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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79

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

80

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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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81

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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Carbonaceous Aerosol

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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83

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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Carbonaceous Aerosol

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 (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

Interaction with Water and its Implications

269

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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271

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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Carbonaceous Aerosol

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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273

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 ( 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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Carbonaceous Aerosol

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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275

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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Carbonaceous Aerosol

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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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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291

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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Carbonaceous Aerosol

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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293

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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295

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

Heterogeneous Reactions on Carbonaceous Aerosol

299

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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Carbonaceous Aerosol

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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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

Copyright Acknowledgements

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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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