. It might be confusing that while the concept of αscat assumes externally mixed aerosol, the hygroscopic growth factor for aerosol scattering is estimated as if the aerosol were internally mixed. 5.1.1.2 Mass Absorption Efficiency of BC The mass absorption (αabs) efficiency of BC is the counterpart of the mass scattering efficiency (αscat) of organic aerosol in radiative transfer calculations. Its determination and application, however, are subject to much larger uncertainty than those of αscat. First, the measurements of the absorption coefficient ba are much less straightforward since they may be loaded with large measurement biases. These biases primarily result from the fact that the most common methods are filter-based, that is ba of the aerosol is not measured for the suspended state, but for particles embedded in a filter matrix. It is well known that filter-based measurements of light absorption also respond to non-absorbing particles (the “scattering artifact”), and there is an orientation-dependent artifact called the “broadside enhancement”. For detailed discussion of these measurement artifacts see sub-chapter 2.2. Therefore the quantity that can be directly measured by filter-based methods can be more correctly termed as the specific attenuation crosssection (σa), which refers to the apparent absorption of aerosol particles constrained by the filter substrate. This parameter, which has the same unit (m2 g−1) as that of αabs, is a useful calibration parameter if BC mass concentrations are determined directly by the specific filter-based optical method. On the other hand, it does not have much to do with αabs of BC in the suspended state unless the above artifacts are corrected for. Nonetheless, there is a great deal of confusion about the use of these parameters in radiative transfer calculations, in which the specific attenuation cross-section, often denoted with σa, is used indiscriminately in place of αabs. The possible consequences of this mix-up will be detailed in this sub-chapter. The second major source of uncertainty stems from the ambiguity of the definition of BC and of its measurement methods. As a large fraction of organic material has been found to be light-absorbing, the distinction between highly absorbing and non-absorbing particles may not be straightforward, except when absorption is dominated by purely black particles (as those emitted by diesel engines) [Bond, 2001]. However, when the sources of absorbing particles are dominated by low-temperature combustion processes (e.g. lignite combustion), the borderline dissipates. These findings imply that the term “black carbon” needs to be reconsidered in atmospheric sciences. A

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more universal term “light-absorbing carbon” seems to be more appropriate, since absorbing material does not always appears black, not even near combustion sources. At this point it should be recalled that in atmospheric science BC is very often seen as a single specific entity throughout the atmosphere, originating from the sole source of combustion, and has a black color and uniform chemical composition and physical properties. Albeit no material constant was assumed for its refractive index, in particular for the imaginary part which is uncertain by a factor of about 2, most of the observed variations in αabs or σa were attributed to changes in mixing state and morphology of BC particles [Liousse et al., 1993]. This overly simplistic view of BC has been recently challenged from many sides, stating that BC covers a wide range of compounds which may not be always black, may also originate from noncombustion sources, a part of them may be organic, and consequently their physical and chemical properties can be rather diverse. However, these recent findings of atmospheric chemistry have not yet infiltrated into global atmospheric radiative transfer modeling. Theoretical calculations The absorption coefficient of a bulk material (babs) can be derived simply from the imaginary part of its complex refractive index (k) as

babs = −4π k λ

(5.5)

However, there are several materials which, in bulk, appear black, but still their absorption as particles is negligible. An example is a substance having transmission of 10−10 in a slab of 1 cm at 550 nm is seen as completely black (babs = 2072 m−1) [Horvath, 1995]. However, using Eq. 5.5 k is obtained as k = −9 × 10−6, an imaginary part which is typical for nonabsorbing aerosol particles. Backward calculation shows that if this substance is placed in a slab of 1 µm thick, the transmission will be 0.998, i.e. practically transparent. This is the “dirty silicate story” by Bohren and Huffman [1983], a piece of hornblende, which was as black as coal, but still had no light absorption when it was assumed to be an aerosol particle. The mass absorption efficiency of a single BC particle can be more rigorously calculated by the Lorentz-Mie theory provided that the refractive index, size and density of the particle are known. These calculations are strictly valid only for homogeneous spherical particles, though their use has been extended to estimate the scattering and absorption properties of nonspherical particles, using volume or surface area equivalent spheres. If the mass absorption efficiency of a carbonaceous sphere is calculated using the Lorenz-Mie theory, the values for αabs will not exceed 7.5 m2 g−1 at

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λ = 0.55 µm unless dubious values for refractive index, size or density are assumed [Fuller et al., 1999]. If the radiative forcing of BC particles is to be calculated from their mass concentration, as it is usually the case, density should represent the material density of the spherules, and not that of their ramiform (branched) or aciniform (packed) aggregates. Its value is typically around 2.0 g cm−3. There is no material constant for the refractive index of black carbon: its value is dependent on the type of the fuel and the conditions of combustion. The refractive index of BC typically used in calculations varies between 1.5 and 2.0 for the real part and between 0.4 and 1.0 for the imaginary part [Chýlek et al., 1995]. D’Almeida et al. [1991] assumed a refractive index of m = 1.75 − 0.44i and a density of 2 g cm−3, which have been used widely in model calculations. The Lorentz-Mie calculations also show that the mass absorption coefficient of pure BC particles decreases steeply with increasing particle radius. It follows that large homogeneous BC particles are much less efficient absorbers of solar radiation. If, however, the large particles are clusters of small spherules rather than homogeneous spheres, their optical properties are also affected by their fractal dimension. An important feature of such fractal structures is that their specific absorption efficiency is expected to be nearly constant, and equal to that of the individual spherules, and not to decrease steeply with increasing radius as would follow from the Mie calculations [Chýlek et al., 1995]. In fact, αabs is thought to increase slightly upon agglomeration of the primary spherules in the atmosphere. For small (d = 10–50 nm) spherules, this enhancement is expected to be less than 15 %, and for larger spherules of 100 nm, even some decrease can be observed [Fuller et al., 1999]. The maximum enhancement (∼30 %) can be seen upon agglomeration in the case of paracrystalline graphitic carbon which is contained in Diesel exhaust. Light absorption by chain aggregates can therefore be approximated as the sum of the absorption by the individual spherules, as long as the fractal dimension remains below 2.0. Branched (ramiform) soot aggregates, which are typical of fresh particles, have fractal dimensions in the order of 1.8. On the other hand, aged, cloud-processed soot forms aciniform aggregates with fractal dimensions > 2 [Sheridan et al., 1992]. Calculations revealed that chains of carbon spheres in water or ammonium sulfate do not exhibit a marked absorption enhancement. This is believed to be due to the lower relative refractive index and the resulting reduced electric polarizability of the individual spherules in the medium. Unfortunately, absorption by chain aggregates is subject to a large measurement artifact which can also be evaluated theoretically. If a chain aggregate consisting of sufficiently small primary spherules is aligned with the direction of the propagation of incident radiation, a minimal value for σa is measured; on the contrary, if the orientation is perpendicular, the value for

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σa will be at its maximum. This effect is known as ‘broadside enhancement’,

and it is an artifact of all filter-based collection and measurement methods. This artifact can be as high as 60 %, and is most pronounced for longer chains of small primary spherules. Fuller et al. [1999] calculated σa in the order of 10 m2 g−1 for chain-like soot aggregate with orientation constrained by the filter substrate. This value was clearly higher than the αabs of the same aggregate (7.5 m2 g−1) randomly oriented in the atmosphere. Due to their more compact morphology, aciniform clusters are expected to be much less affected by such an artifact. Consequently, aged soot might present a lower σa than fresh soot when measured with filter-based optical methods. The light absorption by particles is often measured at only one wavelength, and is extrapolated to other wavelengths using the power-law relationship for the spectral dependence of absorption:

α abs = K × λ− Å

abs

(5.6)

where αabs is the mass absorption efficiency of the particles (m2 g−1), K is a constant, λ (nm) is the wavelength of light, Å is the Ångström exponent. For small spherical particles for which the refractive index is independent of wavelength, Å equals to 1.0. This restriction of size is meant relative to the wavelength of light as 2 rπ |m|/λ «1 [Bond, 2001]. For reasonable particle size distributions and a refractive index of m = 1.55 − 0.5i, the theoretical values of Å varies between 0.2 and 1.2. When absorption is stronger at shorter wavelengths, i.e. the particles appear brownish or yellow, Å will be greater than 1.0. For example, particles from smoldering combustion [Patterson and McMahon, 1984], or from industrial combustion of lignite [Bond et al., 1999b] exhibit such behavior. The relationship between mass absorption efficiency and Å for bituminous coal emission is shown in Figure 5.2. It can be seen that whereas some of the aerosol particles from bituminous coal combustion exhibit a larger mass absorption efficiency and Å value near unity, others have progressively stronger absorption at shorter wavelengths, as follows from Å values greater than unity, and appear yellow or brown. The marked negative correlation between Å and αabs indicates that aerosol with the weakest light absorption has the strongest spectral dependence. It was speculated that the weakly absorbing coal tar, consisting of partially aromatized organic matter, can escape from the coal matrix by devolatilization. Thermal processing of this tar during combustion leads to a progressively greater degree of graphitization. Thus aromatic compounds emitted from bituminous coal burning may cover a continuum between a few conjugated aromatic rings (coal tar) and an extended network of aromatic rings (black carbon). By increasing the degree of graphitization, the energy gap between the highest ground state and the lowest excited state closes

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which enhances absorption efficiency and simultaneously shifts absorption towards longer wavelength (lower energy). If this hypothesis holds true, the traditional classification into light-absorbing BC and purely scattering OC may not be applicable to coal combustion smoke [Bond et al., 2002].

Figure 5.2. Relation between mass absorption efficiencies and Ångström exponents for bituminous coal emissions (after Bond et al. [2002]).

These findings also imply that mass source estimates of global BC emissions from fossil-fuel burning (8 TgC yr−1 by Cooke and Wilson [1996], 6.6 TgC yr−1 by Penner et al. [1993], and 1.9–8.3 TgC yr−1 by Turco et al. [1983]), which were measured largely for regulatory purposes, may not be suitable for quantifying atmospheric light absorption [Bond et al., 1998]. Therefore, instead of the conventional mass emission approach, a new quantity was proposed, the source strength of absorption, Qabs, in units of m2 s−1. Albeit being a derived quantity, it can be calculated directly as a product of two measured quantities, the exhaust absorption coefficient (in units of m2 m−3) and the flow rate (m3 s−1). As a first approximation, it was assumed that the light absorption per emitted mass of aerosol is a conserved quantity. Therefore the source term Qabs can be directly inserted into a dispersion model to calculate the three-dimensional distribution of the light-absorption coefficient ba, analogous to the modeling of BC mass concentration from the mass source term QBC. Alternatively, an absorption emission index (EIabs) can be defined based on this concept, as the absorption cross section per mass of fuel burned (in units of m2 kg−1). This concept was applied to measurement of emissions from a small industrial plant burning lignite, considered to be one of the principal sources of BC emission by Cooke and Wilson [1996]. Bond et al. [1998] observed that the particulate matter collected appeared light brown to yellowish, but

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certainly not black as would be expected for pure BC by definition. The measured average absorption emission index, 0.27 m2 kg−1, can be compared with that calculated as a product of mass emission factor of BC of 1 g kg−1 for lignite burning and the mass absorption efficiency of BC of 5.4 m2 g−1, which is equal to 5.4 m2 kg−1. So the derived quantity is a factor of 20 higher than that based on direct measurements. One reason for this large difference could originate from the possibly much lower mass absorption coefficient of the lignite combustion particles, as indicated by the color of the deposits on the filters. Another factor might be the inclusion of the absorption caused by larger particles, which increases the absorption index by a factor of 4 as compared to that of only fine particles. The values of EIabs is by far the largest from bituminous coal combustion, those from lignite and hard coal burning are lower by a factor of about 50 and 200, respectively. Generally, solid-phase combustion tends to produce very little absorption and scattering. It appears that BC formation is strongly related to the tar content of coal, though the chemical composition of the tar also has an effect. This is especially true for the initial period of the combustion, when yellow particles with Å as high as 2.5 can be released by devolatilization of tar from the solid matrix. The fact that the Ångström exponent may differ significantly from unity, means that caution must be exercised when extrapolating absorption measured at a single wavelength over the solar spectrum. In smoke aerosol BC particles are likely coated with a non-absorbing shell by the condensation of organic compounds or water. Therefore for modeling optical properties, a realistic representation of such particles is a layered sphere with a BC core. Whereas for pure BC particles αabs remains below 10 m2 g−1 over the entire particle size range, αabs will increase significantly if the BC particle is surrounded with a non-absorbing shell [Martins et al., 1998]. This happens because the non-absorbing shell increases the cross-sectional area of the particle that is available to incident light, and focuses light toward the absorbing core thereby the BC inside such a particle can absorb more radiation than a standalone BC particle would. This is shown in Figure 5.3, where BC mass absorption efficiency, αabs is shown at λ = 0.55 µm for such a layered structure. For the calculations, the refractive indices of BC and the non-absorbing shell were taken as m = 2.0 − 1.0i and m = 1.5 − 10−6i, respectively, whereas corresponding densities were 1.85 g cm−3 and 1.5 g cm−3.

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Figure 5.3. Calculated values of BC mass absorption efficiency (αBC) at λ = 0.55 µm for an internal mixture of BC and nonabsorbing material in a layered structure. The particle structure consists of an absorbing BC core with a surrounding nonabsorbing shell. The refractive index of the BC was assumed to be 2.0 – 1.0i, the density 1.85 g cm−3; for the nonabsorbing shell, the refractive index was assumed to be 1.5 – 10−6i and the density 1.5 g cm−3. The amount of BC and the radius of the shell determine the radius of the BC core. The radius of the shell is defined as equal to the radius of the particle (after Martins et al. [1998]).

As it can be seen, the effect of a non-absorbing shell increases with particle radius up to a maximum αabs. For the smallest BC fraction of 0.5 % by mass, at an overall particle radius of 0.25 µm, αabs can be as high as 30 m2 g−1. In general, the absorption enhancement tends to increase with the mean size of the embedding sulfate particles, and with decreasing size of the primary spherules. The dependence of αabs on the particle radius can be explained by the changes in the relative size of BC core. For small particles about 0.05 µm in radius, the αabs will be around 10 m2 g−1, irrespective of the relative volume of the non-absorbing shell. Albeit not shown here, values for αabs will increase with increasing real refractive index of the non-absorbing shell. It was also calculated that specific absorption tends to drop as soot is moved away from the center of the composite particle, being at its lowest when soot was located just beneath the surface [Fuller et al., 1999]. When the soot particle was located on the exterior surface, the surface enhanced absorption was only in the order of 40 % for composite particles above a diameter of 0.2 µm, but rapidly fell off with decreasing particle size. When averaged over realistic particle size ranges, αabs did not exceed 15 m2 g−1, and this value can be obtained only if all soot particles were encapsulated within sulfate particles somewhat larger than usually observed. Thus it may be inferred that cloud-processed aerosol could contain BC with considerably enhanced mass absorption efficiencies [Fuller et al., 1999]. Calculations also

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revealed that the αabs of packed clusters was typically below 20 m2 g−1 over realistic particle size ranges, largely similar in magnitude to those of the layered sphere model [Martins et al., 1998]. Another case which was thoroughly evaluated theoretically was when a BC particle is located at various positions inside a droplet. Chýlek et al. [1995] computed that if a BC particle of 10 nm radius was located within a droplet of 0.5 µm in diameter, its specific absorption efficiency may change from the initial 4 m2 g−1 (characteristic of the individual particle) to over 100 m2 g−1, when the BC particle was positioned just under the surface on the shadow side of the droplet, near the axis parallel to the incident radiation. At selected positions within the droplet, this enhancement can be as high as 20, but there are also locations where the refractive properties of the droplet actually shields soot from incident light [Fuller et al., 1999]. Averaged over the volume of the droplet, this enhancement was about a factor of 2–2.5 relative to that of the individual BC particle. For BC particles, Haywood and Ramaswamy [1998] calculated the specific extinction coefficient, ω0 single scattering albedo, and g asymmetry factor as a function of λ for rm = 0.0118 µm and σ = 2.0, as shown in Figure 5.4.

Figure 5.4. Specific extinction coefficient ke, single scattering albedo ω, and asymmetry factor g for BC aerosol modeled by a lognormal distribution with a geometric mean radius of 0.0118 µm and a geometric standard deviation of 2.0 (after Haywood and Ramaswamy [1998]).

The low single scattering albedo of BC (∼0.2) indicates that absorption dominates over scattering throughout the entire solar spectrum. The very low values of the specific extinction coefficient in the near and thermal infrared indicate very little interaction with radiation in that spectral region.

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For characterization of BC alone, sometimes the co-albedo of the aerosol (1−ω0) is used. In biomass burning smoke its value was found to decrease upon aging from 0.16 near the source to 0.12 after about 210 min downwind [Abel et al., 2003]. The co-albedo of 0.10, a value typical of aged regional haze was expected to require ∼5 hours of aging. The imaginary part of the refractive index derived from Mie scattering calculations were 0.018i and 0.025i for aged regional haze and fresh smoke, respectively, indicating that changes in aerosol absorption rather than changes in particle size distribution were likely responsible. This was supported by theoretical Mie calculations of the co-albedo of aggregates consisting of 65 individual BC spherules with an assumed radius of 11.8 nm and a complex refractive index of 1.75 − 0.44i. Figure 5.5 shows how the modeled co-albedo decreases as the chain-like structure collapses upon aging, with a value of 0.17 for chain-like to 0.10 for a packed cluster.1

Figure 5.5. Evolution of the aerosol co-albedo downwind of a fire. Diamonds represent the mean value for each cross plume transit measured with the PSAP and nephelometer. Error bars are the standard deviation from this mean. Values using the PCASP size distribution for each transit and Mie calculations (m = 1.54 – 0.018i) are shown with an asterisk. The exponential fit suggests that it will take 5 hours for the co-albedo to reach 0.10, which is representative of the aged regional haze (after Abel et al. [2003]).

Experimental data Due to the limitations detailed in the introduction of this sub-chapter, field measurements usually determine σa rather than αabs of BC particles, thus their findings are not directly comparable to—albeit correlate with—the αabs 1

The modeled co-albedo values are much higher than the measurements because only the absorbing component of the aerosol is considered.

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values obtained from theoretical calculations. For example, in airborne measurements Mader et al. [2002] found an average σa of 11 m2 g−1 ± 5.0 m2 g−1, with significant event-to-event variability likely due to the differences in chemical composition, mixing state and/or size distribution of BC. Some of the variability was attributed to the presence of dust particles, since the d50 for absorption measurements was found to be ∼8 µm. The signal of the soot absorption photometer appeared to vary erratically during ascent and descent of the aircraft, ostensibly due to physical or chemical changes in the filter, but unrelated to the levels of light absorbing species. Measured σa values in biomass burning smoke in Brazil spanned a wide range between 5.2 and 58.2 m2 g−1 [Martins et al., 1998]. The absorption measurements were carried out with an aethalometer, whereas the determination of BC was based on the thermal method of Cachier et al. [1989b]. The high values appeared to be outliers, associated with high levels of potassium (the K/BC ratio was higher than 0.78). Since potassium was shown to exert a catalytic effect on the combustion of BC by lowering its volatilization temperature [Novakov and Corrigan, 1995], it may cause underestimation of the amount of BC by the thermal method, and consequently overestimation of σa. By discarding these extreme cases from the data set, the range of σa values reduced to 5.2 and 19.3 m2 g−1, with an average of 12.1 ± 4.0 m2 g−1. Concurrent SEM studies showed that lower σa values were associated with spherical and relatively homogeneous particles, whereas high σa values were determined for high concentrations of nonspherical particles and compact clusters. Mass absorption efficiencies for bituminous coal burning emissions were about half of the value for pure BC, making them less absorptive. Particles emitted by lignite combustion had mass absorption efficiencies about an order of magnitude lower than that of pure BC, which was indicative of the high fraction of organic compounds and mineral matter. Compared with literature values in global BC inventories [Liousse et al., 1996; Cooke and Wilson, 1999], which assigned a single value to BC irrespective of its source, light absorption by bituminous coal emissions seems to be underestimated by a factor of three, whereas that by lignite emissions seems to be grossly overestimated [Bond et al., 2002]. Absorption by organic aerosol components A relatively recent hypothesis that does not fit into the traditional concept of BC aerosol is that reduction in downward UV irradiance— especially in urban areas—may be attributed to absorbing aerosol components including nitrated aromatics, benzaldehydes, benzoic acids, aromatic polycarboxylic acids, phenols, polycyclic aromatic hydrocarbons, and to a lesser extent, nitrated inorganic compounds [Jacobson, 1999]. There are several reasons why this absorption cannot be included in the absorption by

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BC. First, these organic compounds can be determined separately from BC, in other terms they have nothing to do with the tail of the BC signal extending to the organic region in, for example, thermal methods. Second, their absorption has a very strong wavelength-dependence, far beyond that of BC which is approximated by the power-law relationship. Because of absorption by nitrated and aromatic compounds tail off between UV and visible wavelengths, global UV irradiance is attenuated much more than total solar irradiance. The presence of nitrophenols, in particular 4-nitrophenol and 2-methyl4-nitrophenol, originating from the photochemical reactions of benzene and toluene, was observed in urban aerosol [Nojima et al., 1983]. These compounds were also detected in aerosol and precipitation in pristine locations [Rippen et al., 1987]. O’Brien et al. [1975] found that as much as 17 % of organic material in Los Angeles were organic nitrates, with one in every nine organic acid molecules carrying a nitrate group. This hypothesis was supported by calculations based on wavelength-dependent imaginary refractive indices estimated from molar absorption spectral data. In particular, nitrated aromatic compounds, PAHs and benzaldehydes were found to be the most effective UV-A and UV-B absorbers in terms of their imaginary refractive indices and absorption wavelength. Despite the strong absorption of nitrated and aromatic aerosol in the UV, the absorption is not expected to induce significant photochemistry. On the contrary, the decreased UV-flux might even reduce ozone mixing ratios in Los Angeles by 5–8 % [Jacobson, 1999]. Model column simulations with the inclusion of enhanced aerosol absorption still underpredicted the observed UV irradiance reductions by 8 and 25 % for an urban and rural site, respectively. The absorption by organic aerosol represented about ¼ of the total downward UV attenuation, primarily due to non-nitrated and nitrated aromatic compounds in proportions of 10.2–13.5 % and 8.8–14.7 %, respectively. The higher share of nitrated aromatic compounds was found for the more remote site, in accordance with the significantly higher concentration of organic nitrates. 5.1.2 Observations Regarding the Optical Effects of Carbonaceous Aerosol 5.1.2.1 Contribution of Carbonaceous Aerosol to Visibility Impairment A large scale visibility monitoring program was initiated in spring 1988 in the U.S., called IMPROVE (Interagency Monitoring of Protected Visual Environments) [Malm et al., 1994]. Concurrent with the optical measurements, concentrations of the major aerosol species were also measured. Specifically, fine aerosol particles were collected on pre-fired

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quartz filters and analyzed by the thermal-optical reflectance (TOR) method for organic and elemental carbon [Chow et al., 1993]. The mass of organic carbon was multiplied with the factor of 1.4 to obtain the mass of organic aerosol. Light-absorbing carbon was determined from laser integrating plate method measurements, using the specific attenuation coefficient of 10 m2 g−1. A major finding of this study—obtained from detailed statistical analysis of the relationships between the different variables—was that lightabsorbing carbon is not only associated with elemental carbon, but also with ∼25 % of high-temperature organic carbon, a component which had been traditionally considered non-absorbing. The mass concentrations of fine organic aerosol showed a spatial gradient from California and the Pacific Northwest with average concentrations of 2.0–3.1 µg m−3 to the desert areas of the western United States with mass concentrations of 1.0–1.5 µg m−3. The highest average organic aerosol concentrations, 2.0–4.5 µg m−3, were found in the eastern United States, whereas the lowest (0.7–0.9 µg m−3) were measured in Alaska and Hawaii. These values typically referred to background locations. It should be kept in mind, however, that the values reported were organic aerosol mass concentrations (and not organic carbon concentrations), and they possibly carry a significant positive bias due to the adsorption artifact on pre-fired quartz filters. For visibility closure calculations, a dry mass scattering efficiency of 3 m2 g−1 and a mass absorption coefficient of 10 m2 g−1 were assumed for organic aerosol (and not for organic carbon!) and light absorbing carbon, respectively [Malm et al., 1994]. Furthermore, half of the organic aerosol was assumed to be water-soluble whose hygroscopic growth factor was the same as that of ammonium-sulfate. It was found that organic aerosol was the second largest contributor to light extinction throughout the Unites States, with average contribution of about 15–20 %, but can be as high as 40–45 % at the northwest. On the contrary, contribution of light-absorbing carbon to total extinction was much less significant, being about 3 % in the eastern United States, and highest on the Colorado plateau (10–15 %). 5.1.2.2 Contribution of Carbonaceous Aerosol to Aerosol Column Optical Depth In a case study, Hegg et al. [1997] performed a chemical apportionment of aerosol column optical depth off the mid-Atlantic coast of the United States. Airborne measurements and aerosol sampling were performed up to an altitude of about 4 km, accounting for ∼90 % of aerosol optical depths. During sampling of carbonaceous compounds the dual-filter strategy was used to correct for possible positive artifacts. Dry aerosol scattering and absorption coefficients were measured at all altitudes, and at selected altitudes with high

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values of dry scattering coefficients the hygroscopic light scattering coefficients were also measured during descent. The aerosol column optical depth was measured with sunphotometers. The total mass concentration of aerosol was inferred from dry aerosol volume concentrations measured by a passive cavity aerosol spectrometer. To achieve closure, sulfate and carbon mass was determined from the filters, and mass conversion factors of 1.19 and 1.7 were used to convert them into the mass of their most likely molecular forms, respectively. Using this approach a mass closure within ∼10 % was achieved on the basis of carbonaceous and sulfur species alone. The contributions of these species to the dry aerosol light scattering coefficients were determined by multiple linear regression (r2=0.83, with regression line forced through zero). From the regression equation the mass scattering efficiency of carbonaceous aerosol was found to be 4.0 ± 1.1 m2 g−1, higher than that of sulfate species (2.7 ± 1.3 m2 g−1). Recall that this parameter refers to carbonaceous matter and not to aerosol carbon, for which it needs to be multiplied with the conversion factor of 1.7. The contour plot of the dry aerosol light scattering coefficient is shown as a function of sulfate and carbonaceous dry mass concentrations in Figure 5.6.

Figure 5.6. Contour plot of the dry aerosol light-scattering coefficient (shown as parameter in units of m−1) as a function of the sulfate and carbonaceous dry aerosol mass concentrations (after Hegg et al. [1997]).

The pattern of the isopleths reveals that up to a carbonaceous mass concentration of ∼12 µg m−3 the light scattering depends slightly more strongly on the concentration of carbon species. Above this value, however,

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the isopleths bend toward the sulfate axis, suggesting that at higher concentrations its effect is more pronounced. The frequency plot of the contributions of carbonaceous species to dry aerosol light scattering is shown in Figure 5.7.

Figure 5.7. Frequency distribution of the fraction of dry aerosol light scattering due to carbonaceous species (after Hegg et al. [1997]).

The distribution indicates that dry aerosol light scattering is dominated by carbonaceous aerosol and not by sulfate as has been previously suggested. Furthermore, a fraction of carbon species also contributes to aerosol light absorption, which can make up as much as 25 % of total aerosol dry extinction. The contribution of BC aerosol to aerosol optical depth could be even more significant in South-East Asia and Central America where more polluting engine types are in use. At ambient relative humidities, however, condensed water becomes the single most important contributor to aerosol extinction, accounting for on average 35 % of the column optical depth. The share of condensed water increases with increasing optical depth; for optical depths > 0.2 its mean contribution exceeds 50 %. This confirms that aerosol hygroscopicity is a key parameter in regulating aerosol direct radiative forcing. Since the study by Hegg et al. [1997] considered column optical depth and not simply a lightextinction budget at the surface, it was more relevant for radiative transfer

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calculations than ground-based studies in general. The enhanced carbonaceous aerosol contribution to column optical depth can be explained by its observed higher mass concentration aloft, possibly due to its lower nucleation scavenging compared to sulfate [Novakov et al., 1997b]. In another case study, Chiapello et al. [1999] estimated the contribution of carbonaceous aerosol to aerosol optical depth over the northeastern tropical Atlantic during the winter period when the transport of North African dust was significant. Their optical depth measurements were performed at 670 nm at which the effect of Rayleigh scattering was the lowest. They sampled total suspended particulates without a size selective inlet. Among the major aerosol components, mineral dust predominated, sometimes with mass concentrations in excess of 200 µg m−3, and the contribution of sea salt particles was also significant (in the range of 10–30 µg m−3). In contrast, the average carbonaceous mass concentration was very low, on average 1.5 µg m−3, distributed between BC (0.53 ± 0.37 µg m−3) and particulate organic matter (0.98 ± 0.86 µg m−3, assuming a mass conversion factor of 1.3). The mass extinction efficiencies at 670 nm were estimated from previous measurements, and were found to be 7.83 m2 g−1 for BC and 3.34 m2 g−1 for particulate organic matter. In spite of the fact that carbonaceous aerosol can be found in the size range that interacts most efficiently with solar radiation, its contribution corresponded only to an optical depth of ∼0.02, which was about 12 % of the total aerosol optical depth (mean value). The relative contribution carbonaceous aerosol to total aerosol optical depth increased with decreasing wavelength. The contribution of absorption by BC to aerosol optical depth was recently estimated for biomass burning smoke for various regions. Thin African savanna is consumed quickly in intense flaming fires, not much left for the smoldering phase, associated with the emission of vast amounts of black carbon. As a consequence as much as 12 % of the smoke τ can be attributed to BC [Eck et al., 1999]. On the contrary, in boreal forest fires, in which the smoldering phase is predominant, the share of BC absorption in τ was much smaller (5 %) [Dubovik et al., 2002]. Since smoke aerosol is generally less hygroscopic than regional pollution aerosol, the contribution of water to τ is typically low (10–20 %) at relative humidities of 80–85 % in South America and Africa [Ross et al., 1998]. Tegen et al. [1997] calculated the composition of the total aerosol optical thickness for various locations of the world by model simulations, as shown in Figure 5.8.

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Figure 5.8. Aerosol “optical thickness” composition. (a) Minimum contribution of absorbing aerosol (specific extinction: clay, 1 m2 g−1; silt, 0.2 m2 g−1; SO4 and organics, 8.5/14 m2 g−1 (land/ocean); black carbon, 8 m2 g−1; and sea salt, 0.4 m2 g−1). (b) Maximum contribution of absorbing aerosol (specific extinction: clay, 2 m2 g−1; silt, 0.4 m2 g−1; SO4 and organics, 5/7 m2 g−1 (land/ocean); black carbon, 12 m2 g−1; and sea salt, 0.2 m2 g−1) (after Tegen et al. [1997]).

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The model results indicated that the contribution of BC to total aerosol optical thickness was highest in Europe. In general, the contribution of carbonaceous aerosol was most pronounced over the industrialized regions of Europe and Southeast Asia, as well as over areas with significant biomass burning emissions. 5.1.3 Role of Carbonaceous Aerosol in Radiative Transfer in the Atmosphere In the light of the above, it is beyond any doubt that carbonaceous aerosol particles play a large role in the transfer of solar radiation in the troposphere. Of the two major carbonaceous aerosol components, organic aerosol and BC, the atmospheric effects of the latter are much more thoroughly characterized, at least qualitatively. Although the radiative effects of BC are more complex and involve various types of interactions with other aerosol constituents, the better availability of observed concentration data (and global climatologies by the mid-1990s), as well as the simplifying assumptions detailed in sub-chapter 5.1.1.2, have made it a popular subject for radiative transfer models. In this sub-chapter the radiative effects of carbonaceous aerosol components will be discussed in two sections. The first, being more qualitative, will focus on our understanding of the mechanism of the interaction with solar radiation, as well as the on the theoretical background that allows quantification of some of the principal effects, making use of the fundamental radiative transfer parameters discussed previously. At this level, it is neither meaningful nor possible to separate the effects of organic aerosol and BC, and very often the effect of sulfate aerosol has to be also considered. The second part will present some quantitative large-scale or global estimates and model results based on these principles, organized into three sub-sections by the historical trend of available publications. The first calculations and model runs were performed for carbonaceous aerosol from fossil fuel combustion, for which emission inventories were readily available. Biomass burning also received considerable attention, being a major source of carbonaceous aerosol on a global scale. Its emission inventory, however, has been subject to much larger uncertainty, and the properties of smoke particles released by burning of various biomass under various conditions are more diverse and less well characterized. More recently, both source types of carbonaceous aerosol are included in global calculations and radiative transfer models which have the advantage that their predictions can be validated against observational data.

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5.1.3.1 One-dimensional Radiative Transfer Model for the Calculation of Direct Radiative Forcing A box model for the troposphere allowing the estimation of globally averaged aerosol direct radiative forcing was first developed by Charlson et al. [1992]. From this simple model, an equation was derived from radiative transfer equations assuming optically thin aerosol layer (τsc«1) and nonabsorbing aerosol particles. This equation expresses the forcing due to the change in reflectance of the earth-atmosphere system. Although originally applied for purely scattering sulfate aerosol, this equation can equally be used for organic aerosol, as long as it is assumed non-absorbing. The aerosol forcing, ∆FR, is expressed as ∆FR = −

S0 2 2 Tatm (1 − N )(1 − α ) 2 βτ sc 4

(5.7)

where S0 is the solar constant, Tatm is the transmittance of the atmosphere above the aerosol layer, N is the cloud fraction, α is the albedo of the underlying surface, β is the fraction of radiation scattered upwards, τsc is the scattering optical thickness of the aerosol layer. β equals to 0.5 for small particles in the region of Rayleigh scattering, and decreases with increasing size of the particles. In the expression S0/4 represents the globally averaged incident solar flux at the top of the atmosphere, Tatm2 reduces the incident and reflected flux by the transmittance of the atmosphere above the aerosol layer, and N is the fractional cloud cover, in line with the assumption the aerosol effects are significant only in cloud-free areas. Boucher and Anderson [1995] pointed out that this assumption is actually not correct, since aerosol forcing under cloudy sky is about 25 % as efficient. Chýlek and Wong [1995] modified this term to account for the effect of absorbing aerosol:

∆FR = −

[

S0 2 2 Tatm (1 − N )(1 − α ) 2βτ sc − 4ατ abs 4

]

(5.8)

where τabs is the absorption optical thickness of the aerosol layer. For the conservative case of purely scattering aerosol, i.e. when τabs = 0, the formula reduces to Eq. 5.7. It was shown that large error can be introduced into the calculations of top-of-the-atmosphere (TOA) direct forcing when the aerosol with a single-scattering albedo of 0.97 is assumed to be purely scattering (ω0 = 1). The absorbing aerosol component reduces the TOA negative forcing by 30 % at RH = 60 % and by 20 % at RH = 90 % [Pilinis et al., 1995]. Since BC is thought to be the predominant light-absorbing component of fine tropospheric aerosol, it is this modified equation which can adequately

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characterize the direct forcing by carbonaceous aerosol. However, it can also be seen that in this equation the effect of carbonaceous aerosol cannot be separated from those of other aerosol components, which are explicitly included in ω0. Single scattering albedo is typically lower for internal mixtures because all particles absorb solar radiation and not just a fraction of them. A critical value of ω0, ωcrit, indicating a balance between positive and negative forcing, first suggested by Twomey [1977], has been frequently used in the literature. Its value depends on the albedo of the underlying surface and on the boundary conditions of the radiative transfer model in use. Recent estimates of ωcrit are 0.86–0.91 [Hansen et al., 1997], or 0.8–0.9 for clear sky conditions [Haywood and Shine, 1995]. Others have argued that there exists no single value of ωcrit, instead tropospheric aerosol have a net negative and positive forcing at TOA when ω0 > 0.95 and ω0 < 0.85, respectively. In the intermediate range, even the sign of the net effect can vary depending on the albedo of the underlying surface, fractional cloud cover, and vertical distribution of aerosol and clouds [Ramanathan et al., 2001]. Thus it is obvious that absorbing aerosol may even reverse the sign of forcing, causing positive forcing above a certain threshold contribution to aerosol column optical depth. This threshold strongly depends on surface albedo, being 15 % for ocean (albedo 0.06), 10 % for land of an albedo of 0.2, and only 5 % above boundary-layer clouds [Hansen et al., 1997]. In any case, it has been universally accepted that light absorption by aerosol particles does not simply counteract the reduction of incident solar energy by backscattering. A simplistic view behind the concept of balanced forcings is that backscattering reduces solar energy in the Earth-atmosphere system, whereas absorption actually converts it to heat, which warms the system. However, this heating occurs in the atmosphere, thereby reducing surface heating and changing the vertical energy (and temperature) profile in the atmosphere. For example, over the Northern Indian Ocean it might be possible that the absorbed solar radiation by BC, amounting to +14 W m−2, is transported to regions outside the haze layer. Thus the surface experiences cooling in the haze region [Krishnan and Ramanathan, 2002]. In addition, BC can absorb upward solar radiation, which is reflected by the surface and clouds [Haywood and Boucher, 2000]. This effect reduces solar radiation reflected back to space, resulting in a positive aerosol forcing at TOA [Hansen et al., 1997; Jacobson, 2001] but negative forcing at the surface. At the TOA, the absorption by BC counteract the cooling effect of sulfate and organic aerosol, though at the surface all types of aerosol reduces the flux of solar radiation. Therefore aerosol-induced changes on most surfaces far exceed those observed at TOA. This was recently illustrated in the INDOEX campaign, where surface forcing of haze was found to be three and eight times higher than its TOA forcing for clear and cloudy sky,

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respectively [Satheesh and Ramanathan, 2000]. These differences implied that in the presence of highly absorbing aerosol, relatively small values of net TOA forcing can mask a large reduction in surface irradiation. As it was indicated above, surface albedo is a key parameter in determining the radiative effect of tropospheric aerosol. Interestingly, the dependence of the forcing by organic aerosol and BC on surface albedo is just the opposite. Low surface albedo, such as over oceans, enhances the effect of organic aerosol, since most of the solar radiation backscattered by the particles would have been absorbed by the surface. On the contrary, high surface albedo makes absorption by BC (soot) the most effective, since solar radiation reflected from the surface is also absorbed [Haywood and Ramaswamy, 1998]. Therefore BC above ice or snow and above clouds is particularly efficient absorber. Another critical issue is whether the aerosol is externally or internally mixed. Even in the case of externally mixed aerosol the overall effect differs from the sum of the contributions of BC and sulfate aerosol due to the interaction of BC absorption with the scattering by the sulfate particles: for example, the positive forcing of BC becomes more significant as the zenith angle decreases or the surface albedo increases [Schult et al., 1997]. However, internally mixed BC and sulfate seem to result in less negative forcing than those for an external mixture [Haywood and Shine, 1995]. In spite of our ability to explicitly consider the effects of absorbing aerosol in simple radiative transfer calculations, the presence of BC causes additional effects which cannot be handled in a simple radiative model focusing on clear-sky conditions. Under cloudy conditions, when the radiative forcing of scattering aerosol is assumed to be negligible, absorbing aerosol located above clouds can absorb up to three times more solar radiation than in clear-sky conditions [Hansen et al., 1997; Liao and Seinfeld, 1998]. This also implies that soot in the upper troposphere could have more significant climatic effect than its relatively small concentration would suggest. In addition BC, when present inside cloud droplet, may directly affect the shortwave absorption by clouds. Chýlek et al. [1996] derived upper and lower bounds for the mass mixing ratio of BC in cloud water for stratus clouds as 8 × 10−6 and 2.4 × 10−9, respectively. Similarly, for cumulus clouds these values were 8 × 10−7 and 1.2 × 10−10. With an assumed refractive index of 1.75 + 0.44i, the maximum mass absorption efficiency of a small BC particle was 4.4 m2 g−1 [Chýlek et al., 1996]. Regardless of the location of BC particle within the cloud droplet, the extinction efficiency changes less than 0.1 % with respect to that of pure water droplet, which renders the effect of BC on the optical thickness of the cloud negligible. The most realistic approach is to assume a BC particle randomly distributed over the entire volume of the droplet due to its high Brownian motion. With such an assumption the specific absorption of a BC

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particle was calculated to be enhanced by a factor of 2–2.5 relative to that of the same particle suspended in the air [Chýlek et al., 1996]. The increase in the single-scattering co-albedo of the droplet is by a factor of 500 (with respect to that of pure water droplet). Since the absorptivity of an optically thin cloud is proportional to the single-scattering co-albedo, the visible absorption by such a cloud may increase by the same factor.2 This translates into a maximum absorption enhancement over a pure water cloud of about 6 W m−2, equivalent to an annual global average of 1.5 W m−2. It should be noted that these figures were derived from the assumed upper bound of BC mass mixing ratio of 8 × 10−6 in cloud water, while available field observations suggested much lower values (10−8–10−7) [Twohy et al., 1989; Chýlek et al., 1996]. Therefore BC cannot account for the observed anomalous excess absorption of solar radiation by clouds which is in the range of +20–30 W m−2. Unlike other aerosol components, BC can affect atmospheric radiative transfer after it deposited out of the atmosphere under special circumstances. Due to the high surface albedo of fresh snow, even trace concentrations of soot can reduce its albedo and increase the amount of solar radiation absorbed by the surface, and may even enhance snowmelt rates on tundra and sea ice [Clarke and Noone, 1985]. Both the primary effect (decreased surface albedo) and the secondary effect (accelerated snowmelt) can affect regional radiative transfer and thus also climate. Furthermore, deposited soot can become exposed on the surface during snowmelt in summer, long after the springtime Artic haze dissipated from the atmosphere. Calculations showed that a 1 % reduction in the spectrally weighted albedo would require a mass fraction of 40 ng g−1 in fresh snow and only 10 ng g−1 in coarse grained old snow [Warren and Wiscombe, 1985]. Since atmospheric observations show BC mixing ratios well in excess of 10 ng g−1, the expected reduction in weighted albedo is likely to be in the range of a few percent even before the period of accelerated snowmelt [Clarke and Noone, 1985]. The overall scavenging ratio of BC was calculated to be 160, and most BC particles were assumed to be incorporated into snow by wet scavenging. The net gain in energy in the surface-atmosphere system under clear sky conditions was calculated to be 8.19 × 107 J m−2 which included a net loss of 2.95 × 107 J m−2 at the surface. Thus, the overall effect of snow albedo reduction was 6.06 × 107 J m−2, which was approximately equal in sign and magnitude to that of the Arctic haze [Clarke and Noone, 1985]. Jacobson [2002] identified several effects specific to absorbing BC aerosol. In fact, most of these effects were not radiative by nature, but cannot be classified into a single well-established forcing category, so they are discussed here. The daytime stability effect results from the absorption of 2

In the infrared part of the solar spectrum, the absorption by water dominates and the presence of BC particles has a negligible effect.

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solar radiation by BC particles, which reduces the intensity of solar radiation on the surface and increases the vertical stability. As a consequence of enhanced daytime stability, surface winds slow, reducing wind-driven emission of particles and affecting transport of pollutants and energy. At the same time, heating at the top of the boundary layer by absorbing aerosol destabilizes the air above, slightly increasing venting of gaseous compounds and particulates into the free troposphere. The “effect on BC absorption of the first indirect effect” involves the enhancement of atmospheric BC absorption within and above clouds by enhanced cloud scattering due to the first indirect effect. The “semidirect effect” results when the BC-induced solar heating in the boundary layer evaporate some of the clouds and allow more solar radiation to reach the surface [Hansen et al., 2000]. In the tropics and subtropics, due to the emission of highly absorbing particles during the biomass burning season, solar radiation absorbed by the surface can be reduced by as much as 15–35 W m−2, with a corresponding heating of the lower troposphere (up to 3 km) by 60–100 % [Ramanathan et al., 2001]. The “BC-low-cloud positive feedback loop” relates strongly to the semidirect effect through the further enhancement of atmospheric absorption by BC in the boundary layer due to the reduction in low-cloud cover, causing further atmospheric heating and reduction in cloud cover in a positive feedback loop. The “rainout effect” refers to the reduction in cumulus precipitation due to atmospheric absorption by BC, and the resulting increase in the concentration of aerosol particles and gases due to suppressed wet deposition processes. The “BC-water vapor positive feedback” is related to the evaporation of clouds due to the atmospheric heating by BC particles, as well as to the enhanced evaporation rate of water from the surface, since water vapor is a greenhouse gas. The “particle effect through surface albedo” has a component of reducing daytime melting of snow due to reduction in surface irradiation, and another of increasing nighttime melting and sublimation by enhanced downward thermal infrared radiation caused by BC particles [Jacobson, 2002]. As a results of all these complex interactions, the direct relationship between radiative forcing and average global surface temperature may break down for high concentrations of strongly absorbing aerosol [Satheesh and Ramanathan, 2000; Ackermann et al., 2000]. It was long ago when it could be simply calculated that if single scattering albedo of aerosol decreased from 0.95 to 0.75 at a global average optical depth of 0.125, the net radiative effect of the aerosol would change from a cooling of −1.2 °C to a warming of +0.5 °C [Charlock and Sellers, 1980]. Today, as our understanding of these highly complex interactions is still in its infancy, atmospheric radiative transfer

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models rely on very simple parameterization which ignores most of these effects. 5.1.3.2 Radiative Forcing of Carbonaceous Aerosol from Fossil Fuel Combustion The first estimates for the radiative forcing of carbonaceous aerosol were those assessing the modifying effect of BC emitted from fossil fuel combustion on the global clear-sky radiative forcing by sulfate aerosol. These estimates were constrained with the simplifying assumptions that the geographical distribution of fossil-fuel BC follows that of sulfate, and its emission fluxes can also be scaled to those of sulfate. Although in broad terms the source areas and residence times for both species are indeed similar, important differences may arise from the fact that SO2 is mostly emitted by point sources, and converted to sulfate in the atmosphere, whereas soot is emitted directly primarily from mobile sources. By adopting C/SO4 values between 0.05–0.10, Haywood and Shine [1995] found the geographical distribution of the annual mean direct radiative forcing of soot ranging from +0.03 to +0.24 W m−2, depending on the concentration of soot and its mixing state with sulfate. This is comparable in magnitude to the global-mean radiative forcing of sulfate (−0.34 W m−2), which can be calculated using the same model as a base case. The effect of soot is more pronounced on the Northern Hemisphere, where the higher surface albedos made direct aerosol forcing more susceptible to changes in ω0. The result was that hemispheric difference in direct aerosol forcing diminished with increasing soot/sulfate ratio. They found that over areas of high surface albedos the resulting net forcing was positive. Each percent increase in soot/sulfate mass ratio led to a positive forcing increment of +0.01 W m−2 and +0.02 W m−2 for external and internal mixtures, respectively [Haywood and Shine, 1995]. Thus, based on an assumed average BC/SO42− ratio of 0.075, the global mean forcing of BC from fossil-fuel combustion was estimated to be +0.1 W m−2, which was adopted by the Intergovernmental Panel on Climate Change (IPCC) as a very preliminary estimate. Improved estimates which included the effect of BC over clouds, raised its value to +0.2 W m−2 and +0.36 W m−2, depending on the state of mixing [Haywood and Shine, 1997]. These figures still have not included the magnification of the effect of BC by cloud droplets, which could be substantial [Heintzenberg and Wendisch, 1996; Chýlek et al., 1996]. Jacobson [2002] performed model simulations to study the effect of fossil fuel BC and organic aerosol on climate. It was found that removing all fossil fuel BC and organic aerosol cooled global climate by about 0.35 K, which exceeded those resulting from switching off all anthropogenic CH4 and

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CO2 for the forthcoming 100 years and 25–100 years, respectively. Sensitivity studies suggested a range between 0.15 and 0.5 K. The results implied that eliminating all fossil fuel BC + organic aerosol might instantaneously reduce more than 40 % of the net global warming (0.75 K). Most of the globe was warmed by the combined effect of BC + organic aerosol, but not necessarily in the regions where surface BC and OC concentrations were the highest. The major regions of warming were over Eastern Russia, Southeast Asia and Southwestern North America, whereas the major region of cooling was the North Atlantic. To have the same effect anthropogenic CO2 emissions need to be cut by one third, but even then the reduction will occur over a timescale of 50–200 years. Particles have several other effects that the major greenhouse gases do not. Only SO2/SO42− has a greater specific climate response (but opposite in sign) than BC + organic aerosol, since all of its effects lead to cooling, and it also has a greater spatial coverage due to the secondary origin of sulfate particles. It was found that BC warms the air 360,000–480,000 times more effectively per unit mass than does CO2 due to the various effects and feedbacks. The direct forcing by unit mass of organic carbon was found to be −64 W m−2 per g OC when UV absorption by organic compounds was ignored, and −43 W m−2 per g OC when it was accounted for [Jacobson, 2002]. Since a great deal of BC and organic particulates are emitted by diesel vehicles (0.05 g km−1 under current E.U. emission standard, 25–200 times as much as port fuel injection gasoline engines), it was concluded that legislation that favors diesel over gasoline engines on account of their lower specific CO2 emissions, inadvertently promotes global warming due to the enhanced emission of carbonaceous particulates [Jacobson, 2002]. 5.1.3.3 Radiative Forcing of Carbonaceous Aerosol from Biomass Burning The radiative forcing of smoke from biomass burning has long been a subject of scientific research, both on a regional and a global scale. In a strict sense, however, most of these studies do not fit directly into the mainstream of this volume, since they refer to smoke aerosol and not only to their carbonaceous components. It is true that chemically about 70–80 % by mass of smoke aerosol is actually carbonaceous, but the relatively small fraction of inorganic species could exert disproportionately large influence on several of the aerosol properties. Nevertheless, any assessment of the radiative forcing of biomass smoke aerosol could be regarded as a first approximation for the effect of its carbonaceous components. Most recently, there have been specific studies which focus on the effect of carbonaceous species only, in particular of BC, for which measurements have been more readily available.

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More than ten years ago, the global-mean direct radiative forcing ascribed to biomass burning was estimated to be −0.8 W m−2 [Penner et al., 1992]. In the same study, the global effect of absorption by biomass burning aerosol was assessed separately. From an estimated mass absorption efficient for smoke3 of 0.7 m2 g−1, a global annual mean absorption of +0.5 W m−2 was obtained. This atmospheric absorption reduced absorption on the Earth’s surface4 by about −0.3 W m−2. Thus the net heating of the Earth-atmosphere system resulting from the atmospheric absorption by biomass burning smoke aerosol equaled to +0.2 W m−2. The uncertainty of these values was declared to be about a factor of 2. Taking the absorption and scattering optical thickness of a smoke layer τabs = 0.0026 and τscat = 0.030, respectively, the global land and ocean albedo 0.22 and 0.06, respectively, assuming that 80 % of the smoke plume are over the land, S0 = 1370 W m−2 and Tatm = 0.79, N = 0.6, all parameters from Penner et al. [1992], the absorption term was found to reduce the estimated global cooling effect by about +0.17 W m−2. Thus the estimated overall direct effect of biomass burning was in the range of −0.2 W m−2 to −1.1 W m−2, depending on the size distribution of the biomass burning aerosol5 [Chýlek and Wong, 1995]. In a more recent study Hobbs et al. [1997] obtained dry (< 30 % RH) mass scattering efficiency for biomass burning smoke in Brazil in the range of 2.8–3.3 m−2 g−1 for smoke aerosol, 70 % of which consisted of organic compounds. It was also found that mass scattering efficiency increased with the age of the smoke, which was attributed to the growth of the particles by coagulation and further gas-to-particle conversion into the optically more effective size range (median diameter increased from 0.12 ± 0.02 µm to 0.18 ± 0.06 µm). The backscattering factor (β) averaged over the day was found to be 0.24–0.26 (at a wavelength of 550 nm). This value seems to be intermediate between young and aged smoke particles, as for young biomass burning aerosol (r0 = 0.05 µm) β equaled to 0.37, and for the large size end of aged biomass burning aerosol particles (r0 = 0.30 µm) β reduced to 0.11 [Chýlek and Wong, 1995]. The concurrently measured mass absorption efficiencies were in the range of 0.62–0.64 m2 g−1. These values resulted in mean single scattering albedo (ω0) of the dry smoke particles ranging from 0.82–0.84, which was lower than measured by Radke et al. [1988] for boreal fires (0.85 to 0.90). The mean values of fRH(bs) at RH = 80 % (an assumed effective global RH) ranged from 1.1 to 1.4. This

3

Note that this value refers to smoke and not BC. Note that part of the radiation that is actually absorbed in the atmosphere would be reflected by the surface or clouds back to space and would not be absorbed by the surface.

4

5

The absorption term in Eq. 5.8. does not depend explicitly on the assumed aerosol size distribution.

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factor must be applied to the dry mass scattering factor at RH = 80 %, yielding αscat and ω0 of 4.3 m2 g−1 and 0.88, respectively. Using these parameters Hobbs et al. [1997] estimated a global mean direct radiative forcing due to biomass burning aerosol of about −0.3 W m−2, which was a factor of 3 less than that estimated by Penner et al. [1992]. Hobbs et al. [1997] observed optical depths of 0.4 and sometimes as high as 2.5 over large regions in Brazil during the biomass burning seasons, caused by a typical column burden of smoke of about 0.35 g m−2. Under these conditions more sophisticated radiative transfer models would be needed, but beyond any doubt, the direct radiative forcing of biomass burning smoke is substantial on a regional scale. Since biomass burning occurs predominantly in the tropics and is associated with long-range atmospheric transport, its global radiative effects may also be significant. Each year dense smoke plumes can be observed from space downwind of fires in South America between August and October, in Central America in April and May, in South Africa between July and September, in Central Africa between January and March [Liousse et al., 1996]. As an example for long-range transport, smoke from boreal forest fires in Canada was detected as far as over Greece at altitudes of 2–3 km [Formenti et al., 2001], above the boundary-layer clouds. A case study for the evaluation of the radiative forcing of biomass burning smoke was offered recently by the mass forest fires occurred in fall 1997 in Indonesia. These events were triggered by the dry atmospheric conditions associated with a large El Niño event in that year. The magnitude of the fires was so large that they were estimated to contribute to 1–10 % of global biomass burning aerosol emission [Nakajima et al., 1999]. Satellite retrieved optical depth were found to be higher than 0.5 over a fairly large area, extending to the coasts of India and Australia. The aerosol optical depth was especially high in October and November, followed by a rapid decline in early December. The large Ångström exponents found associated with large optical depths indicated the predominance of aged and hygroscopic submicron aerosol. The volume size distribution had a maximum at around 0.25 µm, as retrieved from sky radiances. The single-scattering albedo was high, of about 0.9 during the peak smoke event period, and decreased in time with the reducing intensities of the fires. The high value was slightly above the values reported for forest fire smoke in Amazonia and Africa [Kaufman et al., 1992; Kaufman and Nakajima, 1993; Eck et al., 1998]. The higher ω0 values might be attributed to sulfate emission from simultaneous peat bog fires. These ω0 values correspond to a refractive index of m = 1.5 − 0.005i, consistent with surface measurements.

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5.1.3.4 Global Carbonaceous Aerosol Forcing Estimates Direct radiative forcing is defined as the difference in the net radiation at the top-of-the-atmosphere with the present day and preindustrial aerosol burden. Schult et al. [1997] calculated the direct radiative forcing of BC from the monthly mean BC climatologies of Cooke and Wilson [1996] with a general circulation model. They found the global mean direct forcing to be +1.06 W m−2, distributed between the two hemispheres as +1.70 W m−2 and +0.42 W m−2 for the Northern and Southern Hemisphere, respectively. The geographical distribution of the zonal mean TOA radiative forcing by BC is shown in Figure 5.9.

Figure 5.9. Zonal means of top solar radiative forcing by total BC based on MOGUNTIA calculations for January and July (after Schult et al. [1997]).

As shown in the figure, the peak forcing of +1.8 W m−2 occurred in January at high northern latitudes where the surface albedo was high. Although the maximum optical depth occurred over the savanna regions of Africa, the maximum forcing in the tropical belt (+1 W m−2) can be found farther north over the Sahara where the surface albedo is higher. Local maxima can be as high as +4 W m−2 over the Sahara in July, and +7 W m−2 over the Sahara and Eastern Europe in January. Due to the low surface albedo the effect of BC on the shortwave radiation budget over the ocean was negligible. Considering BC and sulfate as external mixture in model calculations, positive net TOA forcing was found over central Asia and Africa

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in January, which disappeared by July due to lower particulate loadings and the lower surface albedos [Schult et al., 1997]. The global mean normalized direct radiative forcing, i.e. the forcing exerted by unit column mass of BC, was +1.85 W mg−1, but with a very high spatial variation [Haywood and Ramaswamy, 1998]. This quantity indicates how effective BC is radiatively over a given area. Over tropical oceans of low surface albedo its value is almost zero, whereas over ice sheets it can be as high as +5 W mg−1. For comparison, the global mean normalized direct radiative forcing for sulfate is only −0.46 W mg−1, and shows just opposite spatial variations. It might seem surprising that sulfate, which has a mass extinction efficiency of ∼20 m2 g−1 at 92 % relative humidity, more than double that of BC aerosol (9.3 m2 g−1, hygroscopic growth is not considered), is so much less effective than BC in direct radiative forcing. However, this can be very easily understood if we take the simple reflection absorption model of Chýlek and Wong [1995] and assume two hypothetical aerosol with identical optical depth, one for pure scattering, the other for pure absorption, with Rs and β of 0.15 and 0.21, respectively. Then we obtain

DRFabs ≈ 2 × DRFsc

(5.9)

It means that a purely absorbing aerosol will produce a clear-sky direct radiative forcing that is approximately twice as much as that for purely scattering aerosol. The main physical reason behind this is that for scattering, only backscattering reduces the incident solar flux, whereas there is no such dependence for absorption. Another less explicit reason is that while sulfate aerosol exerts significantly reduced direct radiative forcing in cloudy skies, the effect of BC by clouds is even amplified [Haywood and Ramaswamy, 1998]. The fraction of BC that is above highly reflective clouds exerts a much larger effect on the planetary albedo than the other fraction over surfaces of low albedo. Such mechanism can account for a direct radiative forcing of +1.00 W m−2 and +0.27 W m−2 over cloudy regions for the Northern and Southern Hemisphere, respectively, yielding a global contribution from cloudy skies of +0.64 W m−2. This amounts to ∼60 % of the total direct radiative forcing. To put this number into perspective, the contribution from cloudy regions to the total direct radiative forcing of sulfate is only 11 %. Furthermore, the incorporation of BC into cloud droplets, which is not considered in these calculations, can further enhance this effect, making cloudy regions even more important in direct radiative forcing by BC aerosol. The third reason is that direct radiative forcing depends very strongly on the assumed geometric mean radius of the aerosol size distribution. For example, the global mean direct radiative forcing of BC drops only to +0.15

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W m−2 if the same BC climatology is considered with a geometric mean radius of 0.5 µm. Recall that the size distribution of BC is typically shifted to lower radii compared to that of sulfate. Another feature of the direct radiative effect of BC is that it is strongly dependent on the vertical profile of the BC aerosol. For BC present at high altitudes, Tat is higher as there is smaller depth of scattering and absorbing atmosphere above the aerosol layer. But more important is that BC will be located above highly reflective cloud layers, which could increase the direct radiative forcing effect of BC by more than a factor of 10 relative to that in clear sky [Haywood and Shine, 1997]. The magnitude of the increase, however, depends on the cloud optical depth, surface albedo and solar zenith angle. Since sensitivity studies revealed that the direct forcing depends very much on the assumed BC global burden, which is uncertain and being under continuous revision in the literature, Haywood and Ramaswamy [1998] suggested the most probable estimate of direct radiative forcing of BC aerosol of +0.4 W m−2 with a factor of 3 uncertainty. Since normalized direct radiative forcing is free of uncertainties in the BC burden, its range was estimated to be +1.1 to +1.9 W mg−1. The uncertainty of the estimate reflected primarily the uncertainty in the vertical profile of the BC, and to some extent its size distribution, none of which are readily available globally from reliable and standardized measurements. Combination the effect of BC with sulfate in the form of external mixture in the model produces strongly contrasting regions with positive and negative effects, depending on the albedo of the underlying surface. Positive direct radiative effects were seen over central Africa, the Tibetan plateau and high latitudes on both hemispheres with relatively high surface albedo. Overall, the presence of BC reduced the direct radiative forcing exerted by sulfate by ∼50 %. It is likely that the allowing for internal mixing of sulfate and BC aerosol would further increase the relative importance of BC in direct radiative forcing [Haywood and Ramaswamy, 1998]. The reduction of the sulfate negative forcing by internally mixed BC particles is approximately proportional to their mixing ratio, with an average value of +0.034 W m−2 associated with each percent of the BC to sulfate mass mixing ratio [Chýlek et al., 1995]. The direct radiative forcing of organic aerosol has been much less studied that that of BC. In their global model Chung and Seinfeld [2002] found the globally averaged TOA radiative forcing of anthropogenic OC to be only −0.09 W m−2. This is partly due to the relatively low single scattering albedo of OC of ω0 = 0.94 at λ = 550 nm, which toggles the direction of warming or cooling depending on the surface albedo and cloud cover. Including water uptake by OC shifts the tendency toward cooling and doubles the value of the globally averaged annual radiative forcing to −0.18

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W m−2. These values translate into a normalized forcing of OC of −36 to −76 W g−1, which compares to the values given in Cooke et al. [1999] for the radiative effects of fossil fuel OC. Direct comparison can only be made with the predictions by Koch et al. [2001] who assumed externally mixed aerosol with constant mass extinction efficiencies typical of an average relative humidity. The predicted forcing of anthropogenic OC was −0.30 W m−2, which increased to −0.38 W m−2 and −0.64 W m−2 for the year 2030 and 2100, respectively.

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5.2 Hygroscopic Properties of Carbonaceous Particles and their Atmospheric Implications The affinity of various solutes in aerosol particles for water is a key factor controlling the mass and size of the particles, and consequently their scattering and absorption efficiencies, their acidity and rates of multiphase reactions, and—perhaps the most importantly—their ability to act as cloud condensation nuclei. The hygroscopic behavior of pure inorganic compounds making up atmospheric aerosol is thermodynamically well-understood, yet discrepancies arise when one tries to apply these principles to observations of atmospheric aerosol. Most of these discrepancies result from the presence of organic compounds in aerosol, whose hygroscopic properties could be almost as diverse as their chemical structures. Theoretical and laboratory studies on the hygroscopic properties of a few pure organic aerosol constituents have only recently become available, but the understanding of the complex interactions between various organic and/or inorganic species in the atmosphere has advanced little. Basically, organic aerosol can be classified into a hydrophobic fraction, such as primary aerosol from fossil fuel combustion, consisting of highmolecular weight alkanes, alkanoic acids, alkenoic acids, aldehydes and ketones. These molecules often exhibit surfactant behavior, particularly if the acidic group is neutralized by an alkaline ion; and into a hydrophilic fraction which is formed by oxidation of gases and particles in the atmosphere, and to a smaller extent, during combustion. This fraction is made up of lower molecular weight carboxylic acids, dicarboxylic acids, alcohols, nitrates and polyfunctional compounds, which are non-volatile and water-soluble, consequently they likely interact with water. This chapter, which is intended to describe the interaction of carbonaceous aerosol with water, is arbitrarily divided into sub-chapters on hygroscopic growth, CCN-activity, ice-nucleating ability, and on the possible consequences of all these effects on the Earth’s climate. Albeit the fundamental principles underlying all of these processes are very much the same, these effects are important in quite distinct atmospheric phenomena. Hygroscopic growth characterizes the uptake of water at relative humidities below 100 %, the resulting change in physical and chemical properties of the particles, and its effects on atmospheric phenomena such as radiative transfer. From the climate’s point of view, hygroscopic growth is vital for understanding the scattering and absorption of solar radiation by tropospheric aerosol. Within the troposphere this property is more critical in the boundary

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layer, to which most of the aerosol optical depth is confined, especially in marine environment where RH is generally high. On the other hand, cloud condensation studies represent quite a distinct field within atmospheric science, aimed at a better understanding of the nature of clouds, these key elements of the Earth’s atmosphere, which play a vital role in virtually all atmospheric processes. Within the atmosphere, cloud formation typically occurs at higher altitudes, and since it relies on atmospheric aerosol as well as a number of physical processes, it is much less understood than the hygroscopic growth of aerosol particles. In general, our understanding on ice nucleation is at best rudimentary, so we cannot expect that the role of carbonaceous aerosol is well characterized. Nonetheless, there are some studies on ice nucleation on soot particles, which are developed to the extent that even its climatic implications are hypothesized. Each sub-chapter starts with a summary of atmospheric observations regarding the possible role of carbonaceous aerosol, followed by laboratory studies on the effect of soot and pure organic species of atmospheric relevance. The complex interplay between inorganic aerosol and organic species is treated separately. While the former approach provides a sort of theoretical understanding of the type of interaction between carbonaceous species and water, as well as some input for models treating aerosol as an external mixture, the latter is more relevant in the case of internally mixed aerosol. Both approaches rely on the well-established thermodynamic relationships between inorganic solutes and water, and try hard to extend the validity of such relationships to the increasingly complex case of mixing with organic species, sometimes with little success. As a result, the last sub-chapter, which summarizes the postulated climatic effects of carbonaceous aerosol through their effects on cloud formation, is more speculative rather than conclusive. This is not surprising in the light of the fact that such indirect effects are the least understood among all components of climate forcing even for sulfate aerosol. Nonetheless, some challenging hypotheses on the role of carbonaceous aerosol in inadvertent climate modification are presented. Only the indirect effects are considered here: direct effects are treated in sub-chapter 5.1, where aerosol optical properties are discussed.

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5.2.1 Hygroscopic Growth Attributed to Carbonaceous Aerosol The thermodynamic basis for the interaction between inorganic solutes and water as a function of relative humidity below RH=100 % is well understood. Solutes can be classified as deliquescent or hygroscopic, depending on the type of interaction that occurs between the solute and water. If a dry, deliquescent particle is exposed to increasing relative humidity, no change can be observed until the deliquescence humidity of the particle is reached. At that point the particle experiences a sudden increase in mass and transformation of the solid particle into a solution droplet. Upon decreasing humidity, a hysteresis effect occurs and the droplet crystallizes at a much lower RH than its deliquescence humidity. Below the deliquescence point, the particle only adsorbs a thin water layer on its surface, which remains saturated with respect to the solid particle. At the deliquescence RH, solid particle begins to dissolve into the water layer in an unstable situation, thus lowering the vapor pressure over the surface. According to Raoult’s law, reduced vapor pressure favors the condensation of water, which in turn allows more of the particle to dissolve into the aqueous phase. This cycle continues until the solid particle is completely dissolved. Salts such as NaCl or (NH4)2SO4 are good examples of deliquescent particles. On the other hand, a hygroscopic particle will grow monotonically as water sorbs onto its surface. In contrast to deliquescent particle growth, no sudden change of mass can be observed. Furthermore, this type of hygroscopic behavior occurs without hysteresis. In the atmosphere, sulfuric acid particles can be representatives of this class. This simple classification gets more complicated already when several inorganic solutes are mixed within single particles. The situation becomes even worse when the effect of organic species is included. For example, surface active compounds may form an organic film on the surface of inorganic particles and might thus significantly alter the water uptake by the particles. 5.2.1.1 Atmospheric Observations of Water Uptake Apportioned to Organic Aerosol The water content of atmospheric particles was determined as a difference between particle volumes up to an RH of 80–85 % and in dry air [Saxena et al., 1995]. The approach was to compare the total water content of the particles with the water content expected to be associated with the inorganic ionic fraction only, as obtained from simultaneous impactor measurements of particle chemical composition and subsequent calculations

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based on literature data. For Arizona desert aerosol the measured water content was found to be a factor of 1.78 higher for RH=85 % on average than that estimated for inorganic ions only. Large excess of water appeared to be associated with a large organic mass fraction, though the relationship scattered over a very broad range. In addition, relatively more excess water was found for particles of 0.4–0.5 µm than for those of 0.1–0.2 µm. This finding implies that the composition of organic matter varies with particle size, and compounds with higher water affinity tend to be associated with larger particles. The median and average values of excess water mass per unit organic mass were 1.1 and 1.5, respectively, at RH=85 %. Similar results were obtained by in-situ tandem differential mobility analyzer (TDMA) measurements in the Great Smoky Mountain National Park [Dick et al., 2000]. The ambient particles were first dried at ∼5 % RH then their hygroscopic growth was measured in the RH range of 5–85 %. For the model calculations the particles were assumed to be homogeneous mixtures of ammonium-sulfate, organic matter1, elemental carbon, and water. Below RH < 40 %, finite particle growth was measured which could not be predicted neither for crystalline sulfate nor for supersaturated sulfate solution. Similarly, above RH > 70 % significantly more water was absorbed than predicted for the inorganic salts alone. The excess water, i.e. the difference between the measured and calculated water volume ratios at RH=85 % strongly correlated with the organic mass fraction of the particles, as shown in Figure 5.10. The derived estimates for the water uptake by organic compounds, expressed in the ratio of the volume of water to the volume dry particles, are compared with those by several sulfate species in Figure 5.11. The upper and lower ranges of the estimates are given as “high organics” and “low organics”, respectively. It can be seen that on a volume basis, organic species absorb significantly less water than any of the sulfate species. On a mass basis, however, the higher estimate for organic hydration approaches that of both ammonium bisulfate and ammonium sulfate up to RH ∼60 % as a consequence of the lower density of carbonaceous matter. Incorporation of water associated with organic species significantly improves the agreement between the model and observations, with the exception of particles of 0.05 µm. It might well be that the organic composition of these small particles differs from the bulk, actually the organic compounds there might be more hygroscopic, though there were no data to support this hypothesis. Indirect support may be obtained from the analysis of size-resolved aerosol from the same site for functional groups and chemical bonds using FTIR spectroscopy 1 For the calculation of organic matter from organic carbon a mass conversion factor of 2.1 was assumed.

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[Blando et al., 1998]. In the smallest particles, a predominance of carbonyl and organosulfur moieties was found, indicative of the polar character of the nucleation mode organic aerosol. It should be noted, however, that mass concentrations derived from filter samples are about a factor of 2 higher than those obtained from the impactor data, which points to possible adsorption artifacts.

Figure 5.10. Excess water (νwt−νwl) at 85 % RH as a function of the organic fraction of the total mass. The assumed density of the organic compounds is 1.4 g cm−3 (after Dick et al. [2000]).

In contrast, aerosol collected at an urban site (in Claremont, at the northeastern edge of the Los Angeles urban area) showed a dichotomy in hygroscopic behavior, as two distinct types of particles were detected by the Tandem Differential Mobility Analyzer (TDMA), especially above Dp > 2 µm [Saxena et al., 1995]. This finding corroborates previous TDMA measurements showing that atmospheric aerosol particles of a given size at low (∼40 %) relative humidity can be classified into an external mixture of a “more hygroscopic” and a “less hygroscopic” fraction when humidified to RH∼90 % [Zhang et al., 1993]. By studying the elemental composition and morphology of urban aerosol particles separated by hygroscopicity McMurry et al. [1996] found that the “less hygroscopic” particles consisted mostly of carbon in the form of chain agglomerate. About one-third of these particles were irregular in shape, and less than one-tenth were round or flake-like. In contrast, the “more hygroscopic” particles in general appeared to be homogeneous spheres, consisting of sulfur and oxygen, presumably having

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lost some material in the high vacuum and under the electron beam. Occasionally, carbon was found in the “more hygroscopic” particles, but its observed abundance was much less than that reported previously by Zhang et al. [1993], who found that 30–40 % of particulate organic carbon associated with the “more hygroscopic” fraction in Los Angeles and at the Grand Canyon.

Figure 5.11. Comparison of derived organics hydration, in terms of volume, with hydration of sulfuric acid, ammonium bisulfate, and ammonium sulfate (after Dick et al. [2000]).

The total water content of all particles in the Los Angeles urban area was determined simply as the number-weighted sum of the two modes [Saxena et al., 1995]. The water content thus observed was inferior to that calculated from the inorganic species alone (on average 66–75 % of it). Therefore it seems that in an urban location organic compounds tend to diminish water uptake rather than enhance it. In principle, it might also happen that there is insufficient time to reach equilibrium in the measuring system (with a characteristic time constant of 1–5 s). In the case when an organic film is present on the particles, it may impede water uptake and result in underestimation of the water content. This scenario, however, may not be relevant in the atmosphere, where usually sufficient time is allowed for the particles to reach equilibrium. On the basis of such studies it cannot be

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decided whether organic compounds indeed reduce equilibrium water content of the particles or simply retard water condensation, or both. 5.2.1.2 Laboratory and Modeling Studies a) Water uptake by pure organic compounds of atmospheric relevance Hydrophilic organic species themselves are able to interact with water, in a manner similar to pure inorganic compounds. As a rule of thumb, organic compounds which are completely miscible with water, like acetic acid, exhibit hygroscopic behavior similar to sulfuric or nitric acids. On the other hand, organic species which have finite solubility, such as oxalic acid, deliquesce at a given relative humidity. However, the hygroscopic behavior of pure organic compounds cannot be relevant in determining the hygroscopic behavior of atmospheric aerosol particles. The problem is that none of these organic species form individual aerosol particles; they are not even major components of organic aerosol. Furthermore, their properties are markedly dependent on a number of factors, such as pH or ionic strength, and they readily interact with one another as well with inorganic solutes. In addition, some of these species are volatile, consequently their temperature-dependent partitioning between the gas and aerosol phase must be taken into account. Consequently, the hygroscopic behavior of the multitude of organic species forming organic aerosol will not be treated in depth: only properties of a few representative compounds will be described here for illustrating the complexity of the problem. It is well-established that the physico-chemical properties of low molecular weight dicarboxylic acids, such as solubility, melting and boiling point, exhibit a strong alternation with the number of carbon atoms, known as an odd-even effect. Whereas odd carbon number members such as malonic and glutaric acid can deliquesce below RH=100 %, even members such as succinic or adipic acids are only slightly soluble and cannot induce apparent deliquescence below saturation. Several odd carbon-number dicarboxylic acids have higher solubilities in water than ammonium sulfate (e.g. malonic, glutaric, and malic acid, with solubilities 161 g, 116 g, and 145 g in 100 g H2O, respectively)2. Consequently their hygroscopic behavior is very similar to that of ammonium sulfate, obeying the traditional form of Köhler theory. On the other hand, the hygroscopic behavior of the even members, particularly at the initial stage of condensational growth, cannot be described

2

The solubility of ammonium sulfate is 75.7 g in 100 g H2O at 20 °C.

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by the traditional Köhler theory which assumes complete dissolution of solutes in the droplets all the way through its growth [Hori et al., 2003]. However, the salt forms of organic acids exhibit completely different hygroscopic behavior: sodium salts of formic, acetic, succinic, pyruvic and methanesulfonic acid deliquesce already at low relative humidity (e.g. ∼44 % for sodium acetate), whereas sodium malonate and sodium maleate absorb water at all relative humidities and evaporate without a hysteresis [Peng and Chan, 2001]. In contrast, sodium and ammonium oxalate particles do not deliquesce even at RH = 90 %. Compared to studies on pure organic compounds, a recent work on the uptake of water by different organic aerosol particle types—wood smoke, Diesel soot and secondary aerosol from α-pinene-ozone oxidation—can be more relevant in the atmosphere [Jang and Kamens, 1998]. It should be noted, however, that these aerosol types, especially the first two, may contain inorganic inclusions which have an unproportional impact on the hygroscopic behavior of the particles. Based on the earlier work by Vartiainen [1994], who found linear relationship between water uptake (mg) and relative humidity, the proportionality constant was determined for the three aerosol types. The activity coefficient of water in the organic aerosol phase, wγom, was determined from thermodynamic calculations and extrapolated to infinite dilution (wγ∞, Henry’s law range). For wood smoke particles, wγ∞ was close to unity, implying that absorbed water in these particles is homogeneously miscible with the hydrophilic organic aerosol medium. The water activity coefficient for αpinene-O3 aerosol was found to be 2.4, indicating that water had less preference for secondary aerosol than for wood smoke. With an RH of 99 % at 25 °C, the water uptake can be estimated only 3.9 % (m/m) of organic matter, below its calculated saturation solubility. The theoretically calculated water activity coefficient for diesel soot, γ was 32, reflecting the extremely hydrophobic nature of the particles. This would lead to phase separation between the water and the organic layer of soot particles if significant water absorption occurred beyond the saturation point. A recent hygroscopicity TDMA study has recently showed that organic matter isolated from the aqueous extract of rural fine aerosol (HULIS) deliquesced between 30 and 60 % relative humidity, and exhibited hygroscopic growth factors (at 90 % RH) in the range of 1.08 to 1.17 [Gysel et al., 2004]. These growth factors were of similar magnitude that those of SOA in smog chamber experiments, but were lower than those of highly soluble organic acids. When compared to the behavior of humic and fulvic acids, deliquescence occurred at lower RHs and hygroscopic growth factors were slightly higher in the case of HULIS.

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b) Altering the hygroscopic behavior of inorganic aerosol by organic compounds One mechanism for the modification of water uptake of inorganic particles by the organic species is that an organic film forms on the surface of the particles which is expected to impede mass transfer between the gas and particulate phase. There is some controversy in the literature whether the effect is retardation only or also a shift in equilibrium. Dynamic TDMA experiments cannot differentiate between the two effects, since the time constant of the measurements is usually small (5–10 s). The hygroscopic growth rates of sulfuric acid particles were studied by TDMA in the size range of 40–120 nm, coated with lauric, stearic or oleic acid in thickness varying from 2.2 to 30 nm, in the RH range of 20–85 % [Xiong et al., 1998]. Even a monolayer of saturated fatty acids was found to retard hygroscopic growth. As an example, at an RH of 50 %, the average growth factor of H2SO4 particles coated with a monolayer of lauric acid was reduced to 90 % of the values of the uncoated H2SO4 particles over a time scale of 6 seconds. The reduction in growth factor appeared to continue with an increasing number of layers up to six, where the growth factor became independent of the number of layers of the coating. On the contrary, oleic acid, which is expected to form a fluid surface film on droplets and strongly impede mass transfer, showed no observable retardation for a monolayer and much weaker effect for thick multilayers. Nevertheless, the film retardation effects on the particle hygroscopic growth are expected to diminish at higher relative humidity with particle hygroscopic growth, as the film may become thinner or broken at larger sizes. In static experiments, giant NaCl particles (D∼30 µm) coated with a model surfactant (Tween 80, a polyoxyethylene sorbitan ester and nonionic surfactant) showed a decrease in relative mass gain at the deliquescence humidity from 4.1 to 1.8 (at 3:1 salt/surfactant molar ratio) [Andrews and Larson, 1993]. This finding may imply that organic surfactants limit the amount of water uptake in equilibrium. Another mechanism can be that the surfactant coating reduces the deliquescence relative humidity of the particles. It is known for single component particles that the deliquescence humidity decreases as the activity of the solute increases. Therefore the addition of surfactant results in an increase in solute activity via the surface energy of the particles, which corresponds to a decrease in deliquescence relative humidity. For example, organic surfactants were found to slightly reduce the deliquescence RH of NaCl particles from 75 % to 70–73 % [Andrews and Larson, 1993]. It was also demonstrated experimentally that water-soluble dicarboxylic acids shift the deliquescence of ammonium sulfate to lower relative humidity [Brooks et al., 2002].

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On the other hand, recent studies imply that in atmospherically relevant concentrations organic species may not have significant impact on the hygroscopic behavior of inorganic particles. For example, the effect of tetracosane, octanoic acid and lauric acid on the hygroscopic properties of NaCl aerosol particles was studied at relative humidities between 30 and 95 % [Hansson et al., 1998]. The initial dry particle diameter was 100 nm, which was then coated with variable mass of organic compounds. When the mass fractions of the organic species were low (e.g. up to 5 % for tetracosane and 12 % for octanoic acid), the particles exhibited the same size dependence on RH as if they were pure NaCl. The differences became observable only at much higher organic mass fraction. These included a shift in the RH of deliquescence to lower values (from 75 % to 72 %), and a reduction in the growth factor by 22 % at RH=80 %. These values refer to tetracosane, which turned out to be the most effective coating suppressing the uptake of water among the compounds studied. However, the results implied that even if a salt particle contained as much as 50 % by mass surface-active organic compound, it can still deliquesce and grow with increasing relative humidity. c) Water uptake by soot (BC) particles Hygroscopic growth of soot particles is of particular importance in atmospheric science since hydration—well in excess of the increase in particle size—may result in a dramatic increase in the specific absorption of the particles (see sub-chapter 5.1). Hydration also changes the surface characteristics from non-polar to highly polar, thereby influencing adsorption of semi-volatile species and the rate of heterogeneous reactions. The general understanding that is reflected in atmospheric models is that BC is hydrophobic and does not take up water even at RH approaching 100 %. Indeed, observations with carbon black, a surrogate for atmospheric black carbon, have shown no noticeable uptake of water at RHs between 15 and 85 %. However, it was shown that formation of even a thin organic surfactant film on carbon black surface can render it hygroscopic [Andrews and Larson, 1993]. For sodium dodecyl sulfate, the observed mass gain can be 10–20 %, which is somewhat surprising as pure dodecyl sulfate particles do not show any tendency to absorb water. This can be explained by a possible micellelike structure in mixed solutions with carbon black: the hydrophobic end of the dodecyl sulfate attaches to the hydrophobic carbon black, leaving the hydrophilic end to face outward.

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If, however, freshly generated soot is considered instead of carbon black, at least partial hydration of its surface can be observed, as evidenced by its infrared spectrum as a hydrogen-bonded OH band or thermal desorption of water at ∼700 °C from surface functionalities [Akhter et al., 1985]. The hydration of freshly formed and reacted soot was measured with an electrobalance between relative humidities 33 and 52 % [Chughtai et al., 1996]. The results are shown in Figure 5.12.

Figure 5.12. Hydration of soot and increase in P/P0 as a function of time (after Chughtai et al. [1996]).

As it can be seen in the figure, in the initial stage of the isotherms, especially between RHs 30 and 40 %, an activated adsorption process occurs, with the limiting adsorption and the corresponding surface coverage increase in the order of fresh soot < sulfated soot < ozonized soot < nitrated soot. Chemisorption of water consumes about 20 % of the available surface on fresh soot and as much as 52 % on nitrated soot. The results imply that about 40 % of the surface oxygen functionalities are hydrolyzable, most likely carboxylic groups, the remainders can be aromatic-conjugated carbonyls or C–O–C linkages. The experiments can also be extended in time to the point where no further mass gain can be observed; for fresh soot, this occurs when 231 µg of water adsorbed on 20 mg of fresh soot. The water uptake of soot also depends on aging, with larger differences observed over the first few days. If fresh nhexane soot is exposed to pure O2 for 24 hours, significant increase in its water uptake can be observed. Conversely, evacuation of fresh soot for 12

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hours reduces water adsorption. These observations imply that physisorbed (as well as chemisorbed) O2 plays a key role in the adsorption of H2O by hydrogen bonding [Chughtai et al., 1996]. Since dispersion interactions between graphite and water molecules are weak, water adsorption depends strongly on the presence of hydrophilic primary centers. The adsorption isotherm of water on kerosene soot is shown in Figure 5.13 [Ferry et al., 2002].

Figure 5.13. Adsorption isotherm of water on kerosene soot at 295 K and on hydrophobic graphitized soots, graphon, at 291 K taken from Young et al. [1954] (after Ferry et al. [2002]).

It is likely that the active sites in the micropores of kerosene soot can take up water already at low RHs. Between RHs of 10 and 70 %, clusters of water molecules bridge the opposite sites in the pore wall and the supermicropores are filled with a liquid-like adsorbate phase. This process determines the limiting adsorption value indicated by the form of the isotherm. Above p/ps > 0.9 (RH > 90 %) capillary condensation starts in the mesopores and on the external soot surface. At saturation, the fraction of water adsorbed in the micropores was 12 %, in the supermicropores was 39 % and 14 % in the mesopores, whereas 35 % of the total was adsorbed on the external soot surface [Ferry et al., 2002]. Interestingly, water cannot freeze in the micropores, and even at temperature as low as 204 K 15 % of the water was still in supercooled state, since a regular H-network is unable to form in a confined space below 2 nm in diameter.

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At low RH prevailing in the fresh plume, water molecules adsorb on the primary active centers and fill in the soot micropores. As the plume cools and RH approaches 70–80 % the supermicropores of soot become completely filled. Upon saturation, which normally occurs around 240 K, capillary condensation takes place in the soot mesopores (>2 nm) as well as multilayer growth takes place on the external soot surface [Ferry et al., 2002]. Overall, it seems that at least some hygroscopic growth should be attributed to soot particles, with all of its consequences on their optical and surface properties (see sub-chapter 5.1 and 5.3). 5.2.2 Cloud Condensation Nucleus (CCN) Activity of Carbonaceous Aerosol Traditional Köhler theory describes the equilibrium vapor pressure of water over an aqueous droplet. The simplest algebraic form of this theory for dilute aqueous droplets:

Sw ≈

a r

−

b r3

(5.10)

where Sw is the supersaturation of water vapor, r is the droplet radius, a and b are constants, incorporating the influence of the curvature of the droplet (“Kelvin-effect”) and the vapor pressure depression caused by the solute (“Raoult-effect”), respectively. A droplet is activated at the size rc where Sw is at its maximum value, Sc. Above that the droplet can grow spontaneously if ambient supersaturation is maintained at or above the Sw. Below the size of rc, the particle is said to be a haze droplet. The traditional Köhler theory constrains that the growing droplet initially exists as an aqueous solution but with all available solutes already in the solution. These are reasonable assumptions if the cloud condensation nucleus consists of highly soluble inorganic salts such as NaCl or (NH4)2SO4 or highly soluble liquid-like H2SO4, none of which have marked effect on the surface tension. It should be noted that the aerosol particles that serve as CCN also participate in the atmospheric reprocessing of aerosol through multiphase aqueous chemistry leading to the formation of secondary species and droplet coalescence. The water activity coefficient, γw, is usually taken as unity, because the solutions at activation are sufficiently dilute due to either the limited solubility of the organics or to the large water uptake by hydrophilic substances [Raymond and Pandis, 2002]. For most cloud physics application it was previously thought to be sufficient to assume a constant van’t Hoff factor equal to the value at infinite dilution in spite of the fact that the actual

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van’t Hoff factor was shown to depart from this value even in ammonium sulfate solutions already at concentrations of 10−4 M [Young and Warren, 1992]. Later it was clearly shown that this simplifying assumption is not valid, in other terms the limiting value of the van’t Hoff factor should not be used in the Köhler equation around the activation radius [Konopka, 1996]. Unfortunately, this warning went unnoticed in subsequent publications and most authors still keep using values at infinite dilution for slightly soluble multifunctional organic species. When the solubility of the species is low enough, the Raoult term in the Köhler equation becomes negligible, and supersaturation is determined by the Kelvin term. It has been demonstrated that pure organic species may act as CCN, but it has been a subject of controversy in the literature whether organic coatings promote or hinder the activation of inorganic particles. It is particularly important whether organic compounds which do not form CCN in themselves can render active nuclei inactive, can delay their activation, or act only as inert mass. Depending on its chemical and hygroscopic properties, an organic compound may be part of the inactive core in a growing droplet, dissolve gradually into the aqueous phase or be water-soluble. It can either be distributed evenly in the droplet phase or accumulate on the surface. Transfer between the gas and droplet phase is also possible, especially in the case of semi-volatile and water-soluble compounds. 5.2.2.1 Atmospheric Observations of CCN Apportioned to Organic Aerosol Relationships between CCN number concentration and chemical species are usually derived from measured mass concentrations, though in fact most CCN are too small to contribute to even submicron mass concentrations. Since sulfate is clearly not the only component, and sometimes not even the dominant component of submicron aerosol mass, it can be expected that other species, most likely organic compounds also contribute to the CCN concentrations. The earliest observations of this kind were made for biomass smoke particles. For example, smoke from sugar cane fires was found to be a prolific source of CCN, which rather efficiently increased the cloud droplet number concentrations downwind [Warner and Twomey, 1967]. On the other hand, measurements of CCN concentration enhancement downwind of a simulated forest fire indicated a significant increase in relatively inefficient CCN [Hobbs and Radke, 1969]. In other terms, the number of CCN active at 1 % supersaturation increased considerably, but of those active at 0.2 % supersaturation remained unchanged.

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Carbonaceous Aerosol

Since biomass smoke particles contain variable amounts of organic and inorganic constituents, it is unclear whether the CCN activity of organic smoke particles can be attributed to their association with water soluble inorganic species, or their organic component itself is intrinsically CCN active [Hallett et al., 1989]. The former possibility stems from the ability of inorganic species to lower supersaturation. For example, for a dry particle with a diameter of 100 nm, only about 0.3 % (!) by mass of (NH4)2SO4 is sufficient to allow activation at a supersaturation of 1 % [Rogers et al., 1991a]. Possible contributors to the latter may be low-molecular weight organic anions (formate, acetate or oxalate) which were identified in haze layers associated with biomass burning in Amazonia [Andreae et al., 1988], or highly soluble other polar organic species which have critical supersaturation similar to that of ammonium sulfate [Rogers et al., 1991a]. A comparison of the measured CCN concentrations and those calculated from measured size distributions assuming soluble inorganic composition did not infer significant contribution of organic species for oceanic aerosol [Bigg, 1986]. On the contrary, it was found that when the particle concentration exceeded 300 cm−3, the measured CCN concentrations were considerably less than expected. It was assumed that the condensation of organic vapors onto the particle surface forms a surface active coating that retards the uptake of water by the droplet. This effect may cause that cloud formation can be extensively delayed and fully suppressed at low supersaturation. The first explicit indication for the significance of organic aerosol as CCN came from the simultaneous determination of the mass size distribution of nss-SO42− and organic carbon, CN and CCN concentrations at the El Yunque summit in Puerto Rico [Novakov and Penner, 1993]. The results showed that nss-SO42- was associated with a relatively narrow range of fine particles, whereas the OC mass size distribution was more evenly spread over the entire size range extending down to the lowest measured size-cut. This caused that organic particles contributed more to the number concentration of particles below 0.08 µm, in spite of their lower total mass concentration compared to that of nss-SO42−. The average measured CCN concentration (556 cm−3) agreed remarkably well with the sum (578 cm−3) of N0.05(S) and N0.05(C) (total number concentrations of sulfate and organic particles above diameter of 0.05 µm, respectively, as derived from the corresponding mass size distributions assuming externally mixed particles). They also showed that the derived number concentrations did not depend critically on the assumed mixing state, with the assumption of internal mixing the estimated values decreased by less than 10 %. An interesting and much cited conclusion of the authors was that only ∼37 % of the measured CCN concentration can be attributed to nss-SO42−, and the remainder was due to organic aerosol.

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At the time it appeared, this statement was revolutionary in the sense that until then CCN activity had been thought to be explained by the presence of sulfate aerosol. Since the formation of organic aerosol, unlike sulfate, was traditionally assumed not to involve aqueous chemistry, stronger relationship was expected between the mass flux of organic aerosol and aerosol number concentration than in the case of sulfate. The individual contributions of mass concentration of major aerosol species to total CCN concentration were apportioned on the basis of the assertion that CCN concentrations can be estimated from the dry aerosol number size distribution provided that the relationship between the dry size and supersaturation of aerosol particles is known [Rivera-Carpio et al., 1996]. The mass size distributions of major aerosol components, aerosol number size distribution and CCN concentrations (at 0.5 % supersaturation) were measured concurrently. This approach is strictly valid only if the major species are externally mixed, otherwise it can only be regarded as reasonable estimates of the individual contributions. The key assumption was that the CCN concentration can be approximated with the total number concentration of the particles (N > Dc), having dry diameters greater than the critical diameter Dc. It can be readily determined from the cumulative number size distribution N(D), which relates to the aerosol number size distribution through the relationship:

N ( D) = ∫ n( D)d log D

(5.11)

A value of 48 nm was chosen for Dc corresponding to the Dc value of (NH4)2SO4 at a supersaturation of 0.5 %. The choice was confirmed by measurements conducted at a coastal site in California, in which the average CCN/N(>0.048) ratio was found to be 0.98 ± 0.25. This agreement also indicates that the nucleation properties of the multicomponent aerosol are comparable to that of ammonium sulfate or other water-soluble species. The mass contributions of the three major aerosol species were estimated from their cumulative number size distributions. Their findings are summarized in Table 5.1, comparing the measured and estimated CCN concentrations and the percent mass contribution of major species to total CCN.3 The results showed that the measured and derived distributions agreed within about 15 %, but the contributions of individual species were highly variable. SO42− contributed in a range of about 20–65 %, while organic aerosol had an even wider range between 4 and 80 %. The reason for such a high variability lies in the features of the individual mass size distributions. 3

Note that OC* denotes organic matter calculated as 1.2 times the mass of organic carbon.

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Both the SO42− and organic mass size distributions were in all cases bimodal with peaks centered at ∼0.1 and ∼0.3 µm, though with variable mass fractions between the two modes. The relative and absolute contributions of aerosol species to CCN concentration depend primarily on their mass concentration in the first mode, and certainly not on their total mass. Although these calculations were performed for a supersaturation of 0.5 %, similar conclusions applied if the results were recalculated for lower supersaturations (down to 0.1 %) more typical of stratiform clouds. Table 5.1. Measured and estimated total cloud condensation nuclei (CCN) concentration and percent contributions of sulfate, organic, and NaCl aerosols to estimated total CCN (after Rivera-Carpio et al. [1996]). Date

CCN, cm−3

Meas/Est. SO42-,% OC*,% NaCl,% BC(%)

Measured

Estimated

June 27–28, 1994

419

380±117

1.13

54

34

12

---

June 28–29, 1994

229

270±130

0.84

29

62

9

---

July 5–6, 1994

333

308±45

1.08

64

4

31

---

Oct. 20–21, 1993

808

883±113

0.91

28

68

---

4

Oct. 22–23, 1993

506

641±109

0.79

19

78

---

3

Where there are no values given, concentrations were below quantifiable limits.

Simultaneous measurements of aerosol chemical composition and CCN concentration in the Northwest Pacific Ocean, in the lee of the Eastern Asian pollution plume revealed considerable number concentrations of CCN (at 0.5 and 1 % supersaturation) even when nss-SO42− concentrations were nearly zero [Matsumoto et al., 1997]. Under these conditions, with nss-SO42− concentrations below 0.5 µg m−3 and practically constant, significant correlation was found between CCN and water-soluble organic carbon mass concentration. These results suggested that organic compounds actually contributed to CCN formation. Of the two possible explanations for the apparent ability of organic aerosol to serve as CCN, these findings indirectly implied the complex organic material itself is hydrophilic and is therefore CCN active.

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5.2.2.2 Laboratory and Modeling Studies Atmospheric CCN formation was modeled with a view to assessing the role of organic compounds [Kerminen, 2001]. The model lumped organic species into only two groups by their volatility. The first group included all non-volatile species, while the second was restricted to semi-volatile species with Kom in the range of 0.01–1 m3 µg−1. The soluble fraction of organic matter was set to 0.15, and constant production rates were assumed for both groups of lumped organic compounds. 24 hours were allowed for the initial particle population to develop. It was found that both the gas phase production rate and hygroscopicity of the organic matter are important in producing CCN. Apart from the rather unlikely case that the organic compounds are all non-hygroscopic and SO2 mixing ratio is very low, even slightly hygroscopic (ε > 0.15) organic species can produce CCN with efficiency proportional to their gas phase production rate. The capability of non-volatile organic species is much higher than that of semi-volatile compounds, which may be able to assist in CCN production but are unable to create new CCN by themselves. In case a non-volatile substance has too low production rate (<0.1 µg m−3 d−1), its scavenging by preexisting particles and cloud droplets prevents its mixing ratio to build up between two successive cloud events, and its contribution to CCN production remains negligible. The role of non-volatile organic species becomes comparable to sulfate at an atmospheric production rate of about 0.1–0.5 µg m−3 d−1, depending on the assumed initial SO2 mixing ratio and variations in updraft velocities in successive cloud events. It is also clear that the bulk chemistry of the fine aerosol is insufficient to reveal whether the production of new CCN is dominated by sulfate or organic matter. It is actually the chemistry of the dry particle size range between 50 and 80 nm which is directly related to new CCN production. A compound needs to have gas-particle partitioning coefficient in excess of Kom > 102–104 m3 µg−1 in order to be considered non-volatile. Assuming absorptive partitioning, this corresponds to saturation vapor pressures of < 0.01–0.1 ppt. However, organic compounds identified in atmospheric aerosol have Kom < 1 m3 µg−1. These compounds are clearly too volatile to contribute significantly to CCN production. Most likely candidates for non-volatile species are multifunctional compounds, such as oxocarboxylic acids, diols, and hydroxy-carbonyl species, which have just recently been identified in secondary organic aerosol. It is also possible that organic compounds are converted to non-volatile species in multiphase reactions in cloud droplet, though very little is known about the in-cloud production of secondary organic matter.

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New CCN can form either by the condensational growth of very small particles into the so-called “potential activation zone” (where the particles can be activated at high supersaturations associated with strong updrafts), or by further growth of the particles to a size which can activate at more typical supersaturations. Non-volatile organic compounds of a sufficiently large production rate are important in the first step, while the second is thought to be dominated by in-cloud sulfate production [Kerminen, 2001]. a) Modification of the activation of inorganic aerosol by organic aerosol Effect of surface activity Surface tension (σ) is an important factor controlling the equilibrium vapor pressure of small droplets. It arises as a consequence of intermolecular forces that tend to minimize the surface area of the droplets. The surface tension of pure water—which is high due to strong hydrogen bonding—is significantly reduced in the presence of surface-active solutes (surfactants). A possible structure for internally mixed inorganic and organic particles is an aqueous or solid core with organic coating [Andrews and Larson, 1993]. The compounds that make up of this coating tend to be surface active. Surface activity is expected for carboxylic acids with chain lengths of more than 5 carbon atoms [Gill et al., 1983]. Lowering of surface tension by organic compounds dissolved in hydrated aerosol or cloud droplet directly influences the air-water interface of the droplet, which has an effect on several cloud processes such as droplet nucleation, growth, and coalescence. Furthermore, it can affect air-water exchange processes and multiphase chemical reactions within the droplets [Gill et al., 1983; Shulman et al., 1996]. However, earlier measurements on the surface tension of atmospheric waters [Seidl and Hänel, 1983] led to the conclusion that the surface tension depression caused by surface active materials was too small to have a significant effect on cloud microphysics. This might be true for developed cloud droplets, though incipient cloud droplets have typically larger concentrations of organic species. The conditions in incipient cloud droplets were simulated by vacuumevaporating collected fog water samples while measuring σ [Facchini et al., 1999]. At the highest concentrations the relative decrease of surface tension was found to be ∼30 %. According to chemical analyses the observed reduction was mainly due to high molecular weight polyacids, which are known surfactants. This decrease in σ would lead to a 20 % increase in the number of cloud droplets, N. Given constant liquid water content in clouds

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this decrease in N corresponds to a 6 % reduction in the average size of a droplet. The surface tension of the cloud/fog samples and aerosol water extracts markedly decreases with respect to pure water with the increasing concentration of water-soluble organic carbon, as shown in Figure 5.14 [Facchini et al., 2000]. The surface tension depression was highest (−10–20 %) for Po Valley fog water, which exhibited the highest WSOC values. This decrease was attributed largely to polycarboxylic and mono/dicarboxylic acids. The specific surface-activity of the polycarboxylic acids was four times higher than that of mono/dicarboxylic acids, and an order of magnitude higher than that of the neutral compounds. The surface tension lowering effect of polyacidic compounds was estimated to be 0.32 mN m−1 per mgC l−1 [Facchini et al., 2000]. Unfortunately, this value, which merely served for comparison of the surfactant properties of various subsets of WSOC on a basis, was wrongly adapted for subsequent atmospheric studies, as the relationship between surface tension and concentration is logarhytmic (see Eq. 5.13). The high value obtained for the group of polycarboxylic acids was not surprising since humic substances, to which polycarboxylic acids closely resemble, are well known for their surfactant properties.

Figure 5.14. Surface tension measurements of cloud and fog samples from different locations and of aerosol water extracts. Data on surface tension of fog samples collected in Switzerland by Capel et al. [1990] are also reported (after Facchini et al. [2000]).

The dependence of surface tension of cloud water on DOC was also studied in cloud water collected at a mid-level mountain in Central Europe

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[Hitzenberger et al., 2002]. The surface tension is shown in Figure 5.15 as a function of DOC concentration.

Figure 5.15. Relative surface tension vs. OC content of cloud water samples (after Hitzenberger et al. [2002]).

The average surface tension relative to that of pure water was found to be 95.2 ± 3.7 %, with a minimum of 83.8 %. The reduction in surface tension was generally observed at a lower DOC concentration than in the case of other studies. Since surface active compounds are enriched on the surface of the cloud water samples, care has to be taken when considering bulk DOC concentrations. Even if all DOC is considered to be surface active and distributed over the surface of suspended cloud droplets, its amount is far to low to furnish even a monomolecular layer on the surface of all droplets. On the other hand, even a fraction of the DOC can effectively cover haze droplets of 0.1 µm in diameter. The variations therefore most likely reflected compositional differences among the different cloud water samples rather than the effect of different DOC concentrations. When considering the reduced σ in the Köhler equation, the critical supersaturation needed for particles of 0.05 µm diameter was lowered by 7 % on average (by 25 % in the case of the maximum reduction). Such a reduction in Sc may lead to higher cloud droplet concentration and smaller effective droplet radius. Alternatively, it may enable the largest CCN to get activated at lower supersaturations, thereby preventing the more numerous smaller nuclei from being activated by depleting the limited water vapor supply. Li et al. [1998] selected sodium dodecyl sulfate as a representative of surfactants, and studied the change in activation parameters at variable mass ratio of this surfactant and NaCl according to the Köhler theory. At

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sufficiently large concentrations, above the so-called critical micelle concentration (CMC), the molecules of the surfactant form particle-like structures called micelles. Micelles are clusters of surfactant molecules in aqueous solutions which are organized to be isolated from water molecules by grouping their hydrophobic ends together, with their hydrophilic ends facing outwards. A consequence of this behavior for the Köhler theory is that σ becomes practically constant if the concentration of the surfactant exceeds the CMC. Another factor that needs to be considered is the tendency of surfactant molecules to accumulate on the droplet surface, as expressed by the excess surface concentration (γm). These two phenomena make surface tension (σ) dependent on the droplet radius, a fact that is unaccounted for in the traditional formulation of Köhler theory. As a particle, consisting of a mixture of dodecyl sulfate and NaCl, grows by water vapor condensation, σ increased due to changes in concentrations of both the organic and inorganic components. However, in small droplets with large surface to volume ratio, there may not always be enough surfactant to cover the entire surface, even if all surfactant molecules reside on the surface. For example, in a dry particle of 0.05 µm in diameter, initially containing only 20 % sodium dodecyl sulfate by mass, the critical radius (rc) exceeded the maximum radius corresponding to full surface coverage by a factor of 4. Consequently at rc sigma was close to that of pure water (73 kN m−1). Under the same conditions for pure sodium dodecyl sulfate particles, there was nearly 14 times more surfactant at the interface than in the bulk volume of the droplet. It was shown that with increasing mass ratio of NaCl Sc decreased monotonically, consistent with the Raoult term which is larger for NaCl than for sodium dodecyl sulfate. This is principally due to the factor of 5 difference between the molecular weights of the solutes, resulting in almost fivefold decrease in solute molar concentration when NaCl is replaced with the organic surfactant. This translates into a 270 % increase in Sc, which suggests that surface tension depression caused by the surfactant is unable to compensate for the significant change of the Raoult effect. Such response implies that under atmospheric conditions soluble surfactants inhibit the activation of inorganic salts. These conclusions seemingly contradict the findings of Mircea et al. [2002], who calculated a significant decrease in critical supersaturation (Sc) when taking into account water soluble and surface active organic compounds. Upon closer look at the two approaches, however, it becomes evident that they considered different baseline cases. Li et al. [1998] replaced soluble inorganic salt (NaCl) with the surfactant, whereas the other authors assumed an insoluble core in the baseline case, and partly replaced it with

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water-soluble surface active organic compounds. In the latter case, the effects of the organic compounds on both the Kelvin term (reduction in σ, no matter how imperfect it may be) and on the Raoult term (addition of soluble mass, even if it is not very efficient on a molar scale) leads to a reduction of Sc. It is not our competence to decide which approaches can be justified, but at least both can be correct and they are obviously not in contradiction. Effect of an organic surface film The change in CCN activity of pure (NH4)2SO4 particles at 0.3 % supersaturation upon coating with glutaric acid was studied in a laboratory study [Cruz and Pandis, 1998]. A particle was considered activated if—being exposed to a given supersaturation—it grew into cloud droplet within less than 30 seconds. The hygroscopic organic compound was found to actually increase the number of particles activating at the given supersaturation. Since the experimental activation diameters of binary particles were found to fall within the values for pure substances, the effect was thought to be additive. This additive behavior can be adequately described by the formula for two nonvolatile compounds of Shulman et al. [1996]. Theoretically, a difunctional organic acid can either decrease surface tension of the solution (and thus increasing activation), or change the shape of the Köhler curve due to its gradual dissolution, thus delaying activation. Nevertheless, no net effect was observable within the uncertainty of the measurements. The same experiments were repeated with an insoluble organic compound, dioctyl-phtalate, which was either expected to act as an inert mass or hinder activation by forming a hydrophobic coating on soluble particles [Cruz and Pandis, 1998]. Even when dioctyl-phtalate formed a thick coating on the particles, the measurements did not indicate any negative effect on activation. The characteristic time-scale for the diffusion of water through a 0.1 µm dioctyl-phtalate coating was calculated to be in the order of 10−6 s, far too low to cause measurable changes in activation. It is also possible that dioctyl-phtalate decreased the surface tension of the solution, but overall, it behaved as an inert mass during activation. Effect of limited solubility In terms of solubility, four generic classes of potential CCN active material were identified that may cause deviations from the traditional Köhler formulation [Laaksonen et al., 1998]. It should be noted that in the original publication the examples were all inorganic species, but organic compounds can also contribute to each class:

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1) Slightly soluble substances that partly or fully dissolve in a droplet in the intermediate stage of its growth. Such substances, like an insoluble core, effectively increase the radius of the droplet in the initial stage (thereby decreasing Sc), then gradually add solute to the growing droplet (thereby increasing the Raoult effect); 2) Highly soluble gases (vapors) that can add solute to the droplet and effectively depress vapor pressure at a very early stage of the droplet growth, then continue to partition into the droplet and quickly deplete their gas phase concentration within the clouds; 3) Slightly soluble vapors that are mostly in the gas-phase but partition into the droplet providing solute throughout the period of droplet growth, their gas-phase concentrations are not likely to be depleted by the uptake in droplets, 4) Surface active soluble or slightly soluble substances, such as polar organic molecules which reduce the Kelvin term (see above). The inclusion of these classes into the Köhler theory (or rather an extension of it) questions the simple dichotomy of cloud and haze droplets, inasmuch as large but unactivated droplets can exist in clouds. Particles are not necessarily activated to cloud droplets when they have passed the first maximum. In the case of highly soluble gases a second maximum may exist that is even higher than the first local maximum. Thus the condensable gases and/or the slightly soluble substances may cause a minimum or several minima in the equilibrium saturation curve. The growing droplet would have to pass over those maxima before getting activated. This sort of behavior would make the growth/evaporation processes discontinuous with associated sudden increase/decrease in particle radius. Stable cloud droplets in which the soluble mass is either derived from slightly soluble aerosol particles or soluble gases could exist in relative humidity below 100 % up to 10 µm in size [Kulmala et al., 1997]. The incorporation of these two terms in a realistic scenario, along with the surface tension depression caused by the slightly soluble substance, may be sufficient to keep the equilibrium saturation below unity for droplet sizes of up to 10 µm. In fogs formed by such processes, cloud drop-sized particles behave thermodynamically similarly to submicron haze particles, so they are not activated. It should be noted that the original calculations used CaSO4 as a surrogate for slightly soluble substances, but organic compounds may behave similarly. The physical state of slightly soluble organic compounds depends strongly on their solubilities in the aqueous salt solution and the droplet volume. These effects can only be considered in a modified form of Köhler

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equation, since the original theory assumes complete solubility [Shulman et al., 1996].

Sw ≅

2σM w 3M w Φ  v dssc X dssc m ssc v sulf m sulf  − + 3  kTρr M ssc M sulf  4 πρ r     

Kelvin Effect

(5.12)

Raoult Effect

where σ is the surface tension of the solution, Mw is the molecular weight of water, k is the Boltzmann constant, T is the temperature, ρ is the density of the solution, Φ is the osmotic coefficient of the solution, νssc is the number of ions to which a slightly soluble compound dissociates when dissolved, νsulf is the number of ions to which sulfate dissociates when dissolved, mssc is the total mass of the slightly soluble compound, Xdssc is the fraction which dissolves into the droplet during droplet growth, msulf is the total mass of sulfate salt, Mssc is the molecular weight of the slightly soluble compound, Msulf is the molecular weight of sulfate salt. When using this equation, Shulman et al. [1996] made the assumption that only the dissolved material can depress vapor pressure, to the extent proportional to its mole fraction, so that the effects of sulfate and slightly soluble compounds were simply additive. In their calculations, they considered five dicarboxylic acids (C2–C6), phtalic acid and cis-pinonic acid. The solubilities of these compounds in pure water and in a range of salt concentration were measured experimentally. All organic species were present in two phases (in saturated solutions and in solid form) until the droplet acquired sufficient volume of water to fully dissolve them. This point was characterized by the ‘transition radius’ (rT). This radius increased with decreasing organic solubility, from oxalic to phtalic acid. As long as this point is not reached, two competing effects govern the concentration of slightly soluble organic species in the droplet. The growth of the droplet decreases the concentration by simple dilution. On the other hand, with decreasing salt concentration organic species become more soluble. Even if a constant surface tension was assumed, the presence of slightly soluble organic species can significantly alter the shape of the Köhler curve, actually there appeared two maxima separated by a cusp, as shown in Figure 5.16. This cusp was caused by the abrupt termination of the dissolution process at rT, which affected the Raoult term in Eq. 5.12 discontinuously. This implies that the droplet only partially activates at the first maxima if the second maximum has a higher supersaturation than the first.

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Figure 5.16. Breakdown of the modified Köhler curve. The conditions for this calculation are: dry radius of 0.05 µm, 80 % succinic acid by mass, 20 % (NH4)2SO4. (SSC: Slightly soluble compounds) (after Shulman et al. [1996]) .

In the calculations experimentally measured surface tension depression effects were considered which were significant for all organic species but phtalic and oxalic acids. Their magnitude depended on the carbon number of the compounds, and salt concentration had little effect on it except for cispinonic acid. By applying surface tension correction, rT was not shifted but the critical supersaturation decreased with a simultaneous increase in the critical radius. It was concluded that surface active and slightly soluble organic compounds affect cloud droplet growth by both gradual dissolution into the growing droplet which alters the shape of the Köhler curve, and lowering its surface tension, which decreases critical supersaturation. Reduction in the critical supersaturation promotes droplet activation and increases the number of CCN, whereas the presence of a cusp in the equilibrium curve would delay activation. Since many such compounds exist together in an internally mixed CCN, a series of transition radii and cusps would be present on the curve. As a result, there would be no steep negative slope allowing rapid growth after activation. Instead, the curve would show an overall flattening, with a metastable size range around the cusps. The consequence of this is a delayed droplet growth, consistent with the observations in thermal gradient CCN counters for ambient aerosol. This effect was originally attributed to the formation of insoluble organic coating on the soluble core, impeding the uptake of water [Bigg, 1986]. A clear difference was found between the activation rates of equally sized monodisperse organic (pinonic acid) and inorganic (ammonium sulfate) aerosol [Hegg et al., 2001]. The measurements also suggested that the

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presence of succinic acid internally mixed with ammonium-sulfate did indeed retard the activation process, though the uncertainties were quite large. Since the observed delay was at most about 4 min, it is not clear whether the results have atmospheric significance. Eventually, observed values of cloud droplet concentrations caught up with the model predictions based on virtually complete solubilities. A modified form of the Köhler equation, which contained a Raoult term describing the additive effect of the inorganic and organic solutes was used to simulate the growth of a droplet [Mircea et al., 2002]. In addition, the additive effect was extended to a droplet containing an insoluble core. Water-soluble organic compounds can add to the soluble mass (Raoult term) and can decrease the surface tension of the droplet (Kelvin term). As a baseline case, 50 % inorganic solutes (a mixture of 26 % (NH4)2SO4 and 24 % NH4NO3) was considered on an insoluble core (50 % by mass). To study the effect of water-soluble organic species, this inorganic core was reduced to 30 % and the remaining fraction was replaced by a hypothetical dicarboxylic acid having Mw = 100 and ν = 3, and a fulvic acid with Mw = 732 and ν = 5 in proportions of 14 and 6 %, respectively. The use of the limiting values of the van’t Hoff factors, however, is questionable in the light of the results of Young and Warren [1992] and Konopka [1996]. In the first case, only the effect of soluble mass was considered, and the surface tension was regarded equal to that of pure water. The calculations showed that a decrease in critical supersaturation followed the addition of the soluble organic mass. In the second case, the decrease of the surface tension of the droplet was also taken into account according to an empirical relationship as a function of dissolved organic carbon concentration [Facchini et al., 1999]:

σ s = 72.8 − 0.0187T ln(1 + 628.14c )

(5.13)

where T is the temperature (K) and c is the concentration of soluble carbon in mol l−1. In this case a further decrease in the critical supersaturation was calculated. In spite of the fact that Equation 5.13 was obtained for Po Valley fog water measurements, the model was applied to aerosol chemistry and number size distributions for maritime, rural and urban scenarios. The differences against the baseline case in the total number of CCN calculated as a cumulative number concentration (Eq. 5.11), were +13 % for the maritime, +97 % for the typical rural and +110 % for the urban scenario. It was established that the relationship describing the CCN supersaturation spectrum differs considerably from the power law, which has been widely used in

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current cloud models. It was argued that for rural and urban aerosol, an exponential function should be used instead. Nenes et al. [2002] made model simulations with the Köhler theory to predict the variations in total number of cloud droplets, Nd, when taking into account the presence of partially soluble solutes in the growing droplet [Shulman et al., 1996], the surface tension (σ) depression by dissolved organic substances [Shulman et al. 1996; Facchini et al., 1999], and the changes in the accommodation coefficient caused by the formation of organic films on the droplet surface [Feingold and Chuang, 2002]. The baseline aerosol was assumed to be pure (NH4)2SO4. The aerosol organic component was a simplified representation of the measured composition of fog droplets in the Po Valley, consisting of 18 % (m/m) levoglucosan, 41 % (m/m) succinic acid, and 41 % (m/m) fulvic acid. The surface tension depression was calculated by Eq. 5.13. The solubility of the organic mixture was varied between 10−4 and 10−1 M. The accommodation coefficient (α) was assumed to have a low value (10−5) when a continuous film existed, and attained a high value (0.043) when it broke (i.e. its thickness fell below 0.5 nm). The simulations showed that the presence of water-soluble organic species, even without affecting the surface tension of the droplets, can alter the number of droplets by changing the moles of dissolved solutes. Since the activation of the organic mixture did not differ significantly from that of pure (NH4)2SO4, this change was not spectacular. Once effects on σ were taken into account, Nd increased relative to the baseline case, being more pronounced at high updraft velocities for marine aerosol, and at low updraft velocities for urban aerosol. For the marine scenario, chemical effects altogether can approach 50 % of the Twomey effect (see sub-chapter 5.2.4), whereas for the urban scenario, they can even exceed it. It was observed that when solubility was lower than 10−4 M, the organic matter behaved as if it were insoluble. The largest surface tension depression was found in the saturated solution. The surface tension depression of dissolved organic compounds was believed to have a stronger effect on activation than its sheer contribution to the soluble mass. The effect of semi-volatile and water-soluble organic compounds on cloud droplet number concentration was modeled in a recent study [Anttila and Kerminen, 2002]. When the value of the Henry-constant was < 108 atm−1 and the gas-phase concentration of the organic species was < 1 ppb, the change in number concentration relative to the base case remained well below 10 %. Significant increase in concentration was observed when c > 100 pptv and H > 109 atm−1. The effect was most prominent at high updraft velocities and in the marine boundary layer where CCN concentrations are typically much lower.

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For slightly soluble compounds, solubility had an effect only in the range of 0.01–1 M. Below 0.01 M, the amount of dissolved organic compounds was too small to influence cloud droplet activation, and above 1 M all of their mass dissolved in the aqueous phase. To have a notable influence on cloud droplet activation, the value of Kom must be > 0.1 m3 µg−1, in the upper range of reported values. Overall, it seems that slightly soluble, semivolatile organic compounds have only a minor influence in the continental boundary layer, but can potentially exceed 10 % in the marine boundary layer, subject to the conditions that 1 µg m−3 of sufficiently soluble (> 0.1 M) and relatively nonvolatile (Kom > 0.1 m3 µg−1) organic compound is present. It should be noted, however, that according to the atmospheric observations this is not a realistic occurrence. Therefore most semivolatile and slightly soluble compounds are not able to affect cloud droplet number concentrations via mechanisms described by the Köhler theory. The total concentration of dicarboxylic acids are well below 100 ppt in remote areas, and their Henry’s law constants are in the range of 108–109 M atm−1, consequently their contribution to cloud droplet activation is likely to be small. b) Modification of the activation of inorganic aerosol by soot particles It has been shown recently that soot particles embedded in cloud droplets can absorb solar radiation and elevate the temperature by a fraction of a degree, which is sufficient to increase the maximum equilibrium supersaturation of the droplet [Conant et al., 2002]. This effect can be quite substantial at low supersaturations (0.01 %) and for larger particles (above 2– 3 µm in diameter), in which it can sometimes override the Raoult effect leading to an increase in Sc for the largest sizes. The effect of particles containing 10 % BC by mass led to a 1 % reduction in CCN at a supersaturation of Sc=0.03 %. At supersaturations exceeding 0.08 % the reduction in CCN was negligible. It can be surmised that in clouds with slow cooling rates this heating effect suppresses the activation of the largest CCN which might have some impact on the water balance during cloud formation. It can also be speculated that this effect can be even more significant in contrail and high-altitude stratiform cloud formation, where aerosol particles have a large fraction of soot.

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c) CCN-activity of pure organic particles Pure cellulose was burned under controlled conditions to study the CCN activity of pure organic smoke particles [Novakov and Corrigan, 1996]. Thermograms of aerosol from the smoldering combustion of pure cellulose showed overwhelmingly organic species with a wide range of volatility. A large fraction of these species were found to be soluble in water, volatile species the most and refractory species the least. The simultaneously recorded CCN to CN ratios (at supersaturations of 0.3 and 0.5 %) were found to be close to unity, indicating that the smoldering combustion of cellulosic biomass is a source intrinsically CCN active organic particles, attributable to their water soluble fraction and not to their size. Some studies suggested that organic CCN could act quite differently in conventional water-based CCN counters [Bigg, 1986; Chuang et al., 1997]. The reasons for their anomalous behavior might be either changing the growth rate of activated droplets or altering the activation radius of the aerosol particles [Hegg et al., 2001]. The CCN activity of two hydrophilic dicarboxylic acids, adipic acid (C6H10O4) and glutaric acid (C5H8O4), thought to be representative of a much larger groups of secondary compounds, was studied in laboratory experiments [Cruz and Pandis, 1997]. The theoretical activation diameters (dp*) were calculated from available physical and chemical data to be 98 and 103 nm for glutaric and adipic acids, respectively, at a supersaturation of 0.3 % (the corresponding values for supersaturation of 1.0 % were 44 and 46 nm, respectively). The experiments indicated that pure adipic and glutaric acids did act as CCN, and their behavior was analogous to that of the inorganic salts, as observed in the shape of the experimental activation curves. The measurements were in good agreement with the Köhler theory at both supersaturations, except for glutaric acid at a supersaturation of 1.0 %, which showed a larger deviation (+30 %) from the theory. It was assumed that dicarboxylic acids at higher concentrations may either show some nonideality (i.e. Φ ≠ 1 in Eq. 5.12), or may reduce surface tension of the solution by up to 10 %, the incorporation of which into the calculations could reduce the deviations considerably. In contrast, dioctyl-phtalate, a good representative of water-insoluble organic species, was found not to show CCN activity up to a supersaturation of 1.2 %, and up to particle diameters of 150 nm. In general, it was concluded that the dicarboxylic acids studied are less CCN active per unit mass or volume than the inorganic salts, due to their higher molecular weight, lower density and lower equivalence, which are all important parameters in the Köhler theory.

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The critical activation diameters of succinic acid, adipic acid and glucose aerosol particles were determined experimentally at 0.4, 0.5 and 0.8 % supersaturation [Corrigan and Novakov, 1999]. The critical diameter of succinic acid and glucose particles were found to be comparable to that of inorganic salts, being on average only 1.3 and 1.8 times higher, respectively, than that of NaCl. On the other hand, adipic acid had a critical diameter a factor of 3.7 higher than that of NaCl. The measured values were also compared with theoretical values calculated from the Köhler theory, and were in reasonable agreement for succinic acid and glucose, but a factor of two higher for adipic acid. In the case of the latter, this deviation might be caused by the delayed droplet growth due to the slight solubility of adipic acid. The critical diameters as well as the agreement with the theory followed the general trend of solubility. This may not be surprising in the light of the fact that traditional Köhler theory assumes that the particle is infinitely soluble in water. In addition, the van’t Hoff factor and surface tension depression effect of pure organic compounds may deserve attention. The van’t Hoff factor represents the number of dissociated species per solid salt molecule that would form if the salt is dissolved in an infinitely dilute solution. Since the solution droplet is far from being infinitely dilute at activation, this approximation is not valid even for inorganic solutes. Therefore it may not be applicable to organic species, such as succinic acid, for which ν = 3 at infinite dilution, leading to the faulty conclusion that succinic acid is more CCN active than ammonium sulfate. The surface tension (σ), also varies with concentration and can have a large effect on the hygroscopic growth of a particle. However, since at activation the solution is rather dilute (0.02 M for a 100 nm dry particle of adipic acid), the resulting reduction in surface tension is below 5 %. Pinonic acid, pinic acid, and norpinic acid, known photooxidation products of α-pinene (see sub-chapter 3.3) were found to activate to cloud droplets at supersaturation of 0.3 % below dry diameter of 120 nm. Similar results were observed for glutamic and glutaric acids [Raymond and Pandis, 2002]. The addition of a functional group tended to reduce the activation diameter. There were, however, exceptions to this rule. Leucine, a six-carbon molecule with two functional groups, having solubility slightly above those of pinonic acid and glutamic acid, was expected to activate similarly to adipic acid. However, pure leucine particles of 200 nm dry diameter were found to barely activate even at 1 % supersaturation. On the other hand, cholesterol, with low solubility and only one functional group, which was expected not to be activated similarly to myristic acid, did get activated at a diameter of only 48 nm at 1 % supersaturation. It appears that that ability of a pure chemical species to activate is related not only to its solubility in water, but also its ability to spread or form

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a finite contact angle with water. Every species that either readily absorbed water or was barely soluble (i.e. with solubility less than 0.01 g cm−3), but allowed water to spread on its surface (zero contact angle), was found to behave identically in terms of activation. Contrary, species with low solubility that formed a finite contact angle with water (such as those having more than 14 carbon atoms) proved to be inactive under the same conditions. In the Köhler theory the species which is barely soluble but wettable, should be treated as if it were completely soluble to improve agreement between theory and measurements [Raymond and Pandis, 2002]. 5.2.2.3 CCN-activity of BC The water nucleation properties of diesel soot particles were determined experimentally in the laboratory [Lammel and Novakov, 1995]. With a compilation of available literature data the activated fraction of the total particles (i.e. N(CCN)/N(CN) ratios) were compared at a fixed supersaturation of 1 % for various fuel and combustion types. This comparison, which is given in Table 5.2, implies that diesel engine exhaust may have a nucleation activity similar to that of wood smoke from flaming forest fires. Table 5.2. Activated fraction of aerosol particles, N(CCN)/N(CN), immediately after generation at 1% of supersaturation and content of soluble ions in the particulate matter (analyzed for Na+, NH4+, K+ Cl−, NO3−, SO42−, HCOO−, CH3COO− and—in most cases—NO2−) for various fuel and combustion types (after Lammel and Novakov [1995]). Fuel

Type of combustion

Aviation fuel JP-4

Flaming Sweet crude oil, low Ignition, flaming sulfur Wood smoke Smoldering Diesel Diffusion flame Acetylene gas

Welding torch Light crude oil, high Ignition, flaming sulfur Diesel Engine motor, idling Forest and brush fire smoke Flaming *S = 0.8%

N(CCN)/N(CN) Soluble ions References (%) at S = 1% (meq g−1) 1 Not known Hallett et al. [1989] 20

1.25

24

1.3–3.8

42

0.094

49–53

3.13

58

1.22

80*

0.99

80–100

1.3–3.8

Rogers et al. [1991a] Hallett et al. [1989] [Lammel and Novakov, 1995] Hallett et al. [1989] Rogers et al. [1991a] [Lammel and Novakov, 1995] Hallett et al. [1989] and Rogers et al. [1991b]

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It should be noted, however, that the various sources emit variable amounts of soluble inorganic ions, so the observed water nucleation properties refer to a mixture of inorganic and organic species. It was shown, however, that ultrafine pure carbon particles in themselves could act as CCN, in particular after oxidation by ozone, a process understood to produce surface oxygen-containing functional groups [Kotzick et al., 1997]. 5.2.3 Ice-nucleating (IN) Activity of Carbonaceous Aerosol It is well-known that homogeneous freezing-nucleation of water requires a supersaturation of around 450 %, which translates into a temperature of about −40 °C. However, ice formation is frequently observed at much higher temperatures in the troposphere. This process, called heterogeneous nucleation, requires the presence of a minor subset of aerosol particles4 called ice nuclei (IN), which provide specific sites on which ice embryos can grow more readily than purely by the random internal aggregation of water molecules. In this respect, the interaction between the surface and water molecules is of utmost importance. In general, very little is known about the ice nucleating ability of carbonaceous aerosol particles, in spite of the fact that organic ice nucleators were intensively studied soon after the discovery of the ice-nucleating ability of silver iodide. However, most of the compounds tested are of little relevance in the natural atmosphere, with very few exceptions. Most information on atmospheric ice nuclei have come quite recently in studies focusing on contrail formation in the upper troposphere, with particular emphasis on the possible role of soot particles. 5.2.3.1 Atmospheric Observations Ice nucleating particles larger than ∼0.1 µm were collected in the upper troposphere and lower stratosphere. It was found that a major component of these nuclei in aircraft exhaust and contrail influenced air was carbonaceous particles, together with metallic particles [Chen et al., 1998]. The contribution of soot particles to the number of non-volatile residue of ice crystals in aircrafts contrails varied between 5 and 25 % [Twohy and Gandrud, 1998].

4 Typical number concentration of IN in the atmosphere is in the order of 102–104 m−3, as against aerosol concentrations of 103–105 cm−3 (!).

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5.2.3.2 Laboratory and Modeling Studies It was found that aliphatic alcohol monolayer present on water droplet surfaces can induce ice nucleation at temperatures increasing monotonically with chain length, approaching 0 °C for the longest chain homologue (C31) [Gavish et al., 1990]. A systematic difference was revealed between the efficiencies of the even and odd carbon number homologues, the former being more active. The possible reason for the ice-nucleating activity is the hexagonal pattern of the OH-groups in the monolayer5 which closely resembles the structure of water ice. This finding may have atmospheric implications, since hydrated aerosol particles are frequently envisaged to be covered by organic surfactants (e.g. marine aerosol, see Chapter 2.2). It is also well-established that certain strains of bacteria, lichens and fungi can be quite effective atmospheric ice nuclei [Szyrmer and Zawadzki, 1997]. The ice-forming activity of soot particles at temperatures between −5 °C and −20 °C was studied experimentally [Gorbunov et al., 2001]. The fraction of soot particles that act as contact ice nuclei was found to increase with decreasing temperature and increased surface oxidation. The latter can be explained by the hydrogen-bonding ability of the surface functional groups, which reduces the free energy required for embryo formation. At −20 °C, the activated soot fraction may account for 10 % of observed ice nuclei concentrations. 5.2.4 Contribution of Carbonaceous Aerosol to Aerosol Indirect Forcing Aerosol indirect forcing is defined as a suite of complex processes by which aerosol perturb the Earth-atmosphere radiation balance by modulation of cloud albedo and cloud amount [Penner et al., 2001]. The schematic of these processes linking aerosol mass, cloud condensation nuclei (CCN) and ice nuclei (IN) concentrations to cloud physical properties and effects is shown in Figure 5.17. The aerosol indirect effect is usually divided into two basic categories. The first indirect or Twomey effect relates the increased aerosol (and CCN) number concentration to the decrease of effective cloud droplet radius and consequently to the increase mean cloud droplet number concentration and the resulting increase in cloud reflectivity for marine stratus clouds [Twomey, 1977]. The second indirect effect, which is also linked to the decrease of the effective cloud droplet radius, is associated with the suppression of precipitation and the resulting increase in fractional cloudiness. 5

As revealed by X-ray diffraction.

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Carbonaceous Aerosol

Although these terms refer primarily to warm clouds, there have been observational as well as theoretical indications that cirrus clouds may also be affected by aviation-related aerosol [Fahey et al., 1999]. The upward trend in cirrus fractional cloudiness in areas of high air traffic over the last two decades beyond the extent attributable to linear contrails alone was considered as evidence for the perturbation caused by anthropogenic aerosol. Satellite and in-situ aircraft measurements of ship tracks in marine stratocumuli provided the first direct evidences that effluents from ship stacks changed cloud microstructure by redistributing their water content into a larger number of smaller droplets and caused measurable increase in cloud reflectivity [Coakley et al., 1987; Radke et al., 1989]. More recently, such enhancement of cloud reflectivity in polluted warm clouds was firmly established by direct measurements leaving little doubt that aerosol do have an impact on cloud-related radiative forcing [Brenguier et al., 2000]. The large CCN concentrations in the plumes of burning vegetation nucleate many small droplets that coalesce inefficiently into raindrops. Recent satellite observations revealed that smoke from biomass burning did indeed suppress precipitation [Rosenfeld, 1999]. Although similar effects can be expected for urban and industrial air pollution, there are sporadic contrasting views that air pollution might enhance precipitation on a large scale in northeastern America [Cerveny and Balling, 1998]. But the general view is that the drizzle, which normally occurs in clean marine stratocumulus clouds, is strongly suppressed in clouds of reduced droplet size, thereby increasing their water content and longevity [Albrecht, 1989]. Orbiting weather satellites can reveal numerous ship track-like features in clouds over land, created by major urban and industrial pollution sources [Rosenfeld, 2000]. In these pollution tracks, the median effective radius (rc) of the cloud tops was well below the precipitation threshold of 14 µm, showing little growth with decreasing temperature. Outside the plumes, however, rc increased steeply with decreasing T to over 25 µm, indicating coalescence to precipitation. Observations on marked pollution tracks over southeastern Australia revealed that pollution suppressed precipitation also by preventing the formation of ice particles and cold precipitation processes within the clouds. The effect of weather modification by pollution is not restricted to shallow warm clouds, since the precipitation formation in deep tropical clouds is also likely to be affected by the less efficient accretion of smaller supercooled cloud drops by the growing ice particles [Rosenfeld, 1999].

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Figure 5.17. Flow chart showing the processes linking aerosol emissions or production with changes in cloud optical depth and radiative forcing. Bars indicate functional dependence of the quantity on top of the bar to that under the bar. Symbols: CCN (Cloud condensation nuclei); CDNC (Cloud droplet number concentration); IN (Ice nuclei); IP (Ice particles); OD (Optical depth); HC (Hydrometeor concentration); A (Albedo); fc (Cloud fraction); τc (Cloud optical depth); ∆F (Radiative forcing) (after Penner et al. [2001]).

5.2.4.1 Contribution of Organic Aerosol to Aerosol Indirect Forcing A rough estimate of the indirect forcing attributed to anthropogenic carbonaceous aerosol was first given in a global model by Penner et al. [1996]. For the model calculations, a previous BC emission inventory was used, and for fossil fuel burning organic aerosol emissions were scaled to this inventory by measured urban OC/BC concentration ratios. Biomass burning emissions were estimated on the basis of reported emission factors. Secondary organic aerosol from natural sources was also taken into account, assuming a constant 5 % aerosol yield for terpene emission. This yielded an estimated global production of 7.8 Tg yr−1, which was considered as if it were primary emission in the model.

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Although the model apparently underpredicted OC and BC concentrations for polar regions, it was used to give a rough estimate of magnitude of the indirect forcing. A calculated maximum forcing of −4.4 W m−2 was attributed to anthropogenic carbonaceous aerosol alone. The presence of a prescribed marine background aerosol reduced this value to −2.43 W m−2, which might be reduced further if the effects of BC on cloud albedo had also been considered. The temporal distribution of this indirect forcing showed maxima in April and September, the biomass burning seasons in the tropics. The indirect effect resulting from the changes in carbonaceous and sulfate aerosol concentrations since the pre-industrial times was determined in another global model [Lohmann et al., 2000]. In the model half of the organic carbon and 20 % of BC were assumed to be initially hydrophilic. The hydrophobic components were converted into hydrophilic with a timeconstant of 40 h. Since modeling of indirect forcing is highly sensitive to the background conditions, the pre-industrial scenario was assumed to include no fossil fuel emission and 10 % of the biomass burning emissions of today. This yielded a pre-industrial carbonaceous aerosol burden of 0.4 TgC. The OC and BC climatology for fossil fuel and biomass burning were taken from Liousse et al. [1996], and biogenic secondary aerosol from Guenther et al. [1995]. The simulations were performed for two distinct, and to some extent extreme approaches. In the first, it was assumed that anthropogenic sulfate only increased the mass of pre-existing aerosol and did not form new particles. Consequently sulfate was internally mixed with other aerosol components. In the second, external mixture was considered, when sulfate was allowed to form new particles, and the contributions by sulfate and carbonaceous particles were calculated separately. Since in the model sulfate aerosol had larger mode radius than carbonaceous aerosols (69.5 nm as compared to 21.2 nm dry radius), the specific increase in number concentration was more significant for carbonaceous aerosol. The change in shortwave cloud forcing at TOA since the pre-industrial era was calculated to be −1.1 W m−2 for internally mixed aerosol, as a combined effect of the increase in cloud albedo (−0.5 W m−2) and the increased cloud lifetime (−0.7 W m−2). As expected, the indirect effect was largest in the Northern Hemisphere (−1.6 W m−2), being somewhat larger over continental areas (−1.3 W m−2) than over oceans (−1 W m−2). As far as the contributions of the major aerosol components are concerned, sulfate and carbonaceous aerosol accounted for −0.4 W m−2 and −0.9 W m−2, respectively. It should be noted that the effects of the two species are not quite additive in this approach. If external mixture of the species was considered, the indirect effect amounted to −1.5 W m−2, and was almost entirely attributed to carbonaceous aerosol (−1.3 W m−2), as the change caused by anthropogenic sulfate was

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nearly zero. The very small sulfate effect was due to its much larger mode radius and higher density, which resulted in an increase in its number concentration a factor of 20 lower than that of carbonaceous particles when normalized to source strength. In addition, the sulfate burden increased by only one third of that of carbonaceous aerosol. 5.2.4.2 Contribution of BC to Aerosol Indirect Forcing A new mechanism by which aerosol can impact clouds emerged from observations that the dark haze observed over the northern Indian Ocean reduced significantly the areal coverage of trade cumulus [Ackerman et al., 2000].6 The aerosol-induced solar absorption of the typical magnitude observed was shown to drastically alter the properties of trade cumuli. With the simplifying assumption that a soot core of 0.06 µm is embedded in each haze droplet, resulting in a single-scattering albedo of 0.98 (at 0.5 µm) and optical depth of 0.2 for an idealized INDOEX 1998 scenario, the haze layer was calculated to absorb +7.4 W m−2 of solar radiation, causing a diurnally averaged solar heating of +0.5 K d−1 for the cloudless boundary layer. This temperature increase lowered relative humidity and shortened anvil lifetime, amplifying the daytime reductions in fractional cloud coverage (by 25 %) and liquid water path. This effect of soot is termed “cloud burning”, that is the response of clouds to increased atmospheric heating. On a regional scale, reduction in cloud coverage due to solar absorption by absorbing aerosol strongly affected the radiative heat budget at the surface and top-of-atmosphere (TOA). Soot in itself exerted only a small radiative forcing of +0.4 W m−2 under clear-sky conditions, but in the presence of clouds its effect was amplified by nearly a factor of 10. If scattering was also taken in account, the haze exerted a net TOA cooling under clear skies, which was completely offset by cloud burning under cloudy skies. The magnitude and even the sign of the net forcing, however, remained uncertain, depending on assumptions about unpolluted and polluted conditions. Another indirect effect of soot, resulting from the increased soot concentration from pre-industrial times to the present, can be the “glaciation indirect effect”. In this effect the increase in contact ice nuclei number concentration could lead to a more frequent glaciation in clouds and could increase the amount of precipitation via the ice phase mechanism [Lohmann, 2002]. This effect may at least partially offset the indirect aerosol effect on

6

Trade cumuli are usually found over warm tropical water in boundary layers typically 1.5 to 3 km deep, in the conditionally unstable zone between the mixing layer and the trade inversion.

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warm clouds. The schematic comparison of the two effects is given in Figure 5.18. Model calculations revealed that when soot ice contact nuclei were taken into account, the liquid water content decreased in extratropical regions, due to the more efficient freezing of supercooled droplets (a factor of 5–8 more effective compared to the background case without soot, depending on the assumed hydrophilic fraction of soot particles). This initiated precipitation formation in supercooled water clouds which did not precipitate due to the absence of drizzle-size droplets. In turn, the enhanced precipitation scavenged aerosol particles more efficiently, leading to an overall reduction in the global aerosol burden.

Figure 5.18. Schematic diagram of the warm indirect aerosol effect (solid arrows) and glaciation indirect aerosol effect (dotted arrows) (after Lohmann [2002]).

The resulting reduction in cloud cover increased outgoing longwave radiation, causing a net forcing of −1.6 W m−2 for the case when 10 % of soot particles were assumed to be hydrophilic. Interestingly, nearly the same magnitude of forcing (−1.4 W m−2) was calculated between pre-industrial and present times when glaciation effect was not taken into account, resulting from the decrease in shortwave radiation due to the enhanced cloud coverage and cloud albedo. Note that though the magnitude and sign of forcing were very similar in both cases, the signs in shortwave and longwave radiation were completely reversed.

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5.3 Heterogeneous Reactions on Carbonaceous Aerosol in the Atmosphere 5.3.1 Properties of Carbonaceous Particles Relevant to Heterogeneous Chemistry In the stratosphere, the vast majority of carbonaceous particles are identifiable by their fractal geometry, with only a small percent have more amorphous shape [Strawa et al., 1999]. The morphology of stratospheric and upper tropospheric soot is very similar to those observed for particles originating from combustion processes. For their characterization two characteristic dimensions of the particles can be defined: the average diameter of the monomer d0, and the geometric diameter Dge, defined as the diameter of the smallest sphere that encloses the entire aggregate. These dimensions can be used to estimate the number of monomers in a fractal, N, and its aerodynamic diameter, Dge by the following relations [Mandelbrot, 1982]:

(

N = D ge d 0

(

)

f

Dae = d 0 D ge d 0

(5.14)

)(

f −1) 2

(5.15)

where f is defined as the fractal dimension. The fractal dimension of soot particles was found to be 1.91, falling between the typical mean values for aggregates formed by diffusion limited cluster-cluster mechanism [Nyeki and Colbeck, 1995]. This value indicates loosely packed aggregates. Fractal clusters have much less sedimentation velocities than compact spheres of the same mass [Colbeck and Nyeki, 1992]. Another important consequence of the fractal structure of soot is that the actual surface area of a fractal aggregate is about 30 times higher than if the particle were spherical [Blake and Kato, 1995]. The surface area to mass ratio of soot particles is typically a function of the monomer size, fractal dimensions, surface conditions, and other environmental parameters. A theoretical limit (i.e. the total surface area of all monomers) can be approximated by the simple relationship

FSA =

6 ρ0d0

(5.16)

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FSA values of 50 m2 g−1 were observed experimentally for monomer sizes of 50–70 nm [Strawa et al., 1999], which were slightly larger than those found in aircraft exhaust (30–50 nm) [Pueschel et al., 1997]. For the stratosphere, average soot number, surface area, mass concentrations of 0.06 cm−3, 0.03 µm2 cm−3, and 0.64 ng m−3 were obtained, respectively. The total surface area of soot was only 9.5 % of that of sulfuric acid (0.29 µm2 cm−3). This finding was in contrast with previous studies that found comparable magnitudes during volcanically quiescent periods [Blake and Kato, 1995]. Soot from air traffic was found to represent about 0.3 % by mass of the background stratospheric aerosol [Pueschel et al., 1992]. Most of soot particles in the upper troposphere are not entrained in sulfate aerosol. It was speculated, however, that visible contrails form as soot particle acquire a liquid sulfuric acid coating, and soot triggers heterogeneous ice nucleation [Kärcher et al., 1996]. Upon formation of soot spherules many defects, dislocations and discontinuities are formed in the crystal structure of the particles, with a large number of unpaired electrons which make them very reactive. On the surface of soot, nearly all types of oxygen-bearing functional groups were found [Heintzenberg and Winkler, 1991]. These features make soot particles a complex three-dimensional organic polymer with electron-transfer capability. 5.3.2 Heterogeneous Reactions in the Stratosphere on Carbonaceous Aerosol 5.3.2.1 NOx Chemistry In the stratosphere, soot may release nitrogen out of the reservoir HNO3 species into NOx through the renoxification reaction [Rogaski et al., 1997]. BC HNO3 → NO2

(5.17)

BC HNO3 → NO

(5.18)

The uptake coefficient of HNO3 was found to be 3.8 × 102. The products observed were H2O, NO2 and NO. NO2 was found to be 5 times more abundant than NO [Lary et al., 1997]. Under typical stratospheric conditions, however, this reaction was deemed insignificant [Choi and Leu, 1998]. Another possible reaction can be the reduction of NO2 to NO on soot particles [Lary et al., 1997]:

Heterogeneous Reactions on Carbonaceous Aerosol BC NO2 → NO

297

(5.19)

This reaction is particularly important because it is independent of sunlight. The uptake coefficient for NO2 was in the range of 0.04–0.15 [Strawa et al., 1999; Tabor et al., 1993]. If the soot surface area is increased threefold in model simulations corresponding to a 5 % annual increase in air traffic over the next 20 years, measurable increase is predicted in the NOx/NOy ratio due to the renoxification reaction of HNO3 on soot [Strawa et al., 1999]. However, these changes translate into small effects on the ozone loss rate. This is mainly due to the offsetting influences of the NOx catalytic cycles (more effective at higher NOx/NOy ratios) and other catalytic cycles (HOx, ClOx, BrOx), which become less effective because of conversion of the radicals into reservoir species. Similarly low annual ozone loss rates (at most 0.02 % yr−1) are predicted for the NO2 reduction. The highest ozone loss rate is obtained for direct ozone loss (on average 0.07 % yr−1), extending to altitudes above 20 km. It should be noted, however, that these studies focused on Artic summer conditions, when gas-phase reactions are particularly efficient because of extended solar illumination. In other regions of the stratosphere during winter months, heterogeneous reactions on soot may play a larger role. In another study a tenfold increase in the soot loading in the lower stratosphere was predicted to reduce HNO3/NOx ratio by a factor of 100 [Lary et al., 1997]. This finding seemed to account for the hemispheric asymmetry in the observed vertical profiles of NO2 and HNO3/NOx ratio. With no soot particles, simulations predict that approximately 80 % of NOy is in the form of HNO3. When total soot surface area is increased to 1 µm2 cm−3, a value close to the current hemispheric average, the HNO3/NOy ratio is reduced to ∼0.5. At low total soot surface area concentrations, direct ozone loss would be prevalent, but at high surface area concentrations the catalytic loss due to the production of NO would be predominant. The heterogeneous reduction of HNO3 and NO2 was calculated to reduce nighttime ozone lifetime from around 350 months to 52 months [Lary et al., 1997]. On the other hand, the renoxification that occurs on carbon aerosol in the troposphere would tend to increase the ozone concentration. Much depends on at which altitude soot originating from air traffic is transported to. There is a possibility that soot particles having low sedimentation velocities due to their fractal geometry, can be transported up to an altitude of 25 km where they contribute to the ozone loss.

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5.3.2.2 O3 Chemistry Ozone losses may also occur through the more direct interaction of ozone on soot aerosol through the reaction BC O3 → O2

(5.20)

Although the reaction probability of O3 on soot was found to be only about 10−4, the process was deemed a significant sink for ozone in the stratosphere [Fendel et al., 1995]. One study observed a higher reaction probability of 3 × 10−2 [DeMore et al., 1994], but another found a lower value (∼10−5) depending on the soot surface history [Lary et al., 1997]. The major products identified were CO, CO2 and O2, suggesting that the ozone reaction was at least partly non-catalytic. It is critical to account for the extra oxygen in the reaction, the fate of which will determine whether the reactions can be effective over prolonged periods of time. For the most significant ozone reaction, four possible reaction routes were envisaged [Strawa et al., 1999]: Catalysis

2O3 BCA  → 3O2

(5.21)

Oxidation

O3 + C (BCA) → O2 + CO( g )

(5.22)

Poisoning

O3 BCA  → O2 + O(s )

(5.23)

Null reaction

O3 BCA  → O2 + O(g )

(5.24)

Most likely some combination of these four pathways occurs, and the branching ratios depend on environmental conditions and the characteristics of the soot surface. Only Eq. 5.21 represents purely catalytic reaction, which may lead to sustained ozone depletion. In Eq. 5.22, soot particles are gradually destroyed, whereas in Eq. 5.23, their surface is irreversibly altered by the uptake of oxygen, most likely in the form of carboxylic acids. Eq. 5.24 is considered a null reaction, because it regenerates odd oxygen. Available data indicate that ozone loss is not purely catalytic, the branching ratio for carbon consumption was estimated to be between 0.5 and 40 % [Strawa et al., 1999]. The lifetime of soot in the stratosphere can be calculated as the ratio of soot mass to carbon loss rate, approximately proportional to the soot mass to soot surface area ratio. The calculated soot lifetime was in the order of an

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hour, absolutely inconsistent with observed BC loadings. It was speculated that all the reactions compete for the same active sites on soot surfaces, and rapid poisoning of the surface by ozone would prevent HNO3 and NO2 reactions from ever being effective. In addition, soot particles may rapidly acquire a H2SO4 coating in aircraft plumes, which likely poisons the soot surface [Gao et al., 1998]. Atmospheric observations showed that soot particles became more hydrophilic upon oxidation, resembling stratospheric sulfate aerosol [Chughtai et al., 1996]. These hydrophilic particles may even oxidize NOx to nitric acid. In a modeling study, aimed at assessing the role of aircraft-generated soot particles in ozone destruction, soot was assumed to consists of nonvolatile particles of 0.1 µm radius, without sulfuric acid coating, but the surface area enhancement by a factor of 30 due to its fractal structure was taken into account [Bekki, 1997]. The soot specific surface area was about 450 m2 g−1. When ozone decomposition on carbon was considered to be a catalytic reaction, with a reaction probability of 10−3, the model largely reproduced the observed ozone trend at mid-latitudes in the Northern Hemisphere, with the exception of the lower stratosphere, where it underestimated ozone depletion. 5.3.3 Heterogeneous Reactions in the Troposphere on Carbonaceous Aerosol 5.3.3.1 O3 Chemistry

The reaction between ozone and soot in the troposphere can be described with the following reaction sequence [Disselkamp et al., 2000]:

S f + O3 + (1 − y )O2 → S − O + (2 − y )CO2

(5.25)

S − O + O3 → S r + 2CO2

(5.26)

S r + O3 + (1 − y )O2 → S − O + (2 − y )CO2

(5.27)

where Sf and Sr indicate the fresh and reacted soot surface, respectively, and S–O denotes the oxidized soot surface. The stoichiometric parameter y has an average value of 0.45. Reaction 5.25 is assumed to be fast (reaction probability, γ∼10−3), the rate-limiting step is reaction 5.26, the slow decay of ozone (γ∼10−8).

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Extrapolations of these experimental results to typical tropospheric conditions reveal that ozone loss on fresh soot surface by reaction 5.25 is negligible (∼2 ppt), and ozone lifetime calculated from reaction 5.26 is extremely long (∼470 years), indicating that ozone-soot heterogeneous chemistry is unimportant in tropospheric ozone decomposition, even in urban locations with significant soot concentrations. However, soot as UVabsorbing aerosol particles can reduce the actinic flux and thus reduce photochemical ozone production [Dickerson et al., 1997]. 5.3.3.2 SO2 Oxidation on Soot

In the troposphere, soot particles can also play a role in the oxidation of SO2 to sulfate. This oxidation can take place on the surface of soot particles covered with water, by chemisorbed oxygen [Novakov, 1974]. On freshly formed soot particles oxygen is chemisorbed and becomes activated by the adsorption process. The liquid-solid interface area of fresh soot particles is in the order of a few hundred m2 g−1. In the cooling combustion effluent a film of liquid water condenses on the surface of soot particles, into which SO2 can readily dissolve. The resulting aqueous sulfite reacts with the active oxygen at the solid-liquid interface and produce sulfate, which dissolves into the water. This process is very rapid, leading to the formation of quasi-primary sulfate. The surface oxygen complex is then replenished from the gas phase by diffusion of dissolved oxygen and the whole process is repeated. This stage is likely to be diffusion controlled, it therefore constitutes a much slower catalytic oxidation process. As the surface of soot particles gradually degrades by adsorption of sulfate and other species (poisoning), its catalytic activity diminishes. Aged soot is found practically inactive as catalyst for SO2 oxidation [Novakov, 1974]. While this process was later shown to be insignificant in the formation of sulfate in the atmosphere, it is possibly very important in producing internally mixed soot particles [Mamane and Gottlieb, 1989]. These particles as observed over most of the troposphere have fundamentally different surface, optical, and hygroscopic properties from fresh soot particles, which have significant impact on their atmospheric effects such as absorption, lifetime, or reactivity.

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combustion in residential fireplaces, Environmental Science & Technology, 32, 13–22, © 1998 American Chemical Society. Fig. 4.20 Sheesley, R.J., J.J. Schauer, Z. Chowdhury, G.R. Cass, and B.R.T. Simoneit, Characterization of organic aerosols emitted from the combustion of biomass indigenous to South Asia, Journal of Geophysical Research, 108 (D9), 4285, doi: 10.1029/ 2002JD002981. © 2003 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 4.21 reprinted from TracTrends in Analytical Chemistry, 17 (6), Cass, G.R., Organic molecular tracers for particulate air pollution sources, 356–366, © 1998, with permission from Elsevier. Fig. 4.22 reprinted from Atmospheric Environment, 38 (7), Yu, J.Z., H. Yang, H.Y. Zhang, and A.K.H. Lau, Size distributions of water-soluble organic carbon in ambient aerosols and its size-resolved thermal characteristics, 1061–1071, © 2004, with permission from Elsevier. Fig. 5.1 McInnes, L., M. Bergin, J. Ogren, and S. Schwartz, Apportionment of light scattering and hygroscopic growth to aerosol composition, Geophysical Research Letters, 25 (4), 513–516. © 1998 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.2 Bond, T.C., D.S. Covert, J.C. Kramlich, T.V. Larson, and R.J. Charlson, Primary particle emissions from residential coal burning: Optical properties and size distributions, Journal of Geophysical Research, 107 (D21), 8347, doi: 10.1029/2001JD000571. © 2002 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.3 Martins, J.V., P. Artaxo, C. Liousse, J.S. Reid, P.V. Hobbs, and Y.J. Kaufman, Effects of black carbon content, particle size, and mixing on light absorption by aerosols from biomass burning in Brazil, Journal of Geophysical Research, 103 (D24), 32041–32050. © 1998 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.4 Haywood, J.M., and V. Ramaswamy, Global sensitivity studies of the direct radiative forcing effect due to anthropogenic sulfate and black carbon aerosols, Journal of Geophysical Research, 103 (D6), 6043–6058. © 1998 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.5 Abel, S.J., J.M. Haywood, E.J. Highwood, J. Li, and P.R. Buseck, Evolution of biomass burning aerosol properties from an agricultural fire in southern Africa, Geophysical Research Letters, 30 (15), 1783, doi: 10.1029/ 2003GL017342. © 2003 American Geophysical Union. Reproduced by permission of American Geophysical Union. Figs. 5.6 and 5.7 Hegg, D.A., J. Livingston, P.V. Hobbs, T. Novakov, and P. Russell, Chemical apportionment of aerosol column optical depth off the mid-Atlantic coast of the United States, Journal of Geophysical Research, 102 (D21), 25293–25303. © 1997 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.8 Tegen, I., P. Hollrig, M. Chin, I. Fung, D. Jacob, and J. Penner, Contribution of different aerosol species to the global aerosol

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extinction optical thickness: Estimates from model results, Journal of Geophysical Research, 102 (D20), 23895–23915. © 1997 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.9 Schult, I., J. Feichter, and W.F. Cooke, Effect of black carbon and sulfate aerosols on the global radiation budget, Journal of Geophysical Research, 102 (D25), 30107–30117. © 1997 American Geophysical Union. Reproduced by permission of American Geophysical Union. Figs. 5.10 and 5.11 Dick, W.D., P. Saxena, and P.H. McMurry, Estimation of water uptake by organic compounds in submicron aerosols measured during the Southeastern Aerosol and Visibility Study, Journal of Geophysical Research, 105 (D1), 1471–1479. © 2000 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.12 Chughtai, A.R., M.E. Brooks, and D.M. Smith, Hydration of black carbon, Journal of Geophysical Research, 101 (D14), 19505–19514. © 1996 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.13 Ferry, D., J. Suzanne, S. Nitsche, O.B. Popovitcheva, and N.K. Shonija, Water adsorption and dynamics on kerosene soot under atmospheric conditions, Journal of Geophysical Research, 107 (D23), doi:10.1029/2002JD002459. © 2002 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.14 reproduced from Atmospheric Environment, 34 (28), Facchini, M.C., S. Decesari, M. Mircea, S. Fuzzi, and G. Loglio, Surface tension of atmospheric wet aerosol and cloud/fog droplets in relation to their organic carbon content and chemical composition, 4853–4857, © 2000, with permission from Elsevier. Fig. 5.15 Hitzenberger, R., A. Berner, A. KasperGiebl, M. Loeflund, and H. Puxbaum, Surface tension of Rax cloud water and its relation to the concentration of organic material, Journal of Geophysical Research, 107 (D24), doi: 10.1029/2002JD002506. © 2002 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.16 Shulman, M.L., M.C. Jacobson, R.J. Carlson, R.E. Synovec, and T.E. Young, Dissolution behavior and surface tension effects of organic compounds in nucleating cloud droplets, Geophysical Research Letters, 23 (3), 277–280. © 1996 American Geophysical Union. Reproduced by permission of American Geophysical Union. Fig. 5.17 reproduced with permission from Penner, J.E., M. Andreae, H. Annegarn, L. Barrie, J. Feichter, D. Hegg, A. Jayaraman, R. Leaitch, D. Murphy, J. Nganga, and G. Pitari, Aerosols, their Direct and Indirect Effects, in Climate Change 2001: The Scientific Basis. Contribution of Working Group I to the Third Assessment Report of the Intergovernmental Panel on Climate Change, edited by J.T. Houghton, Y. Ding, D.J. Griggs, M. Noguer, P.J.v.d. Linden, X. Dai, K. Maskell, and C.A. Johnson, pp. 881, Cambridge University Press, Cambridge, United Kingdom and New York, Ny, USA. © 2001 Intergovernmental Panel on Climate Change. Fig. 5.18 Lohmann, U., A

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glaciation indirect aerosol effect caused by soot aerosols, Geophysical Research Letters, 29 (4), doi:10.1029/GL014357. © 2002 American Geophysical Union. Reproduced by permission of American Geophysical Union. Table 2.1 reproduced from Atmospheric Environment, 27A (8), Liousse, C., C. Cachier, and S.G. Jennings, Optical and thermal measurements of black carbon aerosol content in different environments: Variation of the specific attenuation cross-section, sigma (σ), 1203–1211, © 1993, with permission from Elsevier. Table 2.2 reprinted with permission from Cachier, H., M. P. Brémond, and P. Buat-Ménard, Determination of atmospheric soot carbon with a simple thermal method, Tellus, 41B (3), 379– 390. © 1989 Blackwell Publishing Ltd. Table 3.1 reprinted from Atmospheric Environment, 35 (30), Hitzenberger, R., A. Berner, H. Glebl, K. Drobesch, A. Kasper-Giebl, M. Loeflund, H. Urban, and H. Puxbaum, Black carbon (BC) in alpine aerosols and cloud water—concentrations and scavenging efficiencies, 5135–5141, © 2001, with permission from Elsevier. Table 3.2 Liousse, C., J.E. Penner, C. Chuang, J.J. Walton, H. Eddleman, and H. Cachier, A global three-dimensional model study of carbonaceous aerosols, Journal of Geophysical Research, 101 (D14), 19411–19432. © 1996 American Geophysical Union. Reproduced by permission of American Geophysical Union. Table 3.3 Andreae, M.O., and P. Merlet, Emissions of trace gases and aerosols from biomass burning, Global Biogeochemical Cycles, 15 (4), 955–966. © 2001 American Geophysical Union. Reproduced by permission of American Geophysical Union. Table 3.4 Cooke, W.F., C. Liousse, H. Cachier, and J. Feichter, Construction of a 1 degrees x 1 degrees fossil fuel emission data set for carbonaceous aerosol and implementation and radiative impact in the ECHAM4 model, Journal of Geophysical Research, 104 (D18), 22137–22162. © 1999 American Geophysical Union. Reproduced by permission of American Geophysical Union. Table 3.5 Griffin, R.J., D.R. Cocker, J.H. Seinfeld, and D. Dabdub, Estimate of global atmospheric organic aerosol from oxidation of biogenic hydrocarbons, Geophysical Research Letters, 26 (17), 2721–2724. © 1999 American Geophysical Union. Reproduced by permission of American Geophysical Union. Tables 3.6 and 3.7 reprinted from Hoffmann, T., J.R. Odum, F. Bowman, D. Collins, D. Klockow, R.C. Flagan, and J.H. Seinfeld, Formation of organic aerosols from the oxidation of biogenic hydrocarbons, Journal of Atmospheric Chemistry, 26 (2), 189–222, © 1997 Kluwer Academic Publishers, with kind permission of Kluwer Academic Publishers and the authors. Table 3.8 Andersson-Sköld, Y., and D. Simpson, Secondary organic aerosol formation in northern Europe: A model study, Journal of Geophysical Research, 106 (D7), 7357–7374. © 2001 American Geophysical Union. Reproduced by permission of American Geophysical Union. Table 4.2 Koch, D., Transport and direct radiative forcing of carbonaceous and sulfate aerosols in the GISS GCM, Journal of

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Geophysical Research, 106 (D17), 20311–20332. © 2001 American Geophysical Union. Reproduced by permission of American Geophysical Union. Table 4.3 reprinted from Atmospheric Environment, 30 (22), Schauer, J.J., W.F. Rogge, L.M. Hildemann, M.A. Mazurek, and G.R. Cass, Source apportionment of airborne particulate matter using organic compounds as tracers, 3837–3855, © 1996, with permission from Elsevier. Table 4.4 reprinted from Saxena, P., and L.M. Hildemann, Water-soluble organics in atmospheric particles: A critical review of the literature and application of thermodynamics to identify candidate compounds, Journal of Atmospheric Chemistry, 24 (1), 57–109, © 1996 Kluwer Academic Publishers, with kind permission of Kluwer Academic Publishers and the authors. Table 4.5 Fuzzi, S., S. Decesari, M.C. Facchini, E. Matta, M. Mircea, and E. Tagliavini, A simplified model of the water soluble organic component of atmospheric aerosols, Geophysical Research Letters, 28 (21), 4079–4082. © 2001 American Geophysical Union. Reproduced by permission of American Geophysical Union. Table 5.1 Rivera-Carpio, C.A., C.E. Corrigan, T. Novakov, J.E. Penner, C.F. Rogers, and J.C. Chow, Derivation of contributions of sulfate and carbonaceous aerosols to cloud condensation nuclei from mass size distributions, Journal of Geophysical Research, 101 (D14), 19483–19493. © 1996 American Geophysical Union. Reproduced by permission of American Geophysical Union. Table 5.2 reprinted from Atmospheric Environment, 29 (7), Lammel, G., and T. Novakov, Water nucleation properties of carbon-black and diesel soot particles, 813–823, © 1995, with permission from Elsevier.

INDEX A Absorption - cross-section, 25 - emission index, 230 - source strength of, 230 Absorptive partitioning - effect of RH on, 115–118 - effect of temperature on, 115–118 ACPM, 34–35 Adsorption artifact (see Positive artifact) Aethalometer, 24–27 Aircrafts - emission of primary OA by, 79–80 Alkenes - SOA formation from, 103 α-pinene - aerosol yield function for, 121–123 - SOA formation from, 91–95 αP-PAN- formation mechanism of, 94 Ambient Carbon Particulate Monitor (see ACPM) Angström-exponent of BC, 229, 231

- surface concentrations of, 58–59 - time-trend in fossil fuel emission of, 67–68 β-pinene- SOA formation from, 95 Bioaerosol, 70–73 Biogenic VOC - annual global flux of, 87 - emission of, 84 Biomass burning - emission factors for, 76–77 - emission of primary organic aerosol from, 77–78 - process of, 76 Biomass burning aerosol - radiative forcing of, 249–251 Black carbon (see BC) Blue haze, 83 BOSS sampler, 18–19 Broadside enhancement, 229 Brown carbon, 47

B

C

Backup filter, 12–14 Bacteria, 71–72 BC - zonal average concentrations of, 65 - concentrations in cloud water and precipitation, 60–61 - definition of, 46–47 - direct radiative forcing of, 252–254 - effect on snow albedo, 246 - emission from aircrafts, 57–58 - emission from biomass burning, 56– 57 - emission from Diesel engines, 57 - global emission estimates, 63–64 - historical records of concentrations of, 61–63 - normalized radiative forcing of, 253– 254 - scavenging efficiency of, 54–55 - scavenging ratio of, 55 - semidirect effect, 247

CCN - contribution of organic aerosol to, 270–272 - organic, modeling of formation of, 273–274 CCN activation - effect of soot on, 284 - effect of wettability, 286–287 - role of limited solubility, 278–284 - role of surface activity, 274–278 - of dicarboxylic acids, 285–286 - of SOA components, 286 Cellulose- concentrations of, 191 Charring in thermal methods, 31–33 - in thermal/optical methods, 39–42 Cloud water - BC concentrations in, 60–61 - organic species in, 218–220 Co-albedo, 234 Coefficient of haze (COH), 24 Colony forming units (CFU), 72

343

344

Carbonaceous Aerosol

D Denuder-filter combination, 18–20 Dicarboxylic acids, 191–195 - as SOA components, 105–106 Diesel- emission of BC, 56–57 Dimer- gas-phase formation of, 93–94 Diterpenoids, 201–202 DRI thermal/optical reflectance (TOR) analyzer, 37–39 Dual filter strategy, 12–14 E EC- definition of, 46 Elemental carbon (see EC) Emission factors - for biomass burning, 76–77 - for fossil fuel combustion, 81 Equilibrium- on quartz filters, 11 Ethylbenzene- SOA formation from, 102 Evaporation artifact (see Negative artifact) F FAC- definition of, 110 Fatty acids (see N-alkanoic acids) Fatty alcohols (see N-alkanols) Film, organic on aerosol, 278 Filter blanks- reduction of, 9 Filter pairs, 14–16 First indirect effect, 289 Foliar emission flux- calculation of, 85 Formaldehyde- in aerosol, 220 Fossil fuel combustion - emission factors for, 81 - emission of primary organic aerosol from, 79–81 Fractal geometry- of soot, 295 Fractional Aerosol Coefficient (see FAC) Fractional Aerosol Yield - definition of, 111 FTIR studies, 174–180 G GC- definition of, 46 Glaciation indirect effect, 293

Graphitic carbon (see GC) H Heterogeneous reactions - acid-catalyzed, 143, 146 - of trimethylbenzene, 144 High-temperature volatility technique, 34 Honeycomb denuder sampler, 20 HULIS - elemental composition of, 177 - formation from isoprene, 144 - formation in multiphase reaction, 145 - isolation of, 177 - MW distribution of, 178 Humic acids, 191 Hygroscopic growth factor - for aerosol scattering, 225–226 Hygroscopic growth - of inorganic compounds- modifying effects of organic compounds, 264– 265 - of organic aerosol, 258–262 - of pure organic compounds, 262–263 - of soot, 265–268 I Impactor- sampling artifacts of, 21 IMPROVE method, 40–41 IN ability of organic aerosol, 288–289 Indirect forcing - contribution of BC to, 293–294 - contribution of organic aerosol to, 291–293 Integrating plate method, 28–29 Integrating sphere method, 29–30 Intercomparison of TC/EC/BC measurement methods, 42–44 Isoprene - atmospheric observations of SOA formation from, 130 - SOA formation from, 97–98, 124 Isotopic measurements (see also Radiocarbon measurements), 182–183

Index

345

L O Lagrangian model - use in SOA estimates, 133 Levoglucosan, 202–203 Lighter fluid effect, 43 Lignans, 201 Lignin derivatives, 199–201 Limited solubility - role in CCN activation, 278–284 Long-chain wax esters, 190 M Malonic acid - formation mechanism of, 106 Mass absorption efficiency (αabs), 25 - for BC, 226 - dependence on wavelength, 229–230 - effect of non-absorbing shell, 213– 233 - inside droplets, 233 - observations, 235 - of organic aerosol, 236–237 - of organic aerosol- dependence on RH, 224–225 Mass scattering efficiency of organic aerosol, 224 Methyltetrols, 130–131 Mixing state- of soot, 53 Morphology- of soot, 49–50 Multiple reflection in fiber filters, 25 N N-alkan-2-ones, 190 N-alkanals, 190 N-alkanes, 184–187 N-alkanoic acids, 187–189 N-alkanols, 189 Negative artifact, 16–18 - in denuder-filter sampler, 17 - temperature dependence of, 17 New particle formation, 127 NIOSH method, 40–41 Nitrogen compounds in aerosol, 215–216 Nopinone- formation mechanism of, 95 Nucleation threshold, 128

OC - normalized radiative forcing of, 255 - direct radiative forcing of, 254–255 - in mass closures, 161–164 - mass concentrations of, 151–153 - mixing state of, 166–168 - modeling globals distribution of, 155–158 - size distribution of, 159–160 - vertical concentration profiles of, 153–155 OC/EC ratio - use in SOA estimates, 131–133 OC/EC split in thermal/optical methods, 41–42 Oleic acid - OH reaction of, 107 - ozonolysis of, 107–108 Oligomerization, 144–146 OM/OC mass conversion ratio, 165 Optical depth- columnar, 237–242 Organic carbon - definition of, 46 Oxalic acid- formation mechanism of, 106 Oxocarboxylic acids - as SOA components, 106 Ozone- loss on soot, 297–298 P PAHs- 196–197 Particle - bounce, 21 - reentrainment, 21 - soot absorption photometer (see PSAP) - trap impactor, 20 Partitioning constant- definition of, 112 Photoacoustic spectrometry, 30–31 Pinonaldehyde - formation mechanism of, 91 Pinonic acid- formation mechanism of, 92 Polycarboxylic acids, 176 Polymerization, 143–146 Polyols, 213 Positive artifact, 10–16 - blank correction for, 11–12 - face velocity dependence, 11

346

Carbonaceous Aerosol

- volatility dependence, 10 Precipitation - BC concentrations in, 60–61 - organic species in, 218–220 Primary organic aerosol - from biomass burning, 77–78 - from fossil fuel combustion, 79–81 PSAP, 27–28 - corrections for scattering in, 28 Q Quartz artifact, 19 Quartz fiber filters, 8–9 - equilibrium on, 9, 11 - saturation of, 13 R Radiative forcing - of biomass burning aerosol, 249–251 - of fossil fuel BC, 248–249 - of OC, 254–255 Radiocarbon measurements, 180–182 Raman - spectroscopy, 31 - spectrum of soot, 51 Reaction artifact, 18 Reactivity-volatility dependence, 110 Reflectance unit of dirt shade (RUDS), 24 Refractive index - of BC, 228 - of organic aerosol, 224 Renoxification reactions, 296–297 Residual organic carbon - definition in TOR analysis, 36 S Saturation- of quartz filter, 13 Scavenging efficiency - of organic species, 217–218 - of BC, 54–55 Sea salt- organic compounds on, 74 Second indirect effect- 289 Semidirect effect- of BC, 247 Sesquiterpenes- ozonolysis products of, 97 SFI- definition of, 109

Single scattering albedo - critical value of, 244 Smoke shade method, 23–24 SOA - CCN activity of, 286 - Formation Index (see SFI) - global formation estimates of, 136– 140 - modeling of size-distribution of, 135 - regional estimate of contribution of, 131 Soil particles, 73 Solubility classification - of organic aerosol, 168–171 Soot - chemistry of, 50–53 - concentrations at high altitudes, 59– 60 - definition of, 45–46 - effect on CCN activation, 284 - fractal geometry of, 295 - glaciation indirect effect of, 293 - graphitic structures in, 51 - mixing state of, 53 - morphology of, 49–50 - nitrogen species in, 52 - pollution history of, 48–49 - Raman spectrum of, 51 - renoxification reactions on, 296–297 - SO2 oxidation on, 300 Source apportionment - based on OC/EC ratio, 131–134 - using tracers, 208–210 Specific attenuation coefficient - historical trend of, 63 Specific attenuation cross section (σ), 25 - variations in, 26–27, 226 Sterols, 203–204 Surface activity - role in CCN activation, 274–278 T Tandem filter strategy (see Dual filter strategy) TC- definition of, 46 Teflon filter, 14–18 Thermal Manganese Oxidation (TMO) method, 34

Index Thermal/optical transmission (TOT) method, 39 Thermograms, 172–174 Thermo-optical reflectance (TOR) method, 35–39 TMO method (see Thermal Manganese Oxidation method) Toluene- SOA formation from, 99 TOR method (see Thermo-optical reflectance (TOR) method) Total carbon (see TC) Tracers - for tobacco smoke, 207 - criteria of, 195–196 - diterpenoids, 201–202 - for charbroilers, 207 - for fossil fuel combustion, 205–206 - for gas-fired applicances, 208 - for terrestrial vegetation, 206–207 - for tire dust, 207–208 - lignans, 201 - lignin derivatives, 199–201 - of cellulose, 202–203 - source apportionment, 208–210 - sterols, 203–204 Twomey effect, 289 Two-product model - for SOA formation, 119–120

347 V VDI methods, 33–34 Visibility degradation, 237 VOC - anthropogenic emission of, 88 - biogenic emission of, 84 Volatile organic carbon - definition in TOR analysis, 36 Volatilization artifact (see Negative artifact) W Water uptake - of organic aerosol, 258–262 - of pure organic compounds, 262–263 - of soot, 265–268 Wax particles- emission of, 70 Wettability - effect on CCN activation of, 286–287 Winnowing effect, 73 WSOC - definition of, 211 - mass size distribution of, 216–217 - nitrogen compounds, 215–216 - retrosynthetic study of, 107–109 - surface activity of, 213 X Xylene, m- SOA formation from, 102

LIST OF ABBREVIATIONS ACPM AED αP-PAN APCI BTX CCN CDNC CFU CMC CN CPI DOC EC EDX EGA EMEP FAC FID FTIR GC GCM GC-MS HMSA HNMR HRTEM HTDMA HULIS IN IP IPCC LWC MW NDIR nss-(sulfate) OC OM ORVOC PAH PAK PAN PC PM

Ambient Carbon Particulate Monitor Aerodynamic Equivalent Diameter α-Pinonyl-Peroxynitrate Atmospheric Pressure Chemical Ionization Benzene, Toluene, Xylene Cloud Condensation Nuclei Cloud Droplet Number Concentration Colony Forming Unit Critical Micelle Concentration Condensation Nuclei (Particle Number Concentration) Carbon Preference Index Dissolved Organic Carbon Elemental Carbon Energy-Dispersive X-Ray Evolved Gas Analysis European Monitoring and Evaluation Programme Fractional Aerosol Coefficient Flame Ionization Detector Fourier-Transform Infrared Graphitic Carbon General Circulation Model Gas Chromatography-Mass Spectrometry Hydroxyalkyl Sulfonate Adduct Proton Nuclear Magnetic Resonance (Spectrometry) High Resolution Transmission Electron Microscopy Hygroscopicity Tandem Differential Mobility Analyzer Humic-Like Substances Ice Nuclei Ice Particles Intergovernmental Panel on Climate Change Liquid Water Content Molecular Weight Non-Dispersive Infrared Non-sea-salt-(sulfate) Organic Carbon Organic Matter Other Reactive Volatile Organic Compounds Polycyclic Aromatic Hydrocarbon Polycyclic Aromatic Ketone Peroxy Acetyl Nitrate Particulate Carbon Particulate Matter

349

350 POC PSAP RH ROG RSD SFI SOA SOAM TC TDMA TMO TOR TOT TSP UV-VIS VOC WINSOC WSOC

Carbonaceous Aerosol Particulate Organic Carbon Particle Soot Absorption Photometer Relative Humidity Reactive Organic Gas Relative Standard Deviation SOA Formation Index Secondary Organic Aerosol Secondary Organic Aerosol Model Total Carbon Tandem Differential Mobility Analyzer Thermal Manganese Oxidation Thermo-Optical Reflectance Thermo-Optical Transmittance Total Suspended Particulate Ultraviolet-Visible Volatile Organic Compounds Water Insoluble Organic Carbon Water-Soluble Organic Carbon

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