Atmospheric Chemistry and Physics

Preface

xxv

1 The Atmosphere

1

1.1 1.2 1.3 1.4 1.5

History and Evolution of the Earth'sAtmosphere 1 Climate 4 The Layersof theAtmosphere 6 Variationof Pressurewith Height in the Atmosphere 9 Large-ScaleMotion of the Atmosphere 10 1.5.1 The GeneralCirculation 10 1.5.2 Troposphere-Stratosphere Transport 14 1.6 TemperatureandWaterVapor 16 1.7 ExpressingtheAmount of a Substancein the Atmosphere 18 1.8 Compositionof theAtmosphere 21 1.9 Radiation 23 1.9.1 SolarandTerrestrialRadiation 26 1.9.2 Absorptionof Radiationby Gases 29 1.10 EnergyBalancefor Earth andAtmosphere 33 1.10.1 SolarVariability 37 1.10.2 Earth'sEnergyBalance 38 1.11 SpatialandTemporalScalesof AtmosphericProcesses 40 Appendix 1 Derivation of the GeostrophicWind Speed 43 References 47 Problems 47 2 Atmospheric Composition, Global Cycles,and Lifetimes 2.1 2.2

2.3

AtmosphericResidenceTimes 50 2.1.1 ResidenceTime 51 Sulfur-ContainingCompounds 55 2.2.1 Dimethyl Sulfide (CH3SCH3) 2.2.2 CarbonylSulfide (OCS) 62 2.2.3 Sulfur Dioxide (SO2) 63 2.2.4 The AtmosphericSulfur Cycle Nitrogen-ContainingCompounds 67 2.3.1 Nitrous Oxide (N2O) 67 2.3.2 Nitrogen Oxides(NOx = NO + 2.3.3 ReactiveOdd Nitrogen (NOy)

49

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4 Chemistry of the Stratosphere 4.1 4.2 4.3

ChapmanMechanism 164 HO x Cycles 171 NO x Cycles 173 4.3.1 N2O StratosphericSourceof NO x 173 4.3.2 NOx Cycles 175 4.4 ClOx Cycles 177 4.5 ReservoirSpeciesand Coupling of the Cycles 180 4.6 StratosphericSpeciesObservationsand Predictions 184 4.6.1 Upper Stratosphere 184 4.6.2 Lower Stratosphere 186 4.7 OzoneHole 189 4.7.1 Polar StratosphericClouds 192 4.7.2 PSCsand the OzoneHole 194 4.7.3 Antarctic OzoneHole Measurements 198 4.7.4 Arctic OzoneHole 199 4.7.5 Summary 202 4.8 Heterogeneous (Nonpolar)StratosphericChemistry 203 4.8.1 Heterogeneous Hydrolysis ofN2Os 203 4.8.2 Effect of Volcanoeson StratosphericOzone 207 4.8.3 Summaryof Midlatitude andTropical Stratospheric OzoneChemistry 210 4.9 TransportBetweenthe Tropical andMidlatitude Stratosphere 210 4.10 Ozone-DepletingPotentialof Halocarbons 212 4.11 Effect of Aircraft Emissionson StratosphericOzone 215 4.12 CarbonylSulfide (OCS)and the Stratospheric Aerosol Layer 216 4.12.1 AtmosphericChemistryofOCS and OCSLifetime 217 4.12.2 StratosphericAerosol Layer 217 4.13 Projectionsof FutureOzoneChange 218 Appendix 4 Sensitivity/UncertaintyAnalysis of Atmospheric ChemicalMechanisms 219 4.A.l SensitivityCoefficients 222 4.A.2 The Direct DecoupledMethod 223 4.A.3 Adjoint Methods 224 4.A.4 Green'sFunctionMethods 224 References 226 Problems 230 5 Chemistry of the Troposphere 5.1 5.2 5.3

Basic PhotochemicalCycle of N02, NO, and03 235 AtmosphericChemistryof CarbonMonoxide andNOx 239 AtmosphericChemistryof FonnaldehydeandNOx 244

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5.15 AtmosphericChemistry(GasPhase)of Sulfur Compounds 314 5.15.1 Sulfur Oxides 314 5.15.2 ReducedSulfur Compounds(Dimethyl Sulfide) 315 5.16 TroposphericChemistryof HalogenCompounds 317 5.16.1 ChemicalCyclesof HalogenSpecies 318 5.16.2 TroposphericChemistryofCFC Replacements: Hydrofluorocarbons(HFCs)and Hydrochlorofluorocarbons (HCFCs) 319 References 323 Problems 331 6 Chemistry of the Atmospheric Aqueous Phase 6.1 6.2 .6.3

6.4 6.5

6.6

Liquid Waterin the Atmosphere 337 6.1.1 Cloud Typesand Liquid WaterContent 338 AbsorptionEquilibria and Henry's Law 340 6.2.1 Gas/Aqueous-Phase Distribution Factor 343 Aqueous-Phase ChemicalEquilibria 344 6.3.1 Water 344 6.3.2 CarbonDioxide/WaterEquilibrium 345 6.3.3 Sulfur Dioxide 348 6.3.4 Ammonia/WaterEquilibrium 353 6.3.5 Nitric Acid/Water Equilibrium 355 6.3.6 Equilibrium of OtherImportantAtmosphericGases 356 Aqueous-Phase ReactionRates 361 S(IV) to S(VI) Transformationand Sulfur Chemistry 363 6.5.1 Oxidation of S(IV) by Dissolved03 363 6.5.2 Oxidation of S(IV) by HydrogenPeroxide 366 6.5.3 Oxidation of S(IV) by OrganicPeroxides 367 6.5.4 UncatalyzedOxidation of S(IV) by O2 368 6.5.5 Oxidationof S(IV) by 02 Catalyzedby Iron andManganese 369 6.5.6 S(IV) Oxidationby the OH Radical 372 6.5.7 Oxidation ofS(IV) by Oxidesof Nitrogen 374 6.5.8 Reactionof DissolvedS02 with HCHO 376 6.5.9 Comparisonof Aqueous-Phase S(IV) Oxidation Paths 378 Aqueous-Phase Nitrite andNitrate Chemistry 380

6.6.1 NOx Oxidation 380 6.7 6.8 6.9

6.6.2 Nitrogen Radicals 381 Aqueous-Phase OrganicChemistry 381 Oxygenand HydrogenChemistry 383 Dynamic Behaviorof Solutionswith Aqueous-Phase ChemicalReactions 384 6.9.1 ClosedSystem 385

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CONTENTS xiii

8.5.2 Aerosol Mobility andDrift Velocity 475 8.5.3 MeanFreePathof an Aerosol Particle 478 8.6 PhoreticEffects 480 8.6.1 Thermophoresis 480 8.6.2 Diffusiophoresis 483 8.6.3 Photophoresis 483 8.7 Aerosol andFluid Motion 484 8.7.1 Motion of a Particlein an IdealizedFlow (900 Corner) 485 8.7.2 StopDistanceand StokesNumber 486 8.8 Diametersof NonsphericalParticles 488 References 488 Problems 489 9 Thermodynamics of Aerosols

491

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ThermodynamicPrinciples 491 9.1.1 Internal Energyand ChemicalPotential 491 9.1.2 The GibbsFreeEnergy,G 493 9.1.3 Conditionsfor ChemicalEquilibrium 495 9.1.4 ChemicalPotentialsof Ideal Gasesand Ideal GasMixtures 499 9.1.5 ChemicalPotentialof Solutions 501 9.1.6 The Equilibrium Constant 506 9.2 Aerosol Liquid WaterContent 507 9.2.1 ChemicalPotentialof Waterin AtmosphericParticles 508 9.2.2 TemperatureDependenceof the DRH 510 9.2.3 Deliquescenceof Mu1ticomponent Aerosols 514 9.2.4 Crystallizationof Single andMulticomponentSalts 519 9.3 Equilibrium VaporPressureOver a CurvedSurface: The Kelvin Effect 519 9.4 Thermodynamicsof AtmosphericAerosol Systems 523 9.4.1 The H2SO4-H2OSystem 523 9.4.2 The Sulfuric Acid-Ammonia-Water System 529 9.4.3 The Ammonia-Nitric Acid-Water System 531 9.4.4 The Ammonia-Nitric Acid-Sulfuric Acid-Water System 537 9.4.5 OtherInorganicAerosol Species 539 References 541 Problem 543

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10 Nucleation 10.1

ClassicalTheory of HomogeneousNucleation: KineticApproach 547 10.1.1 The ForwardRateConstant.Bi 550 10.1.2 The ReverseRateConstantYi 551

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11.2.5 CharacteristicTime for Aqueous-Phase ChemicalReactions 617 11.3 MassTransportandAqueous-Phase Chemistry 617 11.3.1 Gas-Phase Diffusion andAqueous-Phase Reactions 618 11.3.2 Aqueous-Phase Diffusion and Reaction 620 11.3.3 Interfacial MassTransportandAqueous-Phase Reactions 621 11.3.4 Application to the S(IV)-Ozone Reaction 623 11.3.5 Application to the S(IV)-Hydrogen Peroxide Reaction 626 11.3.6 Calculationof Aqueous-Phase ReactionRates 627 11.3.7 An Aqueous-Phase Chemistry/Mass TransportModel 634 11.4 MassTransferto Falling Drops 635 11.5 CharacteristicTime for AtmosphericAerosol Equilibrium 636 11.5.1 Solid Aerosol Particles 636 11.5.2 AqueousAerosol Particles 638 Appendix 11 Solution of the TransientGas-Phase Diffusion ProblemEquations(11.4)to (11.7) 641 References 643 Problems 645 12 Dynamics of Aerosol Populations 12.1

12.2 12.3

12.4 12.5 12.6

MathematicalRepresentations of the Aerosol SizeDistributions 648 12.1.1 DiscreteDistribution 648 12.1.2 ContinuousDistribution 649 Condensation 649 12.2.1 Solutionof the CondensationEquation 652 Coagulation 656 12.3.1 Brownian Coagulation 656 12.3.2 Coagulationin Laminar ShearFlow 664 12.3.3 Coagulationin TurbulentFlow 665 12.3.4 Coagulationfrom GravitationalSettling 665 12.3.5 Brownian Coagulationand ExternalForceFields 666 12.3.6 The CoagulationEquation 672 12.3.7 Solutionof the CoagulationEquation 676 The DiscreteGeneralDynamic Equation 680 The ContinuousGeneralDynamic Equation 682 Evolution of an Aerosol SizeDistribution During Gas-to-ParticleConversion 684 12.6.1 Diffusion-ControlledGrowth 685 12.6.2 SurfaceReaction-ControlledGrowth 686 12.6.3 VolumeReaction-ControlledGrowth 688 12.6.4 DimensionlessSizeSpectraEvolution 689

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14.1.2 TemperatureChangesof a Rising (or Falling) Parcelof Air 770 14.2 AtmosphericStability 772 Problems 775 15 Cloud Physics

777

15.1

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18.3 18.4

Summaryof GaussianPoint SourceDiffusion Fonnulas 923 DispersionParametersin GaussianModels 926 18.4.1 Correlationsfor uyand UzBasedon Similarity Theory 926 18.4.2 Correlationsfor uyand UzBasedon PasquillStability Classes 929 18.5 PlumeRise 931 18.6 Analytical Propertiesof the GaussianPlumeEquation 933 18.7 FunctionalFonns of MeanWind SpeedandEddy Diffusivities 938 18.7.1 MeanWindSpeed 938 18.7.2 Vertical Eddy Diffusion Coefficient Kzz 938 18.7.3 Horizontal Eddy Diffusion Coefficients Kxx and K yy 942 18.8 Solutionsof the Steady-State AtmosphericDiffusion Equation 943 18.8.1 Diffusion from a Point Source 943 18.8.2 Diffusion from a Line Source 944 References 947 Problems 949 19 Dry Deposition

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GeneralRepresentation of Atmospheric Wet RemovalProcesses 997

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22 Radiative Effects of Atmospheric Aerosols: Visibility and Climate 22.1

ScatteringandAbsorptionof Light by Small Particles 1114 22.1.1 RayleighScatteringRegime 1120 22.1.2 GeometricScatteringRegime 1122 22.1.3 ScatteringPhaseFunction 1123 22.1.4 Extinction by an Ensembleof Particles 1123 22.2 Visibility 1126 22.3 Scattering,Absorption,and Extinction Coefficientsfrom Mie Theory 1131 22.4 CalculatedVisibility ReductionBasedon AtmosphericData 1135 22.5 Direct Effect of Aerosolson Climate 1139 22.5.1 Optical Depth 1143 22.5.2 UpscatterFraction 1146 22.5.3 ScatteringModel of an Aerosol Layer 1147 22.5.4 Cooling VersusHeatingof anAerosol Layer 1152 22.5.5 ScatteringModel of anAerosol Layer for a NonabsorbingAerosol 1154 22.5.6 Direct AerosolForcing of Climateby SulfateAerosols 1156 22.5.7 Effect of Mineral Dust on RadiativeForcing of Climate 1160 22.5.8 Effect of Carbonaceous Aerosolson Radiative Forcingof Climate 1166 22.5.9 Internaland ExternalMixtures 1166 22.6 Indirect Effect of Aerosolson Climate 1170 22.6.1 RadiativeModel for a Cloudy Atmosphere 1173 22.6.2 Sensitivityof Cloud Albedo to Cloud Drop NumberConcentration 1175 22.6.3 Relationof Cloud Drop NumberConcentrationto Aerosol Concentrations 1177 22.6.4 Estimatesof Indirect RadiativeForcing of Aerosols 1179 22.7 Summary:Estimatesof Contributionsto RadiativeForcing ~ 1180 22.8 Climate Responseto AnthropogenicAerosol Forcing 1182 Appendix 22 Calculationof ScatteringandExtinction Coefficients by Mie Theory 1184 References 1185 Problems 1190 23 Atmospheric Chemical Transport Models 23.1

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24.6

Alternative Formsof Air Quality Standards 1277 24.6.1 Evaluationof Alternative Formsof the OzoneAir Quality Standardwith 1971Pasadena, California, Data 1279 24.6.2 Selectionof the AveragingTime 1281 24.7 RelatingCurrentandFutureAir Pollutant StatisticalDistributions 1281 References 1283 Problems 1285 Appendix A

Units and Physical Constants

1289

A.l SI BaseUnits 1289 A.2 SI Derived Units 1289 A.3 FundamentalPhysicalConstants 1292 A.4 Propertiesof theAtmosphereandWater 1293 A.5 Units for RepresentingChemicalReactions 1294 A.6 Concentrationsin the AqueousPhase 1295 A.7 Symbolsfor Concentration 1295 References 1295 Appendix B

Rate Constantsof Atmospheric Chemical Reactions

1297

References 1307 Index

1309

The study of atmospheric chemistry as a scientific discipline goes back to the 18th century, when the principal issue was identifying the major chemical components of the atmosphere, nitrogen, oxygen, water, carbon dioxide, and the noble gases. In the late 19th and early 20th centuries attention turned to the so-called trace gases,speciespresent at less than 1 part per million parts of air by volume (1 JLmol per mole). We now know that the atmosphere contains a myriad of trace species, some at levels as low as 1 part per trillion parts of air. The role of trace species is disproportionate to their atmospheric abundance; they are responsible for phenomena ranging from urban photochemical smog, to acid deposition, to stratospheric ozone depletion, to potential climate change. Moreover, the composition of the atmosphere is changing; analysis of air trapped in ice cores reveals a record of striking increases in the long-lived so-called greenhouse gases, carbon dioxide (CO2), methane (CH4), and nitrous oxide (N 20). Within the last century, concentrations of tropospheric ozone (03), sulfate (SO~-), and carbonaceous aerosols in the Northern Hemisphere have increased significantly. There is evidence that all these changes are altering the basic chemistry of the atmosphere. Atmospheric chemistry occurs within a fabric of profoundly complicated atmospheric dynamics. The results of this coupling of dynamics and chemistry are often unexpected: witness the unique combination of dynamical forces that lead to a wintertime polar vortex over Antarctica, with the concomitant formation of polar stratospheric clouds that serve as sites for heterogeneous chemical reactions involving chlorine compounds resulting from anthropogenic chlorofluorocarbons-all leading to the near total depletion of stratospheric ozone over the South Pole each spring; witness the nonlinear, and counterintuitive, dependence of the amount of ozone generated by reactions involving hydrocarbons and oxides of

nitrogen(NOx) at the urbanandregionalscale-although both hydrocarbonsandNOx are ozoneprecursors,situationsexist wherecontinuing to emit more and more NO x actually leads to less ozone. The chemical constituents of the atmosphere do not go through their life cycles independently; the cycles of the various species are linked together in a complex way. Thus a perturbation of one component can lead to significant, and nonlinear, changes to other components and to feedbacks that can amplify or damp the original perturbation. In many respects, at once both the most important and the most paradoxical trace gas in the atmosphere is ozone (03). High in the stratosphere ozone screens living organisms from biologically harmful solar ultraviolet radiation; ozone at the surface, in the troposphere, can produce adverse effects on human health and plants when present at levels elevated above natural. At the urban and regional scale, significant policy issues concern how to decreaseozone levels by controlling the ozone precursors-hydrocarbons and oxides of nitrogen. At the global scale, understanding both the natural ozone chemistry of the troposphere and the causes of continually increasing background tropospheric ozone levels is a major goal.

xxv

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JAxx

PREFACE xxvii

for the benefit of the studentand the instructor.The designation(A) indicatesa problem that involves a straightforwardapplicationof materialin the text. Thoseproblemsdenoted (B) requiresomeextensionof the ideasin the text. Problemsdesignated(C) encouragethe readerto apply conceptsfrom the book to currentproblemsin atmosphericscienceandgo somewhatbeyondthe level of (B) problems.Finally, thoseproblemsdenoted(D) are of a degreeof difficulty correspondingto (C) but generallyrequiredevelopmentof a computer programfor their solution. This book isa successorto John H. Seinfeld'sAtmosphericChemistryand Physicsof Air Pollution (Wiley, 1986),which has been widely usedfor a decade.Substantialadditions havebeenmadein virtually all areascoveredby the earlier book, reflectingthe significant increasein our understandingof atmosphericprocessesthat hasemergedover the last decadeand reflecting the truly global scopeof the scienceof atmosphericchemistry and physics. We extendsincereappreciationto Marta GoodmanandLaura Shaheenwho providedso much valuableassistancein making this book a reality and to Cecilia Lin who draftedthe figuresso skillfully. JOHN H. SEINFELD SPYROSN. PANDIS

1 1.1 HISTORY AND EVOLUTION OF THE EARTH'S ATMOSPHERE It is generally believed that the solar system condensed out of an interstellar cloud of gas and dust, referred to as the "primordial solar nebula," about 4.6 billion years ago. The atmospheres of the Earth and the other terrestrial planets, Venus and Mars, are thought to have formed as a result of the release of trapped volatile compounds from the planet itself. The early atmosphere of the Earth is believed to have been a mixture of carbon dioxide (COz), nitrogen (Nz), and water vapor (HzO), with trace amounts of hydrogen (Hz), a mixture similar to that emitted by present day volcanoes. The composition of the present atmosphere bears little resemblance to the composition of the early atmosphere. Most of the water vapor that outgassed from the Earth's interior condensed out of the atmosphere to form the oceans. The predominance of the CO2 that outgassed formed sedimentary carbonate rocks after dissolution in the ocean. It is estimated that for each molecule of COz presently in the atmosphere, there are about 105 COz molecules incorpor~ted as carbonates in sedimentary rocks. Since Nz is chemically inert, non-water soluble, and noncondensable, most of the outgassed Nz accumulate~ in the atmosphere over geologic time to become the atmosphere's most abundant constituent. The early atmosphere of the Earth was a mildly reducing chemical mixture, whereas the present atmosphere is strongly oxidizing. The dramatic rise of oxygen (Oz) as an atmospheric constituent over time was the result of the production of Oz as a by-product of photosynthetic activity. It has been estimated that the current level of Oz in the atmosphere was achieved approximately 400 million years ago (Cloud, 1983). The present level of Oz is maintained by a balance between production from photosynthesis and removal through respiration and decay of organic carbon. If Oz were not replenished by photosynthesis, the reservoir of surface organic carbon would be completely oxidized in about 20 years, at which time the amount of Oz in the atmosphere would have decreased by less than 1% (Walker, 1977). In the absenceof surface organic carbon to be oxidized, weathering of sedimentary rocks would consume the remaining O2 in the atmosphere, but it would take approximately 4 million years to do so. The Earth's atmosphere is composed primarily of the gases Nz (78%), Oz (21%), and Ar (1 %), whose abundancesare controlled over geologic time scales by the biosphere, uptake and release from crustal material, and degassing of the interior. Water vapor is the next most abundant constituent; it is found mainly in the lower atmosphere and its concentration is highly variable, reaching concentrations as high as 3%. Evaporation and precipitation control its abundance. The remaining gaseous constituents, the trace gases, comprise less than 1% of the atmosphere. These trace gases playa crucial role in the Earth's radiative

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OF THE EARTH'S ATMOSPHERE

3

The atmosphere is the recipient of many of the products of our technological society. These effluents include products of combustion of fossil fuels and of the development of new synthetic chemicals. Historically these emissions can lead to unforeseen consequences in the atmosphere. Classical examples include the realization in the 1950s that motor vehicle emissions could lead to urban smog and the realization in the 1970s that emissions of chlorofluorocarbons from aerosol spray cans and refrigerators could cause the depletion of stratospheric ozone. The chemical fates of trace atmospheric species are often intertwined. The life cycles of the trace species are inextricably coupled through the complex array of chemical and physical processesin the atmosphere. As a result of these couplings, a perturbation in the concentration of one species can lead to significant changes in the concentrations and lifetimes of other trace species and to feedbacks that can either amplify or damp the original perturbation. An example of this coupling is provided by methane. Methane is the predominant organic molecule in the troposphere and it is the sec.ondmost important greenhouse gas after CO2. Methane sources such as rice paddies and cattle can be estimated and are increasing. Methane is removed from the atmosphere by reaction with the hydroxyl (OH) radical, at a rate that depends on the atmospheric concentration of OR. But, the OH concentration dependson the amount of carbon monoxide (CO), which itself is a product of C~ oxidation as well as a result of fossil fuel combustion and biomass burning. The hydroxyl concentration also depends on the concentration of ozone and oxides of nitrogen. Change in C~ can affect the total amount of ozone in the troposphere, so methane itself affects the concentration of the species, OR, that governs its removal. Depending on their atmospheric lifetime, trace species can exhibit an enormous range of spatial and temporal variability. Relatively long-lived species have a spatial uniformity such that a handful of strategically located sampling sites around the globe are adequate to characterize their spatial distribution and temporal trend. As species lifetimes become shorter, their spatial and temporal distributions become more variable. Urban areas,for example, can require tens of monitoring stations over an area of hundreds of square kilometers in order to characterize the spatial and temporal distribution of their atmospheric components. The extraordinary pace of the recent increases in atmospheric trace gases can be seen when current levels are compared with those of the distant past. Such comparisons can be made for CO2 and C~, whose histories can be reconstructed from their concentrations in bubbles of air trapped in ice in such perpetually cold places as Antarctica and Greenland. With gasesthat are long-lived in the atmosphere and therefore distributed rather uniformly over the globe, such as CO2 and C~, polar ice core samples reveal global average concentrations of previous eras. Analyses of bubbles in ice cores show that CO2 and C~ concentrations remained essentially unchanged from the end of the last ice age some 10,000 years ago until roughly 300 years ago, at mixing ratios close to 260 ppm by volume and 0.7 ppm by volume, respectively. (See Section 1.4 for discussion of units.) About 300 years ago methane levels began to climb, and about 100 years ago levels of both gases began to increase markedly. Before the large-scale production of chlorofluorocarbons the natural level of chlorine in the stratosphere was 0.6 part per billion (ppb) by volume; now it is 3 ppb, a factor of 5 increase. Activities of humans account for most of the rapid changes in the trace gases over the past 200 years-combustion of fossil fuels (coal and oil) for energy and transportation, in-

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century. It has been estimated that a doubling of CO2 from its pre-Industrial Revolution mixing ratio of 280 ppm by volume could lead to a rise in average global temperature of 1.5 to 4.5°C. A 2°C warming would produce the warmest climate seen on Earth in 6000 years. A 4.5°C rise would place the world in a temperature regime last experienced in the Mesozoic Era-the age of dinosaurs. Although an average global ~arming of a few degrees does not sound like much, it could create dramatic changes in climatic extremes. It has been estimated, for example, that, in the event of an average global warming of 1.7°C, the frequency of periods of 5 days or more exceeding 35°C (95°F) in the Com Belt of the United States would increase threefold. Such conditions at critical stages of the growing season are known to harm com and lead to reduced yields. With a doubling of CO2, the number of days exceeding 38°C (100°F) and nights above 27°C (80°F) have been estimated to rise dramatically in many major American cities. Changes in the timing and amount of precipitation would almost certainly occur with a warmer climate. Soil moisture, critical during planting and early growth periods, will change. Some regions would probably become more productive, others less so. The North American Grain Belt, according to at least one climate model, will shift northward into Canada as warming produces hotter, drier conditions in the American Midwest. Of all the effects of a global warming, perhaps none has captured more attention than the prospect of rising sea levels. This would result from the melting of land-based glaciers and volume expansion of ocean water as it warms. Prevailing opinion is that a sea level rise of about 0.5 m could occur by 2100. In the most dramatic scenario, the West Antarctic ice sheet, which rests on land that is below sea level, could slide into the sea if the buttress of floating ice separating it from the ocean were to melt. This would raise the average sea level 5 to 6 m (Bentley, 1997). Even a 0.3 m (1 foot) rise would have major effects on the erosion of coastlines, salt water intrusion into the water supply of coastal areas, flooding of marshes, and inland extent of surges from large storms. To systematically approach the complex subject of climate, the scientific community has divided the problem into two major parts, climate forcings and climate responses. Climate forcings are changes in the energy balance of the Earth that are imposed upon it; forcings are measured in units of heat flux-watts per square meter (W m-2). An example of a forcing is a change in energy output from the sun. Responsesare the meteorological results of these forcings, reflected in temperatures, rainfall, extremes of weather, sea level height, and so on. Much of the variation in the predicted magnitude of potential climate effects resulting from the increase in greenhouse gas levels hinges on estimates of the size and direction of various feedbacks that may occur in response to an initial perturbation of the climate. Negative feedbacks have an effect that damps the warming trend; positive feedbacks reinforce the initial warming. One example of a greenhouse warming feedback mechanism involves water vapor. As air warms, each cubic meter of air can hold more water vapor. Since water vapor is a greenhouse gas, this increased concentration of water vapor further enhances greenhouse warming. In turn, the warmer air can hold more water, and so on. This is an example of a positive feedback, providing a physical mechanism for multiplying the original impetus for change beyond its initial amount. Some mechanisms provide a negative feedback, which decreasesthe initial impetus. For example, increasing the amount of water vapor in the air may lead to forming more clouds. Low-level, white clouds reflect sunlight, thereby preventing sunlight from reaching the

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9

'HE LAYERS OF THE ATMOSPHERE

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Stratosphere. Extends from the tropopause to the stratopause (-45 to 55 kIn altitude); temperature increases with altitude, leading to a layer in which vertical mixing is slow. Mesosphere. Extends from the stratopause to the mesopause (-80 to 90 kIn altitude); temperature decreaseswith altitude to the mesopause, which is the coldest point in the atmosphere; rapid vertical mixing. Thermosphere. The region above the mesopause; characterized by high temperatures as a result of absorption of short wavelength radiation by N2 and O2; rapid vertical mixing. The ionosphere is a region of the upper mesosphereand lower thermosphere where ions are produced by photoionization. Exosphere. The outermost region of the atmosphere (>500 kIn altitude) where gas molecules with sufficient energy can escape from the Earth's gravitational attraction.

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3'M3HdSOW.LV 3ffi

8

VARIATION OF PRESSURE WITH HEIGHT IN THE ATMOSPHERE

9

Absorption of solar ultraviolet radiation by 03 causesthe temperature in the stratosphereto be much higher than expected, based on simply extending the troposphere's lapse rate into the stratosphere.

1.4 VARIATION OF PRESSUREWITH HEIGHT IN THE ATMOSPHERE The variation of pressurewith height in the atmospherecan be addressedwith the hydrostaticequation,! dp(z)

dz

=

.p(z)g

(1.1)

where p(z) is the massdensityof air at height z and g is the accelerationdue to gravity. From the ideal gaslaw,

p(z)=

MairP(Z) RT(z)

(1.2)

where Mairis the averagemolecularweight of air (28.97g mol-I). Thus dp(z)

dz

-

Mairgp(z) RT(z)

(1.3)

which we can rewrite as d lnp(z) - -~

dz

where H(z)

= RT(z)/

-

H(z)

(1.4)

Mairg is a characteristic length scale for decrease of pressure with

height. The temperaturein the atmospherevariesby lessthan a factor of 2, while the pressure changesby six ordersof magnitude(seeTableA.8). If the temperaturecan be takento be approximatelyconstant,just to obtain a simple approximateexpressionfor p(z), then the

'Units of pressure. Because instruments for measuring pressure, such as the manometer, often contain mercury, commonly used units for pressure are based on the height of the mercury column (in millimeters) that the gas pressure can support. The unit mm Hg is often called the Torr in honor of the scientist, Evangelista Torricelli. A related unit for pressure is the standard atmosphere (abbreviated atm): 1 standard atmosphere

= 1 atm = 760

mm Hg = 760 Torr

The unit of pressure in the International System of Units (SI) is newtons per meter squared (N m-'), which is called the pascal (Pa). In terms of pascals, the standard atmosphere is 1.01325 x 10s Pa. Another commonly used unit of pressure in atmospheric science is the millibar (mbar), which is equivalent to the hPa (see Tables A.5 and AR) Thp .t"nti"rti "tmn.nhprP i. 101, ?~ mh"r

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to the tropics must pass through more atmosphere and intercept a larger surface area. The uneven distribution of energy resulting from latitudinal variations in insolation and from differences in absorptivity of the Earth's surface leads to the large-scale air motions of the Earth. In particular, the tendency to transport energy from the tropics toward the pola( regions, thereby redistributing energy inequalities on the Earth, is the overall factor governing the general circulation of the atmosphere. In order to visualize the nature of the general circulation of the atmosphere, we can think of the atmosphere over either hemisphere as a fluid enclosed within a long, shallow container, heated at one end and cooled at the other. Because the horizontal dimension of the "container" is so much greater than its vertical dimension, the curvature of the Earth can be neglected, and the container can be considered to be rectangular. If such a container were constructed in the laboratory and the ends differentially heated as described above, one would observe a circulation of the fluid, consisting of rising motion along the heated wall and descending motion along the cooled wall, flow in the direction of warm to cold at the top of the box, and flow in the direction of cold to warm along the bottom of the box. In the atmosphere, then, the tendency is for warm tropical air to rise and cold polar air to sink, with poleward and equatorward flows to complete the circulation. However, the general circulation of the atmosphere is not as simple as just described. Another force arises because of the motion of the Earth, the Coriolis force. At the Earth's surface an object at the equator has a greater tangential velocity than one in the temperate zones. Air moving toward the south cannot acquire an increased eastward (the Earth rotates from west to east) tangential velocity as it moves south and thus, to an observer on the Earth, appears to acquire a velocity component in the westward direction. Thus air moving south in the Northern Hemisphere appears to lag behind the Earth. To an observer on the Earth it appears that the air has been influenced by a force in the westward direction. To an observer in space, it would be clear that the air is merely trying to maintain straight-line motion while the Earth turns below it. Friction between the wind and the ground diminishes this effect in the lower atmosphere. From the standpoint of air motion, the atmosphere can be segmented vertically into two layers. Extending from the ground up to about 1000 m is the planetary boundary laye!; the zone in which the effect of the surface is felt and in which the wind speed and direction are governed by horizontal pressure gradients, shear stresses, and Coriolis forces. Above the planetary boundary layer is the geostrophic laye!; in which only horizontal pressure gradients and Coriolis forces influence the flow. . To predict the general pattern of macroscale air circulation on the Earth we must consider both the tendency for thermal circulation and the influence of Coriolis forces. Figure 1.2 shows the nature of the general circulation of the atmosphere. At either side of the equator is a thermal circulation, in which warm tropical air rises and cool northern air flows toward the equator. The circulation does not extend all the way to the poles becauseradiative cooling of the upper northward flow causesit to subside (fall) at about 30° Nand S latitude. The Coriolis force acting on these cells leads to easterly winds, called the trade winds. The same situation occurs in the polar regions, in which warm air from the temperate zones moves northward in the upper levels, eventually cooling by radiation and subsiding at the poles. The result is the polar easterlies. In the temperate regions, between 40° and 55° latitude, influences of both tropical and polar regions are felt. The major feature of the temperate regions is large-scale weather systems, which results in the circulation shown in Figure 1.2. The surface winds in the Northern Hemisphere are westerlies because of the Coriolis force.

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the averagetemperatureexceeds295 K. The annualmeantemperaturestructureis largely symmetricabout the equator(although the highest temperatureoccursat about 10° N), with the differencein meansurfacetemperaturebetweenthe tropics and the poles about 35°C.At the top of the troposphere,aboutthe 200 mbar level, the meridionaltemperature gradientbeginsto reverse;at 100 mbar, for example,temperatureincreasespolewardby about20°C. Watervapor is distributedthroughoutthe lower troposphere,at highly variablelevels. The watervaporcontentof the atmospherecanbe expressedin a variety of ways: 1. Mole (or volume) mixing ratio--moles of water vaporper mole of air (seeSection 1.7). 2. Amount of water vaporto dry air by mass[g H2O (kg dry air)-']. 3. Specifichumidity-proportion of watervapor to total air [g H2O (kg air)-']. 4. Relativehumidity-ratio of the specifichumidity to the maximumspecifichumidity possibleat a given temperatureandpressure(dimensionless).(SeeSection1.7.) 5. Massconcentrationg H2O (m3air)-'. 6. Massmixing ratio g H2O (g air)-'. Figure 1.5showsthe zonalmeanspecifichumidity for the period 1980to 1989.Maximum specifichumidities are reachedin the tropics, about 16 g kg-I. By the 500 mbar altitude, the valueover the tropics hasdecreasedto 2 g kg-I. This value of 2 g kg-I also holdsin the northernand southernpolar regions.

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21

OF THE ATMOSPHERE

The atmosphere is composed primarily of nitrogen, oxygen, and several noble gases, the concentrations of which have remained remarkably fixed over time. Also present are a number of trace gases that occur in relatively small and sometimes highly variable amounts. In spite of its apparent unchanging nature, the atmosphere is in reality a dynamic system, with its gaseous constituents continuously being exchanged with vegetation, the oceans, and biological organisms. The so-called cycles of the atmospheric gasesinvolve a number of physical and chemical processes. Gases are produced by chemical processes within the atmosphere itself, by biological activity, volcanic exhalation, radioactive decay, and human industrial activities. Gases are removed from the atmosphere by chemical reactions in the atmosphere, by biological activity, by physical processes in the atmosphere (such as particle formation), and by deposition and uptake by the oceans and earth. The average lifetime of a gas molecule introduced into the atmosphere can range from secondsto millions of years, depending on the effectiveness of the removal processes. Most of the species considered air pollutants (in a region in which their concentrations exceed substantially the normal background levels) have natural as well as man-made sources. Therefore, in order to assessthe effect man-made emissions may have on the atmosphere as a whole, it is essential to understand the atmospheric"cycles of the trace gases,including natural and anthropogenic sources as well as predominant removal mechanisms. The important atmospheric gasesare listed in Table 1.1 arranged according to the nature of their global cycles. The total quantity of a species both in the atmosphere and dissolved

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RADIATION

23

erals at the Earth's surface. Thus the oxygen concentration in the atmosphere is probably a result of both accumulation and geochemical cycles. While in the air, a substancecan be chemically altered in one of two ways. First, the sunlight itself may contain sufficient energy to break the molecule apart, a so-called photochemical reaction. The more frequently occurring chemical alteration, however, takes place when two molecules interact and undergo a chemical reaction to produce new species. Atmospheric chemical transformations can occur homogeneously or heterogeneously. Homogeneous reactions occur entirely in one phase; heterogeneous reactions involve more than one phase, such as a gas interacting with a liquid or with a solid surface. During transport through the atmosphere, all but the most inert substancesare likely to participate in some form of chemical reaction. This process can transform a chemical from its original state, the physical (gas, liquid, or solid) and chemical form in which it first enters the atmosphere, to another state that may have either similar or very different characteristics. Transformation products can differ from their parent substance in their chemical properties, toxicity, and other characteristics. These products may be removed from the atmosphere in a manner very different from that of their precursors. For example, when a substancethat was originally emitted as a gas is transformed into a particle, the overall removal is usually hastened since particles often tend to be removed from the air more rapidly than gases. In spite of the fact that the atmosphere is composed predominantly of relatively inert molecules such as Nz and Oz, it is actually a rather efficient oxidizing medium. One reason for the atmosphere's oxidizing capacity arises because the atmosphere contains minute amounts of very reactive molecular fragments, called free radicals. The most important free radical in the chemistry of the troposphere is the hydroxyl (OH) radical, which reacts with nearly every molecular species in the atmosphere. In addition, the atmosphere contains trace amounts of species less reactive than free radicals but nonetheless reactive enough to attack a variety of airborne compounds. Ozone (03) is one important oxidizer, which also participates in the formation of the hydroxyl radical. Once emitted, species are converted at various rates into substances generally characterized by higher chemical oxidation statesthan their parent substances.Frequently this oxidative transformation is accompanied by an increase in polarity (and hence water solubility) or other physical and chemical changes from the precursor molecule. An example is the conversion of sulfur dioxide (SOz) into sulfuric acid (HzS04). Sulfur dioxide is moderately water soluble, but its oxidation product, sulfuric acid, is so water soluble that even single molecules of sulfuric acid in air immediately become associated with water molecules. The demise of one substance through a chemical transformation can become another species in situ source. In general, then, a species emitted into the air can be transformed by a chemical process to a product that may have markedly different physicochemical properties and a unique fate of its own.

1.9 RADIATION Basically all the energy that reaches the Earth comes from the Sun. The absorption and loss of radiant energy by the Earth and the atmosphere are almost totally responsible for the Earth's weather, both on a global and local scale. The average temperature on the Earth re-

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As can be seenfrom Figures 1.7 and 1.8, the higher the temperature, the greater is the emissive power (at all wavelengths). We also see that, as temperature increases, the maximum value of FB(A) moves to shorter wavelengths. The wavelength at which the maximum amount of radiation is emitted by a blackbody is found by differentiating (1.14) with respect to A, setting the result equal to zero, and solving for A. The result with A expressed in nm and T in kelvin units is

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Thushot bodiesnot only radiatemore energythan cold ones,they do so at shorterwavelengths.The wavelengthsfor the maximaof solar andterrestrialradiationare480 nm and about 10,000nm, respectively.The Sun, with an effective surfacetemperatureof about 6000 K, radiatesabout2 X 105more energyper squaremeterthan the Earth at 300 K.

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If (1.14) is integratedover all wavelengths,the total emissivepower FB (W m-1 of a blackbody is found to be

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Absorption of Radiation by Gases

Absorption of radiation by gases is one of the most important aspects of both global meteorology and atmospheric chemistry. The solar spectrum is radically altered by absorption as the radiation traverses the atmosphere. It is important to note that the molecules that are responsible for the most pronounced absorption of both solar and terrestrial radiation are the minor constituents of the atmosphere, not N2 and O2. Thus ozone in the upper atmosphere effectively absorbs all solar radiation below 290 nm, whereas water vapor and carbon dioxide absorb much of the long-wave terrestrial radiation. The most significant absorbing gasesin the atmosphere are O2, 03, H2O, and CO2, Figure 1.9 shows the solar irradiance at the top of the atmosphere and that at sea level. The absorption spectra are quite complex, but they do indicate that absorption is so strong in some spectral regions that no solar energy in those regions reaches the surface of the Earth. For example, absorption by O2 and 03 is responsible for removal of practically all the incident radiation with wavelengths shorter than 290 nm. However, atmospheric absorption is not strong from 300 to about 800 nm, forming a "window" in the spectrum. About 40% of the solar energy is concentrated in the region of 400 to 700 nm. Water vapor absorbs in a complicated way, and mostly in the region where the Sun's and Earth's radiation overlap. From 300 to 800 nm, the atmosphere is essentially transparent. From 800 to 2000 nm, terrestrial long-wave radiation is moderately absorbed by water vapor in the atmosphere. Table 1.2 summarizes the attenuation of solar radiation by the atmosphere. Figure 1.10 shows the penetration of radiation as a function of height in the atmosphere. Wavelengths shorter than about 100 nm are absorbed by O2 and N2 and do not penetrate below 100 krn. O2 absorbs strongly in the range 100 to 175 nm, the so-called Schumann-Runge continuum, and also in the range 175 to 200 nm, the Schumann-Runge bands. Wavelengths in the 200 to 245 nm range are absorbed in the stratosphere, mainly by O2 (the weak Herzberg continuum). Wavelengths between 200 and 230 nm do penetrate as low as 30 krn altitude. Ultraviolet absorption by ozone, which peaks near 254 nm (the Hartley band of 03 absorption), attenuates solar radiation over the entire range of 230 to 300 nm. As a result, solar radiation of wavelengths shorter than about 290 to 300 nm does not reach the Earth's surface. Why molecules absorb in particular regions of the spectrum can be determined only through quantum chemical calculations. In general, the geometry of the molecule explains, for example, why H2O, CO2, and 03 interact strongly with radiation above 400 nm but N2 and O2 do not. In H2O, for instance, the center of the negative charge is shifted toward the oxygen nucleus and the center of positive charge toward the hydrogen nuclei, leading to a separation between the centers of positive and negative charge, a so-called electric dipole moment. Molecules with dipole moments interact strongly with electromagnetic radiation

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because the electric field of the wave causes oppositely directed forces and therefore accelerations on electrons and nuclei at one end of the molecule as compared with the other. Similar arguments hold for ozone; however, nitrogen and oxygen are symmetric and thus are not strongly affected by radiation above 400 nm. The CO2 molecule is linear but can easily be bent, leading to an induced dipole moment. A transverse vibrational mode exists for CO2 at 15 JLm,just where the Earth emits most of its infrared radiation. Considering the outgoing long-wave infrared radiation, the spectral region from about 7 to 13 JLmis also a window region; nearly 80% of the radiation emitted by the Earth in this region escapesto space. Most of the non-CO2 greenhouse gases, including 03, CH4, N2O, and the chlorofluorocarbons, all have strong absorption bands in this window region. For this reason, relatively small changes in the concentrations of these gasescan produce a significant change in the net radiative flux. As the concentration of a greenhouse gas continues to increase, it can absorb more of the radiation in its energy bands. Once an absorption wavelength becomes saturated, further increases in the concentration of the gas have less and less effect on radiative flux. This is called the band saturation effect. For CO2, for example, the 15 JLmband is already close to saturated. In addition, if a gas absorbs at wavelengths that are also absorbed by other gases, then the effect of increasing concentrations on radiative flux is less than in the absence of band overlap. For example, there is significant overlap between some of the absorption bands of CH4 and N2O; this overlap must be carefully accounted for when calculating the effect of these gases on radiative fluxes. Even with the band saturation effect, it is incorrect to conclude that because there is already so much CO2 in the atmosphere, more CO2 can have no additional effect on absorption of outgoing radiation. When gases are present in small concentrations, doubling the concentration of the gas will approximately double its absorption. When an absorbing gas is present in high concentration the effect of further addition is not one-to-one but it is not

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in this portion of the spectrum. Away from this band, however, where CO2 is less strongly absorbing, the increase in CO2 does have an effect. As more and more CO2 is added to the atmosphere, more of its spectrum will become saturated, but there will always be regions of the spectrum that remain unsaturated and thus capable of continuing to absorb infrared radiation. For example, the 10 .urn absorption band is about 106times weaker than the peak of the 15 .urn band, but its contribution to the irradiance change in the lower frame is important. And as CO2 concentrations increase, the importance of the 10 .urn band will continue to increase relative to the 15 .urn band.

1.10 ENERGY BALANCE FOR EARTH AND ATMOSPHERE The Earth's climate is controlled by the amount of solar radiation intercepted by the planet and the fraction of that energy that is absorbed. The flux density of solar energy, integrated over all wavelengths, on a surface oriented perpendicular to the solar beam at the Earth's orbit is about 1370 W m-2. This is called the solar constant.3 Let the solar constant be denoted by So = 1370 W m-2. The cross-sectional area of the Earth that intercepts the solar beam is 1TR2,where R is the Earth's radius. The surface area of the Earth that receives the radiation is 41TR2.Thus the fraction of the solar constant received per unit area of the Earth is (1TR2 /41TR1 = 1/4 of the solar constant, about 343 W m-2. Of this incoming solar radiation, a fraction is reflected back to space; that fraction, which we can denote by Rp, is the global mean planetary reflectance or albedo. Rp is about 0.3 (Ramanathan, 1987; Ramanathanet al., 1989). Contributing to Rp are clouds, scattering by air molecules, scattering by atmospheric aerosol particles, and reflection from the surface itself, the surface albedo (the surface albedo is denoted as Rs). The fraction 1 - Rp represents that fraction of solar short-wave radiation that is absorbed by the Earth-atmosphere system. For Rp = 0.3, this corresponds to about 240 W m-2. This amount is matched, on an annual and global average basis, by the long-wave infrared radiation emitted from the Earth-atmosphere system to space(Figure 1.12). The infrared radiative flux emitted at the surface of the Earth, about 390 W m-2, substantially exceeds the outgoing infrared flux of 240 W m-2 at the top of the atmosphere. Clouds, water vapor, and the greenhouse gases(GAGs) both absorb and emit infrared radiation. Since these atmospheric constituents are at temperatures lower than that at the Earth's surface, they emit infrared radiation at a lower intensity than if they were at the temperature of the Earth's surface and therefore are net absorbers of energy. The equilibrium temperature of the Earth can be estimated by a simple model that equatesincoming and outgoing energy (Figure 1.13). Incoming solar energy at the surface of the Earth is Fs

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For Rp = 0.30, this equation gives Te = 255 K. If the Earth were totally devoid of clouds,

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1.10.1

Solar Variability

The amount of solar radiation reaching Earth and Earth's changing orientation to the Sun have been the major causes for climatic change throughout its history. If the Sun's radiation intensity declined 5 to 10% and there were no other compensating factors, ice would engulf the planet in less than a century. Although no theory exists to predict future changes in solar output, the effect of changes in Earth's orbit as it travels around the Sun is beginning to be understood. During the past million years, Earth has experienced 10 major and 40 minor episodes of glaciation. All appear to have been controlled by three so-called orbital elements that vary cyclically over time. First, Earth's tilt changes from 22° to 24.5° and back again every 41,000 years. Second, the month when Earth is closest to the Sun also varies over cycles of 19,000 and 24,000 years. Currently, Earth is closest to the Sun in January. This month-of-closest-approach factor can make a difference of 10% in the amount of solar radiation reaching a particular location in a given season. Last, the shape of Earth's orbit varies from being nearly circular to being more elliptical with a period of 100,000 years. The climatic cycles caused by these orbital factors are called Milankovitch cycles after the Serbian mathematician Milutin Milankovitch, who first described them in 1920. Superimposed on the Milankovitch cycles are changes in the Sun that occur over days or months or a few years. Over the period 1979 to 1990, for example, total solar irradiance varied by about 0.1 % (Hickey et al., 1988; Willson and Hudson, 1988). Even though studies of ocean cores have shown that these orbital changes are the principal determinant of the times of glaciation, the exact mechanisms by which Earth responds to the orbital changes have not been established. Orbital changes alone appear not to have caused the vast climate shifts associated with glaciation and deglaciation. Feedbacks, such as changes in Earth's reflectivity, amount of particles in the atmosphere, and the carbon dioxide and methane content of the atmosphere, act together with orbital changes to enhance global warming and cooling. The levels of carbon dioxide and

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from the Earth, andnonradiativeprocesses.Of the 100units of incoming solar flux: -26 units are absorbedwithin the atmosphere(-22 by cloud-free air and -4 by clouds)

- 30unitsarereflectedbackto space 4 (- 7 from the cloud-free atmosphere, -17 from the cloudy atmosphere,and -6 from the Earth's surface) -44 units absorbedby the Earth's surface The Earth-atmosphere system emits thermal infrared radiation. The upward flux from the Earth's surface is -115 units. The cloud-free atmosphere emits -33 units back to the Earth's surface and -34 units out to space. The cloudy atmosphere emits -67 units back to Earth and -36 units out to space. Thus -70 units of infrared radiation leave the top of the atmosphere, balancing the net -70 units of solar radiation penetrating the top of the atmosphere.The net upward flux of infrared radiation at the surface of the Earth is -15 units, consisting of -115 units emitted by the Earth and -100 units radiated back to Earth by the cloud-free and cloudy atmosphere. The incoming solar energy absorbed by the Earth is -44 units; this is balanced by the net upward flux of infrared radiation of -15 units, plus -6 unit loss by sensible heat conduction, and - 23 unit loss by latent heat. The Earth emits -115 units of infrared radiation to the atmosphere, whereas the atmosphere emits -170 units of infrared radiation, a net deficit of -55 units. Since the atmosphere absorbs -26 units of solar radiation, the net radiative loss from the atmosphere is -29 units; this is made up for by the sensible and latent heat fluxes. The net radiative cooling of the atmosphere is thus balanced by the latent heat of condensation released in precipitation processes and by the convection and conduction of sensible heat from the surface. The average annual ratio of sensible to latent heat loss at the surface is called the Bowen ratio. With -6 units of sensible heat loss and -23 units of latent heat loss, the Bowen ratio is -0.27. Figure 1.16 shows the zonally annual averaged absorbed solar and emitted infrared fluxes, as observed from satellites. We note a net gain of radiative energy between about 40° Nand 40° S, and a net loss of energy in the polar regions. This pattern results largely from the decrease in insolation to the polar regions in winter and from the high surface albedo in the polar regions. The outgoing infrared flux displays only a small latitudinal dependence. The tendency for the outgoing infrared flux to be greatest in the tropics where the surface temperature is largest is muted by the larger amount of atmospheric water vapor and the higher and colder clouds in the tropics. As a result of the net gain of radiative energy in the tropics and the net loss in the polar regions, an equator-to-pole temperature gradient is generated.

'The average value of the albedo, the incoming radiation that is reflected or scattered back to space without absorption, is usually taken to be somewhere in the range of 30 to 34%. It is important to note that the albedo varies considerably, depending on the surface of the Earth. For example, in the polar regions, which are covered by ice and snow, the reflectivity of the surface is very high. On the other hand, in the equatorial regions, which are largely covered with oceans, the reflectivity is low, and most of the incoming energy is absorbed by the surface.

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dimensions. Four rough categories have proved convenient to classify atmospheric scales of motion:

1. Microscale. Phenomena occurringon scalesof the orderof 0 to 100m, suchasthe meanderinganddispersionof a chimneyplumeandthe complicatedflow regimein the wakeof a largebuilding. 2. Mesoscale. Phenomenaoccurring on scalesof tens to hundredsof kilometers, such as land-seabreezes,mountain-valley winds, and migratory high- and lowpressurefronts. 3. SynopticScale. Motions of whole weathersystems,on scalesof hundredsto thousandsof kilometers. 4. Global Scale. Phenomenaoccurringon scalesexceeding5 X 103kin. Spatial scales characteristic of various atmospheric chemical phenomena are given in Table 1.3. Many of the phenomena in Table 1.3 overlap; for example, there is more or less of a continuum between (1) urban and regional air pollution, (2) the aerosol haze associated with regional air pollution and aerosoVclimate interactions, (3) greenhouse gas increases and stratospheric ozone depletion, and (4) tropospheric oxidative capacity and stratospheric ozone depletion. The lifetime of a species is the averal!:etime that a molecule

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atmosphere's scales of motion. Much of this book will be devoted to understanding the exquisite interactions between chemical and transport processes in the atmosphere. In anticipation of much of the remainder of this book, Table 1.4 summarizes atmospheric effects of trace gases.

APPENDIXl

DERIVATION OF THE GEOSTROPffiC WIND SPEED

The direction of winds in the geostrophic layer is determined by horizontal pressure gradients and Coriolis forces. As we have discussed, an air parcel moving southward in the Northern Hemisphere as a result of pressure gradients is accelerated toward the west by the Coriolis force. We can actually compute the wind speed and direction at any latitude as a function of the prevailing pressure gradient if we assume that only pressure and Coriolis forces influence the flow. It can be shown that the acceleration experienced by an object on the surface of the Earth (or in the atmosphere) moving with a velocity vector u consists of two components, -0 x (0 x r) and - 2(0 x D), where 0 is the angular rotation vector for the Earth and r is the radius vector from the center of the Earth to the point in question. The first term is simply the centrifugal force, in a direction that acts normal to the Earth's surface and is counterbalanced by gravity. The second term, 0 x u, is the Coriolis force. This force

arisesonly when an object, suchas an air parcel,is moving; that is U :;t: O. Even though the Coriolis force is of much smaller magnitude than the centrifugal force, only the Coriolis force has a horizontal component. Since the winds are horizontal in the geostrophic layer, the Coriolis acceleration is given by the horizontal component of the Coriolis term, namely, 2uGQ sin cf>,where Q is the rate of rotation of the Earth and cf> is the latitude.The direction of the Coriolis force is perpendicular to the wind velocity, as shown in Figure I.A.I. Wind speed UG at latitude cf> lies in the horizontal plane.

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Low p

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Weusuallydenote20 sin r/> by f, calledthe Coriolis parameter.From the continuity equation (I.A.I), we seethat au/ax = 0, since v = w = O.Thusfrom (I.A.5) op/ox = 0, and the direction of flow is perpendicular to the pressure gradient op/oy. In addition, from

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is exactly balancedby the

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The approach to the geostrophic equilibrium for an air parcel starting from rest, accelerated by the pressure gradient and then affected by the Coriolis force, is shown in Figure I.A.2. The geostrophic balance determines the wind direction at altitudes above about 500 m. In order to describe the air motions at lower levels we must take into account the friction of the Earth's surface. The presence of the surface induces a shear in the wind profile, as in a turbulent boundary layer over a flat plate generated in a laboratory wind tunnel. In analyzing the geostrophic wind speed we found that for steady flow a balance exists between the pressure force and the Coriolis force. Consequently, steady flow of air at levels near the ground leads to a balance of three forces: pressure force, Coriolis force, and friction force due to the Earth's surface. Thus, as shown in Figure 1.A.3, the net result of these three

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REFERENCES Bentley,C. R. (1997)Rapidsea-levelrise soonfrom WestAntarctic Ice Sheetcollapse?Science,275, 1077-1078. Boering,K. A. et al. (1995) Measurements of stratosphericcarbondioxide andwater vapor at northern midlatitudes:implications for troposphere-to-stratosphere transport,Geophys.Res.Lett., 22, 2737-2740. Brasseur,G., and Solomon,S. (1984)Aeronomyof the Middle Atmosphere.Reidel, Dordrecht,The Netherlands. Cloud,P.(1983).The biosphere.Sci.Am., 249, 176-189. Graedel,T. E., and Crutzen,P.J. (1989)The changingatmosphere,Sci.Am., 261,58-68. Hickey,J. R., Alton, B. M., Kyle, H. L., andHoyt, D. (1988)Total solarirradiancemeasurements by ERB/Nimbus-7:a review of nine years,SpaceSci. Rev.,48,321-342. Holton, J. R., Haynes,P. H., McIntyre, M. E., Douglass,A. R., Rood, R. B., and Pfister,L. (1995) Stratosphere-troposphere exchange,Rev.Geophys.,33,403-439. IntergovernmentalPanel on Climate Change (IPCC) (1995) Climate Change 1994: Radiative Forcing of ClimateChangeand an Evaluationof the 1PCCIS92 EmissionScenarios.Cambridge University Press,Cambridge,UK. Iqb~l, M. (1983)An Introduction to Solar Radiation.AcademicPress,Toronto. Liou, K. N. (1992) Radiation and Cloud Processesin the Atmosphere.Oxford University Press, Oxford, UK. Mecherikunnel,A. T., Lee, R. B., Kyle, H. L., andMajor, E. R. (1988)Intercomparisonof solartotal irradiancedatafrom recentspacecraftmeasurements, J. Geophys.Res.,93,9503-9509. Prather,M., McElroy, M., Wofsy, S., Russell,G., and Rind, D. (1987) Chemistryof the global troposphere:fluorocarbonsastracersof air motion, J. Geophys.Res.,92,6579-6613. Ramanathan, V. (1987)The role of earthradiationbudgetstudiesin climate and generalcirculation research,J. Geophys.Res.,92,4075-4095. Ramanathan, V., Cess,R. D., Harrison,E. F., Minnis, P.,Barkstrom,B. R.,Ahmad,E., andHartmann, D. (1989)Cloud-radiativeforcing andclimate:resultsfrom the earthradiationbudgetexperiment, Science,243,57-63. Salstein,D. A. (1995) Mean propertiesof the atmosphere,in Composition,Chemistry,and Climate of theAtmosphere,editedby H. B. Singh.VanNostrandReinhold,New York, pp. 19-49. Trenberth,K. E., andGuillemot, C. J. (1994)The total massof the atmosphere,J. Geophys.Res.,99, 23079-23088. Walker,J. C. G. (1977)Evolution of theAtmosphere.Macmillan, New York. Willson, R. C., and Hudson,H. S. (1988) Solar luminosity variationsin solarcycle 21, Nature,332, 810-812.

PROBLEMS l.lA Calculatethe concentration(in moleculescm-3) andthe mixing ratio (in ppm) of water vapor at groundlevel at T = 298 K at RH valuesof 50%, 60%, 70%, 80%, 90%, 95%, and 99%. The vapor pressureof pure wateras a function of temperaturecanbe calculated with the following correlation: PH10(T)= Ps exp[13.3185a- .976a2- O.6445a3- O.1299a4]

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2 Virtually every element in the periodic table is found in the atmosphere; however, when classifying atmospheric species according to chemical composition it proves to be convenient to use a small number of major groupings such as: 1. Sulfur-containing compounds. 2. Nitrogen-containing compounds. 3. Carbon-containing compounds. 4. Halogen-containing compounds. Obviously these categories are not exclusive; many sulfur-containing compounds, for example, also include atoms of carbon. And virtually all the atmospheric halogens involve a carbon atom backbone. We do not include in the above list species of the general formula HxOy; with the exception of water and hydrogen peroxide (H2O2), these are all radical species (e.g., hydroxyl, OH) that play key roles in atmospheric chemistry but do not necessarily require a separate category. Every substance emitted into the atmosphere is eventually removed so that a cycle of the elements in that substanceis established. This is called the biogeochemical cycle of the element. The biogeochemical cycle of an element or a compound refers to the transport of that substance among atmospheric, oceanic, biospheric, and land compartments, the amounts contained in the different reservoirs, and the rate of exchange among them. The circulation of water among oceans, atmosphere, and continents is a prime example of a biogeochemical cycle. The term biogeochemical cycle is often used to describe the global or regional cycles of the "life elements," C, 0, N, S, and P, with reservoirs including the atmosphere, the ocean, the sediments, and living organisms (Rodhe, 1992). A condition of "air pollution" may be defined as a situation in which substancesthat result from anthropogenic activities are present at concentrations sufficiently high above their normal ambient levels to produce a measurable effect on humans, animals, vegetation, or materials. This definition could include any substance, whether noxious or benign; however, the implication is that the effects are undesirable. Traditionally, air pollution has been viewed as a phenomenon characteristic only of large urban centers and industrialized regions. It is now clear that dense urban centers are just "hotspots" in a continuum of trace speciesconcentrations over the entire Earth. Both urban smogs and stratospheric ozone depletion by chlorofluorocarbons are manifestations of what might be termed in the broadest senseas air pollution. The first recognized type of air pollution was that typified by high concentrations of sulfur compounds (S02 and sulfates) and particles, resulting from combustion of coal and high-sulfur-containing fuels. Cities with this characteristic type of air pollution are often in cold climates where electric power generation and domestic heating are major sources of emissions. 49

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from direct emissionsof particlesbut also from emissionsof certaingasesthat either condenseas particlesdirectly or undergochemicaltransformationto a speciesthat condenses as a particle.A full descriptionof atmosphericparticlesrequiresspecificationof not only their concentrationbut also their size,chemicalcomposition,phase(i.e., liquid or solid), andmotphology. Onceparticlesare in the atmo~phere, their size,number,and chemicalcompositionare changedby severalmechanismsuntil ultimately they are removedby natural processes. Someof the physicaland chemicalprocessesthat affect the "aging" of atmosphericparticles are more effective in one regimeof particle size than another.In spite of the specific processesthat affect particulateaging,the usualresidencetime of particlesin the lower atmospheredoesnot exceedseveralweeks.Very closeto the ground,the main mechanisms for particle removalaresettling anddry depositionon surfaces;whereasat altitudesabove about 100m, precipitationscavengingis the predominantremovalmechanism. As air risesthrougha cloud andbecomesslightly supersaturated with water vapor (i.e., asits relative humidity exceeds100%),cloud dropletsform on condensationnuclei-usually solubleaerosolparticles(e.g., microscopicparticlesof various salts)that exist in the atmosphereat concentrationsof 100 to .1000cm-3-and grow by condensationof water vapor.As the droplets grow and collide with each other they becomeraindrops,which grow rapidly asthey fall and accretecloud droplets. 2.1.1

Residence Time

The fundamental physical principle governing the behavior of a chemical in the atmosphere is conservation of mass. In any imaginary cube of air the following balance must hold:

Rateof the species flowing in

Rateof the Rateof introduction (emisspecies + flowing out sion) of the species

-

Rateof removal of = the species

Rateof accumulation of the speciesin the imaginaryvolume

This balance must hold from the smallest cube of air all the way up to the entire atmosphere. If we let Q denotethe total massof the substancein the volume of air, Fin and Foulthe massflow ratesof the substancein and out of the air volume,respectively,P the rateof introductionof the speciesfrom sources,andR the rate of removalof the species,thenconservationof masscanbe expressedmathematicallyas dQ dt

(Fin

-

Fout)+ (P - R)

(2.1)

If the amountQ of the substancein the volume or reservoiris not changingwith time, then Q is a constant and dQ/dt

= O. In

order for Q to be unchanging,

all the sources of the

substanceto the reservoirmust be preciselybalancedby the sinks of the substance.This meansthat Fin

+ P = Foul+ R

In such a case steadv-state conditions are said to hold

(2.2)

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ATMOSPHERIC RESIDENCE llMES

53

The stratosphere can be considered well mixed vertically only for atmospheric species with residence times greatly exceeding 50 years. In fact, one of the only examples of such a long-lived species is He, which has its source at the Earth's surface and its sink as escape through the very top of the atmosphere into space. Thus the stratosphere is poorly mixed vertically for essentially all atmospheric trace constituents. Frequently the rate at which a chemical is removed from the atmosphere is proportional to its concentration (first-order loss)-the more that is present, the faster its rate of removal. This is generally the case for both dry deposition at the Earth's surface and scavenging by cloud droplets. Consider a species for which steady-state conditions hold and which is removed at a rate proportional to its concentration with a proportionality constant A. Such a species is 85Kr,the only significant removal process for which is radioactive decay. For 85Kr,then

Q -)..Q

1 -)..

Thus to estimate the residence time of 85Kr does not even require knowledge of its atmospheric abundance Q, but only of the radioactive decay constant. Consequently, it does not matter whether 85Kris uniformly mixed throughout the entire atmosphere or not to estimate its lifetime. In a case where the removal process is first order, then even for a poorly mixed species a simple and accurate estimate for its residence time can be obtained provided that its removal rate constant can be accurately estimated. Now consider a species with mass Q in the atmosphere that is removed by two independent processes,the first at a rate kl Q and the second at a rate k2 Q, where kl and k2 are the first-order removal coefficients. Its overall residence time is given by

or ~ = kl + k2 1:

Process1. for example.could be dry depositionand process2 cloud scavenging.We can actually associatetime constantswith the two individual removalprocesses,

.. = kl

T2= k2

where ., canbe thoughtof asthe residencetime of the speciesif the only removalprocess is process1, which is alsotrue for .2. From (2.8) and (2.9) we canexpressthe overall residencetime. in termsof the two individual removaltimes ., and .2 by

..:.=..:.+..:. . .1 .2

(2.10)

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SULFUR-CONTAINING

COMPOUNDS

55

The loss processesare usually representedas first order; for example, Rt = kt Qi' wherethe first-orderrate constants,which we will denoteby k's, must be specified.Thus (2.13)becomes dQi dt

= p.n , + p.Q I + p.C , - (k~ I + k!lJ , + k'! , + k~ , )Q,'

(2.14)

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p.n+ p.a+ p.C,

,

,

( kd

+ ,

k!1'

+ ,

k':

+ ,

k~

)

Q,

'

=

0

I

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-

.,.,=

kd, + k!V , + k,:I + k~I

or

t";=

Qi

p!'I + p.a I + p.C I

(2.17)

To use (2.16) the individual first-order rate constants for removal must be estimated, whereas in (2.17), estimates for the total number of moles in the troposphere, which can be derived from a concentration measurement, and for the source strength terms are needed. If the k; values are difficult to specify, mean residence times are often estimated from (2.17).

2.2 SULFUR.CONTAINING COMPOUNDS Sulfur is present in the Earth's crust at a mixing ratio of less than 500 parts per million by mass and in the Earth's atmosphere at a total volume mixing ratio of less than I ppm. Yet, sulfur-containing compounds exert a profound influence on the chemistry of the atmosphere and, likely, on climate. The main questions that we will seek to answer with respect to sulfur compounds, and indeed for all classes of atmospheric compounds, are: 1. What are the species present in the atmosphere? What are their natural and anthropogenic sources? 2. What chemical reactions do they undergo in the atmosphere? How fast are these reactions? 3. What are the products of atmospheric transformations? 4. What effect does the presence of the compound and its chemical transformation products have on the atmosphere? In this chapter we focus on the first question.

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TABLE 2.3 Observed Mixing Ratios of Atmospheric Sulfur Gases CompoundandLocation H2S Marine surfacelayer Coastalregions Forests Wetlands Urban areas Freetroposphere(2-5 kin) CH3SCH3 Marine surfacelayer Continentalsurfacelayer Freetroposphere(2-5 kin) CS2 Marine surfacelayer Continentalsurfacelayer Freetroposphere(2-5 kin) OCS Total troposphere Marine surfacelayer Continentalsurfacelayer SO2 Marine surfacelayer Freetroposphere(>5 km)Europe/NorthSea/Arctic North America cleancontinental CoastalEurope Pollutedcontinentalair

Average Mixing Ratio (ppt)

3.6-7.5 65 35-60 450-840 365 6-8.5 80-110 8-60 1.5-15 2-18 35-120 5-7 500 500 545 20 50 160 260 1500

Source: Berresheim et al. (1995) (detailed references given by the authors).

2.2.1

Dimethyl Sulfide (CH~CH3)

Dimethyl sulfide (DMS) is the dominant sulfur compound emitted from the world's oceans. DMS was discovered in the surface ocean by Lovelock et al. (1972), who suggested that DMS may be the biogenic sulfur species that was needed at the time to balance the global sulfur budget. DMS is produced in oceanic waters by both benthic and planktonic marine organisms (Dacey and Wakeham, 1986), suggesting that it may be ubiquitous in the surface ocean (Barnard et al., 1982). It is thought to originate from the decomposition of dimethyl-sulfoniopropionate produced by marine organisms, in particular, phytoplankton (Andreae, 1990). Its concentration in the upper layer of the ocean varies between a few nanograms of S per liter to a few micrograms of S per liter (Lovelock et al., 1972; Barnard et al., 1982; Andreae and Raemdonck, 1983; Cline and Bates, 1983; Nguyen et al., 1984, 1988). The DMS surface seawater concentration is highly nonuniform; its average concentration is approximately 100 nanograms (ng) of S per liter. It has been observed that the concentration of DMS is dependent on diurnal (Andreae and Barnard, 1984) and seasonal variations (Turner and Liss, 1985), and on depth and location (Andreae and Raemdonck, 1983).

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and 0.57 Tg in the stratosphere.Basedon the estimatedglobal OCS sourcestrengthof 0.73Tg yr-I, the global atmosphericlifetime of OCS is estimatedto be about7 years.We will return to the global cycle and chemistryof OCS in Chapter4 in connectionwith the stratosphericaerosollayer. 2.2.3 Sulfur Dioxide (S02) Sulfur dioxide is the predominant anthropogenic sulfur-containing air pollutant. Mixing ratios of S02 in continental background air range from 20 ppt to over I ppb; in the unpolluted marine boundary layer levels range between 20 and 50 ppt. Urban S02 mixing ratios can attain values of several hundred parts per billion. We will consider the atmospheric chemistry of S02 in Chapter 5.

2.2.4 The Atmospheric Sulfur Cycle Figure2.2 depicts the major reservoirsin the biogeochemicalcycle of sulfur. with estimatedquantities(in Tg(S» in eachreservoir.Directions of fluxes betweenthe reservoirs areindicatedby arrows.The major pathwaysof sulfur compoundsin the atmosphereare depictedin Figure 2.3. The numberson eacharrow refer to the descriptionof the process givenin the captionto the figure (not to fluxes).All the processesindicatedschematically in Figure2.3 will be studiedlater in this book.

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The meanresidencetime of SO"is 7:S02

= P;02QS02 + PS02 a k~o, + k~O2+ kso,

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(2.23)

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2.3 NITROGEN-CONTAINING COMPOUNDS Aside from N2, which is extremely stable chemically and is not involved in the chemistry of the troposphere or stratosphere, the important nitrogen-containing trace species in the atmosphere are nitrous oxide (N2O), nitric oxide (NO), nitrogen dioxide (NOJ, nitric acid (HNO3), and ammonia (NH3). The first of these, nitrous oxide (N2O), is a colorless gas that is emitted almost totally by natural sources, principally by bacterial action in the soil. The gas is employed as an anesthetic and is commonly referred to as "laughing gas." The second, nitric oxide (NO), is emitted by both natural and anthropogenic sources. Nitrogen dioxide (NOJ is emitted in small quantities from combustion processesalong with NO and is also formed in the atmosphere by oxidation of NO. The sum of NO and NO2 is usually designated as NOx' Nitric oxide is the major oxide of nitrogen formed during high-temperature combustion, resulting from both the interaction of nitrogen in the fuel with oxygen present in the air and the chemical conversion of atmospheric nitrogen and oxygen at the high temperatures of combustion. Other oxides of nitrogen, such as NO3 and N20S, exist in the atmosphere in relatively low concentrations but nonetheless participate importantly in atmospheric chemistry. Nitric acid is an oxidation product of NO2 in the atmosphere. Ammonia (NH3) is emitted primarily by natural sources. Finally, nitrate and ammonium salts are not emitted in any significant quantities but result from the atmospheric conversion of NO, NO2, and NH3' Nitrogen is an essential nutrient for all living organisms. The primary source of this nitrogen is the atmosphere. However, N2 is not useful to most organisms until it is "fixed" or converted to a form that can be chemically utilized by the organisms. (Nitrogen fixation refers to the chemical conversion ofN2 to any other nitrogen compound.) The "natural" fixation of N2 occurs by two types of processes.One is the action of a comparatively few microorganisms that are capable of converting N2 to ammonia, ammonium ion (NHt), and organic nitrogen compounds. The other natural nitrogen fixation process occurs in the atmosphere by the action of ionizing phenomena, such as cosmic radiation or lightning, on N2. This process leads to the formation of nitrogen oxides in the atmosphere, which are ultimately deposited on the Earth's surface as biologically useful nitrates. In addition to natural nitrogen fixation, human activities have led to biological and industrial fixation and fixation by combustion. Humans have increased the cultivation of legumes, which have a symbiotic relationship with certain microorganisms capable of nitrogen fixation. Legumes provide an increase in the soil nitrogen and serve as a valuable food crop. Industrial nitrogen fixation consists primarily of the production of ammonia for fertilizer use. Combustion can also lead to the fixation of nitrogen as NOx' In the process of nitrification, ammonium is oxidized to N°"i and NO3 by microbial action. N2O and NO are by-products of nitrification; the result is the release of N2O and NO to the atmosphere. Reduction of NO3 to N2, NO2, N20, or NO is called denitrification. Denitrification is accomplished by a number of bacteria and is the process that continually replenishes the atmosphere's N2. Figure 2.4 depicts the atmospheric nitrogen cycle. 2.3.1

Nitrous Oxide (N2O)

Nitrous oxide (N2O) is an important atmospheric gas that is emitted predominantly by biological sources in soils and water. Although by comparison to CO2 and H2O, N2O has a far lower concentration, it is an extremely influential greenhousegas. This is a result of its long

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TABLE 2.4 Estimated Sourcesand Sinks of N2OTypical of the Last Decade Sources

Range(Tg(N) yr-l)

Likely (Tg(N) yi

NATURAL

Oceans Tropical soils Wet forests Dry savannas Temperate soils Forests Grasslands

1-5

3

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3 1

0.1-2.0 0.5-2.0 9

Total natural sources ANTHROPOGENIC

Cultivated soils Biomassburning Industrial sources Cattle and feedlots

1.8-5.3 0.2-1.0 0.7-1.8 0.2-0.5

3.5 0.5 1.3 0.4

Implied total sources (atmospheric increase + total sinks)a "The observed atmospheric increase implies that sources exceed sinks by 3.9 Tg(N) yr-i. Source: IPCC (1995).

the stratosphere.We will return to this processin Chapter4. Sourcesof N2Oexceedestimatedsinksby 2.4 Tg(N) yr-i. Estimatesfor the atmosphericlifetime of N2Ocomefrom stratosphericchemicaltransport modelsthat havebeentestedagainstobservedN2Odistributions.The bestcurrentestimate for the lifetime ofN2O is 120 :t 30 years.Becauseof its long lifetime N2Oexhibits more or lessuniform concentrationsthroughoutthe troposphere.Ice core recordsof N2O show a preindustrialmixing ratio of about276 ppb. N2O levels haverisen approximately 15% since preindustrialtimes, reaching 311 ppb in 1992 (IPCC, 1995; Machida et al., 1995)(Figure 2.5). This observedatmosphericincreaseis consistentwith a differenceof 3.9 Tg(N) yr-1 excessof sourcesover sinks, which is in reasonableagreement,given the uncertainties,with the mismatchbasedon attemptingto estimatesourcesand sinks indeoendentlv.

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Detailedemissionsinventoriesareavailablefor Canada,the United States,and western Europethat describethe spatialpatternsof NOxemissionsfrom combustionof fossil fuels andfrom industrialprocesses(Lubkert andZierock, 1989;Placetet al., 1990).Between40 and 45% of all NOxemissionsin the United Statesareestimatedto comefrom transportation, 30 to 35% from power plants,and about20% from industrial sources.About half the NOx emissionsassociatedwith transportationcome from light-duty gasolinetrucks and carsand approximatelyone-quarterare from heavy-dutygasolineanddieselvehicles.3 2.3.3

Reactive Odd Nitrogen (NO,)

Reactive nitrogen, denoted NO" is defined as the sum of the two oxides of nitrogen (NOx = NO + NOJ and all compounds that are products of the atmospheric oxidation of NOx' These include nitric acid (HNO3), nitrous acid (HONO), the nitrate radical (NO3), dinitrogen pentoxide (N2Os), peroxynitric acid (HNO4), peroxyacetyl nitrate (PAN) (RC(O)OONOJ and its homologues, alkyl nitrates (RONOJ, and peroxyalkyl nitrates (ROONOJ. Nitric acid (HNO3) is the major oxidation product of NOx in the atmosphere. Because of its extreme water solubility, HNO3 is rapidly deposited on surfaces and in water droplets. Also, in the presence of NH3, HNO3 can form an ammonium nitrate (NH4NO3) aerosol. The nitrate radical (NO3) is an important constituent in the chemistry of the troposphere, especially at night. NO3 is present at night at mixing ratios ranging up to 300 ppt in the boundary layer. Nitrous oxide (N2O) and ammonia (NH3) are not considered in this context as reactive nitrogen compounds. Measurement of total NO, in the atmosphere provides an important measure of the total oxidized nitrogen content. Concentrations of individual NO, species relative to the total indicate the extent of interconversion among species. NO, is indeed closer to a conserved quantity than any of its constituent species (Roberts, 1995). There is a sizable body of data on the concentrations of NOx in the atmosphere, but caution must be exercised in drawing conclusions from these measurements. Many measurements of NO x have been made by devices that convert NO2 to NO, which is then measured by the phenomenon of chemiluminescence. Comparison of these measurements with more specific techniques suggests that surface converters that can convert NO2 to NO also convert other reactive nitrogen oxide species, such as peroxyacetyl nitrate (PAN), to NO, thereby causing interference. In urban locations, where the local NO sources are typically large, NO and NO2 are probably the dominant constituents of the total reactive nitrogen NO,. Thus, in urban areas, interference from PAN and other oxides of nitrogen is believed to be relatively small. In rural and remote locations, however, the interference can be substantial. For this reason, all nonurban NOx measurements made with surface converters must be considered upper limits (biased toward a high measurement). Given the dominant role of anthropogenic emissions in the budget of atmospheric NOx and the fact that the sources of these emissions tend to be located in or near urban areas,elevated concentrations of NOx are to be expected in these locations. Observations of NOx support this expectation. The range and variability of NOx measurements are reflected in measurements made in 29 cities across the eastern and southern United States during the 3It is estimated that in 1994 there were 147,000,000 light-duty motor vehicles in the United States, 48,000,000 trucks (85% light-duty), and 676,000 buses. Total vehicle miles traveled were estimated as 235 X 10'2, with 1.4 X 10'igallons of gasoline consumed.

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TABLE 2.6 Typical Boundary Layer NOx Mixing Ratios

Urban-suburban Rural Remotetropical forest Remotemarine Source: National Research Council (1991).

rural sites and generally range from a few tenths to 1 ppb. Measurements of NOx in the atmospheric boundary layer and lower free troposphere in remote maritime locations have generally yielded mixing ratios of 0.02 to 0.04 ppb (20 to 40 ppt). Although the databaseis still quite sparse, mixing ratios in remote tropical forests (not under the direct influence of biomass burning) appear to range from 0.02 to 0.08 ppb (20 to 80 ppt); the somewhat higher NOx concentrations found in remote tropical forests, as compared with those observed in remote marine locations, could result from biogenic NOx emissions from soil. A summary of the NOx measurements made in the four regions of the globe mentioned above is presented in Table 2.6. It can be seen that NOx concentrations decreasesharply as one moves from urban and suburban to rural sites in the United States and then to remote sites over the ocean and tropical forests. The striking difference of three orders of magnitude or more between NOx concentrations in urban-suburban areas and remote locations is compelling evidence for the dominant role of anthropogenic emissions of NOx over strong anthropogenic source regions such as North America. Because the ability to measure Nay was developed only recently, the rural and remote Nay databaseis even more limited than that for NOx' However, there are enough data to establish a rough indication of the Nay distribution. Average Nay concentrations observed at many sites in the United States are quite similar; median mixing ratio values range from 3 to 10 ppb. These are somewhat lower than NOx mixing ratios typically observed in urban and suburban locations, which range from 10 to 1000 ppb. The contrast in Nay concentrations found in rural areas of the continental United States with those observed in the remote troposphere is illustrated in Figure 2.7. The measurement sites are Scotia, Pennsylvania, a rural site in the eastern United States; Niwot Ridge, Colorado, an isolated inland site in the western United States; Point Arena, California, a site on the West Coast that often receives air from the Pacific Ocean; and Mauna Loa, Hawaii, a remote maritime site. Two of the sites, Mauna Loa and Niwot Ridge, are at high elevations (approximately 3 kin), and thus the air sampled there is not necessarily representative of the boundary layer. There is a progressive decreasein the contribution of NOx to Nay as one moves toward more remote regions. On average, NOx at Scotia accounted for 59% of the observed Nay. At Niwot Ridge in 1987, NOx accounted for 32% of the Nay, and at Mauna Loa, NOx accounted for only 15% of the Nay. Because Nay enters the atmosphere as NOx, the decrease in the ratio of NOx to Nay as one moves to more remote sites can be understood in terms of the increasing chemical conversion of NOx to organic nitrates (principally PAN) and to inorganic nitrates (principally HNO3) with increasing distance of the site from major anthropogenic sources.The most remote sites are characterized by the lowest ratios of NOxINOv' Those sites at hi.e:haltitudes have the lar.e:estratio of

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TABLE 2.7 Estimated Global Ammonia Emissions Emission

Sourceof Ammonia

(Tg(N) yr-1

,NTHROPOGENIC

Dairy cattle Beef cattle/buffalo Pigs Horses Sheep/goats Poultry Fertilizer Biomassburning Subtotal NATURAL Wild animals Vegetation Ocean Subtotal Total

5.5 8.7 2.8 1.2 2.5 1.3 6.4 2.0 3OA 2.5 5.1 -1.:Q 1.4,.Q 45.0

Source: Dentener and Crutzen (1994).

are animal waste, ammonification of humus followed by emission from soils, losses of NH3-based fertilizers from soils, and industrial emissions (Table 2.7). The ammonium (NHt) ion is an important component of the continental tropospheric aerosol. Because NH3 is readily absorbed by surfaces such as water and soil, its residence time in the lower atmosphere is expected to be quite short, about 10 days. Wet and dry deposition of NH3 are the main atmospheric removal mechanisms for NH3' In fact, deposition of atmospheric NH3 and NH: may represent an important nutrient to the biosphere in some areas. Atmospheric concentrations ofNH3 are quite variable, depending on proximity to a sourcerich region. NH3 mixing ratios over continents range typically between 0.1 and 10 ppb.

2.4 CARBON-CONTAINING COMPOUNDS 2.4.1

Classification of Hydrocarbons

Let us review briefly the classifications of carbon-containing compounds, particularly those of interest in atmospheric chemistry. The carbon atom has four valence electrons and can therefore share bonds with from one to four other atoms. The nature of the carbon-carbon bonding in a hydrocarbon molecule basically governs the properties (as well as the nomenclature) of the molecule.

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Moleculeswith two doublebondsarecalled alkadienes,an exampleof which is CH2=CH-CH=CH2 1,3-Butadiene

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Individual Estimate

Total

NATURAL

Wetlands Termites Oceans Other Total identified natural sources

115(55-150) 20 (10-50) 10 (5-50) 15 (10-40) 160 (110-210)D

ANTHROPOGENIC

Fossil-fuelrelatedsources Natural gas Coal mines Petroleumindustry Coal combustion Total fossil-fuel related Biosphericcarbon Enteric fermentation Rice paddies Biomassburning Landfills Animal waste Domesticsewage Total biospheric

40 (25-50) 30 (15-45) 15 (5-30) ?(1-30) 100(70-120)b 85 (65-100) 60 (20-100) 40 (20-80) 40 (20-70) 25 (20-30) 25 (15-80) 275 (200-350)

~

375 (300-450)"

Total identified anthropogenic sources Total identified sources

Sinks TroposphericOH Stratosphere Soils Total sinks

445 (360-530) 40 (32-48) 30 (15-45) 515 (430-600)

Total global burden:4850Tg (C~) aA preindustrial level of 700 ppb would have required a source of2l0 Tg(CH.) yr-1 if the lifetime has remained constant, and 280 Tg (CH.) yr-1 if current tropospheric chemical feedbacks can be extrapolated back. The total anthropogenic emissions of CH. based on identified sources, 375 (300--450), is slightly higher than the inferred range from preindustrial levels, 270-340, but is well within the uncertainties. bpractional source from fossil carbon based on a measure of the atmospheric ratio of "CH. to 12CH.. Note: The observed increases in methane show that sources exceed sinks by about 35 to 40 Tg each year. All data are rounded to the nearest 5 Tg. Source: IPCC (1995).

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TABLE 2.10 Estimated Global Anthropogenic Emissionsof NonmethaneVolatile Organic Compounds

FUEL PRODUCTION/DISTRIBUTION Petroleum

Natural gas Oil refining Gasoline distribution

8 2 5 2.5

FUEL CONSUMPTION

Coal Wood Crop residues(including waste) Charcoal Dung cakes Roadtransport Chemicalindustry Solventuse Uncontrolledwasteburning

3.5 25 14.5 2.5 3 36 2 20 8

OTHER

10

Total

142

Source: Middleton (1995).

pogenic VOC emissions included, for example, about 600 different compounds (Placet et al., 1990). Motor vehicles are the dominant contributor to VOC emissions in the United States. VOCs emitted from motor vehicles are mainly hydrocarbons that result from the incomplete combustion of fuel or from its vaporization. These contributions are generally categorized and reported as exhaust and evaporative emissions. Within the exhaust emissions category are included the unburned and partially burned fuel and lubricating oil in the exhaust and gases that leak from the engine. The evaporative emissions category includes fuel vapor emitted from the engine and fuel system that can be attributed to several sources: vaporization of fuel as a result of the heating of the fuel tank, vaporization of fuel from the heat of the engine after it has been turned off (hot-soak emissions), vaporization of fuel from the (uel system while the vehicle is operating (running losses), fuel losses due to leaks and diffusion through containment materials (resting losses), and fuel vapor displacement as a result of filling fuel tanks (refueling losses). Motor vehicles are the major sources of alkane and aromatic emissions. Estimates of global anthropogenic nonmethane VOC emissions in 1990 are given in Table 2.10. In general, a breakdown by chemical compound is not yet available for global anthropogenic VOC emissions. As seen in Table 2.10, transportation is the largest source ofVOC emissions worldwide, with solvent use following as the second largest source. As an illustration of the large number of organic compounds identified in the atmosphere, Table 2.11 lists the median concentrations of the 25 most abundant nonmethane organic species measured in the 1987 Southern California Air Quality Study.

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Isoprene

=~>==

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~ IlJ

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3

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TABLE 2.13 Global Biogenic VOC Emission Rate Estimates by Sourceand Class of Compound (Tg yr-1

821 120 194 5 9

127

260

.Other reactive biogenic VOCs (ORVOC). bThese totals include additional non-reactive VOCs not reflected in the columns to the left. Source: Guenther et al. (1995).

biogenic hydrocarbon emissions in the southeastern United States is predicted to be as large as that in the tropics. An estimate of global biogenic VOC emissions appears in Table 2.13. On a global basis, biogenic hydrocarbon emissions far exceed those of anthropogenic hydrocarbons. 2.4.5

Carbon Monoxide

The global sources and sinks of CO are given in Table 2.14. Methane oxidation (by OH) is a major source of CO, as are technological processes (combustion and industrial processes), biomass burning, and the oxidation of nonmethane hydrocarbons. Uncertainties in each of these estimated sources are large. It is estimated that about two-thirds

TABLE 2.14 Estimated Sourcesand Sinks of CO Typical of the Last Decade

Technological Biomassburning Biogenics Oceans Methaneoxidation NMHC oxidation

Sinks

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As a group of atmospheric chemicals, halogen-containing compounds have a wide variety of anthropogenic and natural sources. They are produced by biological processesin the oceans, from sea salt, from biomass burning, and from industrial synthesis. Their atmospheric lifetimes vary considerably depending on their mechanism of removal, ranging from a few days to several centuries. Table 2.15 lists atmospheric halogenated organic species with global average concentrations, atmospheric burdens, lifetimes, sources, and sinks.4 Of the exclusively man-made organic halogenated species, the chlorofluorocarbons are used as refrigerants (CFC-12, HCFC-22), blowing agents (CFC-ll, HCFC-22), and cleaning agents (CFC-113). Methyl chloroform (CH3CCI3), methylene chloride (CHzClz), and tetrachloroethene (CzCI4) are used as degreasersand as dry cleaning and industrial solvents. Methyl bromide (CH3Br) is a widely used agricultural and space fumigant. All the monomethyl halides in Table 2.15 have natural sources. Methyl chloride (CH3Cl) and CH3Br are also products of biomass burning. Lovelock (1971) first detected SF6 and CFCl3 in the atmosphere using the electron capture detector. In landmark work in atmospheric chemistry for which they received the 1995 Nobel Prize in Chemistry, Molina and Rowland (1974) showed that CFCs that are immune to removal in the troposphere could decompose photolytically in the stratosphere to release Cl atoms capable of catalytic destruction of stratospheric ozone. The very lack of chemical reactivity that makes chlorofluorocarbon molecules so intrinsically useful also allows them to survive unchanged in most commercial applications and eventually to be released to the atmosphere in their original gaseous form. The usual tropospheric sinks of oxidation, photodissociation, and wet and dry deposition are ineffective with the chlorofluorocarbons. The only important sink for CFC13and CFzClz is photodissociation in the midstratosphere (25 to 40 kill) by solar ultraviolet radiation with wavelengths shorter than 230 nm. These same CFCs that lead to stratospheric ozone depletion are efficient absorbers of infrared radiation and potentially important greenhouse gases. There is a sharp demarcation in atmospheric behavior between fully halogenated halocarbons and those containing one or more atoms of hydrogen. Halocarbons containing at least one hydrogen atom, such as CFzHCl, CHCI3, and CH3CCI3, are effectively broken down in the troposphere by reaction with the hydroxyl radical before they can reach the stratosphere. Atmospheric lifetimes of these species range from months to decades. Some of these gases also react with seawater; it is estimated that 5 to 10% of the removal of CH3CCl3 occurs by absorption into the ocean. The hydrohalogenated species such as

'The tenD hydrochlorofluorocarbons is the collective name given to a series of chemicals with varying number of carbon, hydrogen, chlorine, and fluorine atoms. The somewhat arcane system of numbering these compounds was proposed by the American Society of Heating and Refrigeration Engineers in 1957. For the simpler hydrochlorofluorocarbons, the numbering system may be summarized as follows: 1. The first digit on the right is the number of fluorine (F) atoms in the compound. 2. The second digit from the right is one more than the number of hydrogen (H) atoms in the compound. 3. The third digit from the right, plus one, is the number of carbon (C) atoms in the compound. When this digit is zero (i.e., only one carbon atom in the compound), it is omitted from the number. 4. The number of chlorine (CI) atoms in the compound is found by subtracting the sum of the fluorine and hydrogen atoms from the total number of atoms that can be connected to the carbon atoms.

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methyl chloroform (CH3CCI3) and the HCFCs are also significant infrared absorbers but their shorter atmospheric lifetimes reduce their radiative impact relative to the fully halogenated CFCs. Discovery of the stratospheric ozone-depleting potential of CFCs in the mid-1970s led to a ban in the United States (announced in 1976 and effective in 1978) on the use ofCFCs as aerosol propellants and to similar restrictions in Canada and Scandinavia. When the Antarctic ozone hole was discovered, subsequently, an international protocol outlining proposed actions to protect the stratospheric ozone layer was signed in Montreal in September 1987. The so-called Montreal Protocol specifies a 20% reduction from 1986 emissions of fully halogenated CFCs by 1994 and a further 30% reduction by 1999. Global emissions of CFCI3, CF2CI2, C2F3C13,CF2HC1, and CH3CC13up to 1992 are shown in Figure 2.9. The turn-down in emission rates reflects the results of the Montreal Protocol. Atmospheric levels of CF2CI2, CFCI3, and CH3CC13from 1978 to 1992 are shown in Figure 2.10. Mixing ratios increased steadily until about 1990. Methyl chloride (CH3Cl), at an atmospheric abundance of about 600 ppt, is the dominant halogen compound in the atmosphere. To maintain the steady-state CH3Cl concentration, with an atmospheric lifetime of order 2 years, requires a source strength of about 3.5 Tg yr-1 , most of which comes from the ocean. Southern Hemisphere (SH) CFC concentrations lag behind those in the Northern Hemisphere (NH) by about 1 year, reflecting the predominant source of CFCs in the NH and the approximate 1 year mixing time between the NH troposphere and the SH troposphere. At present, the atmosphere contains approximately 20 ppt of bromine, about half of which is methyl bromide (CH3Br). Methyl bromide is an ubiquitous component of the atmosphere, arising from both man-made and natural sources.A calculated atmospheric lifetime of 1.7 to 1.9 years, based solely on removal by reaction with OH radicals, is consistent with a global source of 90 to 110 Gg (109 g) yr-l (Singh and Kanakidou, 1993). A shorter lifetime of about 1.3 years (see Table 2.15) results if deposition/hydrolysis losses are also considered. Available data provide an estimate of global sources of CH3Br that divide 35% (20 to 50%) man-made and 65% (80 to 50%) natural. Oceans are supersaturated with CH3Br and constitute the major natural source of CH3Br of about 60 Gg yr-1, which could

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range from 40 to 80 Gg yr-l. CH3Br mixing ratios in the Northern Hemisphere ar,ound 1992 were about 14 ppt, with a ratio of NH to SH mixing ratios varying between 1.2 and 1.4 depending on season.The seasonalvariation is a result of source strength and OH variations. CH3Br levels have been increasing at a rate of 0.1 to 0.2 ppt yr-i. Despite the fact that CH3Br is largely removed in the troposphere by OH reaction, enough survives to enter the stratosphere and contribute to halogen-induced stratospheric ozone depletion.

2.6 ATMOSPHERIC OZONE Ozone(03) is a reactiveoxidant gasproducednaturally in traceamountsin the Earth's atmosphere.Ozonewas discoveredby C. F. Schonbeinin the middle of the last century;he alsowas first to detectozonein air (Schonbein,1840,1854).Schonbein(1840) suggested the presenceof an atmosphericgashaving a peculiarodor (the Greekword for "to smell" is ozein).Spectroscopicstudiesin the late 19th century showedthat ozoneis presentat a higher mixing ratio in the upper atmosphericlayersthan closeto the ground.Attemptsto explainthe chemicalbasisof existenceof ozonein the upperatmospherebegannearly 70 yearsago.Within the last 30 years,however,while increasedunderstandingof the role of other trace atmosphericspeciesin stratosphericozonewas unfolding, it becameapparent that anthropogenicallyemitted substanceshave the potential to seriouslydepletethe natural levels of ozonein the stratosphere. At aboutthe sameperiod, ironically, it was realized that anthropogenicemissionscould lead to ozone increasesin the troposphere. Whereasstratosphericozoneis essentialfor screeningof solar ultraviolet radiation,ozone at groundlevel can,at elevatedconcentrations,lead to respiratoryeffectsin humans.This paradoxicaldual role of ozonein the atmospherehas, on occasion,led to the dubbing of stratosphericozoneas "good" ozoneandtroposphericozoneas "bad" ozone. Most of the Earth's atmosphericozone(about90%) is found in the stratospherewhere it plays a critical role in absorbingultraviolet radiation emitted by the Sun. Figure 2.11

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from 1979to 1991have shownthat broadly similar midlatitudelosseshavealso occurred at equivalentlatitudesin the SouthernHemisphere(IPCC, 1995).Figure 2.12 showsthe annual,latitudinal variation of the total ozonetrendscalculatedfrom the Dobsongroundbasedmeasurementrecord and satellite observations.The effect of the Antarctic ozone hole is seenin the data from 60° S to 90° S. Analysis of the total ozonemappingspectrometer(TOMS) dataover the period 1979to 1989between69° S and69° N revealsa total global averageozonedecreaseof 3.5% over that 11year period (Hermanet al., 1991). Statisticaily significant decreasesin total ozoneare now being observedin all seasonsin both the Northernand SouthernHemispheresat middle andhigh latitudes(Stolarskiet al., 1992).

2.6.2

Ozone Flux from the Stratosphere to the Troposphere

Ozone from the stratosphere is transported down across the tropopause by so-called tropopause folding events (Danielson, 1968) in which tongues of stratospheric air intrude into the troposphere, usually at extratropicallatitudes (recall Figure 1.3 and associated discussion). Estimates of the amount of 03 transported from the stratosphere to the troposphere on an annual basis rely on measurements of conserved tracers or on general circulation models but are quite uncertain. In the NH ozone fluxes seem to maximize in spring, being as much as five times the value in the fall. The NH stratosphere-totroposphere ozone flux has been estimated to fall in the range of (3 to 8) X 1010molecules cm-2 S-l (Crutzen, 1995). The SH flux may be about half as large. An estimate for the total 03 production in the stratosphere is about 5 X 1013molecules cm-2 S-I; only about 0.1 % of all 03 produced in the stratosphere leaks down to the troposphere (Crutzen, 1995). The estimated global 03 loss from photolysis in the troposphere is 14 X 1010molecules cm-2 S-I, which generously exceeds the amount of 03 from stratosphere-to-troposphere exchange.

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downwind of large urban areas,under certain meteorological conditions, emissions of NO x and VOCs can result in ozone concentrations as high as 200 to 400 ppb. Such production of ozone and related oxidant species is called photochemical air pollution; it was first recognized in the Los Angeles basin in the 1940s. The database of ozone observations suggests a systematic pattern of decreasing daily maximum concentrations as one moves from urban-suburban locations to rural locations and then to remote locations. Daily maximum ozone concentrations within the atmospheric boundary laver tend to be lar2est in the

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2.7 PARTICULATE MATTER (AEROSOLS) Particles in the atmosphere arise from natural sources, such as windbome dust, sea spray, and volcanoes, and from anthropogenic activities, such as combustion of fuels. Whereas an aerosol is technically defined as a suspension of fine solid or liquid particles in a gas, common usage refers to the aerosol as the particulate component only (Table 2.17). Emitted directly as particles (primary aerosol) or formed in the atmosphere by gas-to-particle conversion processes (secondary aerosol), atmospheric aerosols are generally considered to be the particles that range in size from a few nanometers (nm) to tens of micrometers

TABLE 2.17 Terminology Relating to Atmospheric Particles Aerosols,aerocolloids, aerodispersesystems Dusts Fog

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Suspensions of solid particlesproducedby mechanicaldisintegrationof materialsuchascrushing,grinding, andblasting.Dp > I JLm. A looseterm appliedto visible aerosolsin which the dispersedphaseis liquid. Usually,a dispersionof water or ice, closeto the ground. The solid particlesgeneratedby condensationfrom the vapor state,generally after volatilization from meltedsubstances, and often accompanied by a chemicalreactionsuchas oxidation.Often the material involved is noxious.Dp < I JLm. An aerosolthat impedesvision and may consistof a combinationof water droplets,pollutants,and dust.Dp < 1 JLm. Liquid, usually water in the form of particlessuspendedin the atmosphereat or nearthe surfaceof the Earth; small water dropletsfloating or falling, approachingthe form of rain, and sometimesdistinguished from fog asbeing more transparentor as havingparticlesperceptibly moving downward.Dp > I JLm. An aerosolparticle may consistof a single continuousunit of solid or liquid containingmany moleculesheld togetherby intermolecular forcesandprimarily largerthan moleculardimensions(> 0.001 JLm). A particle may alsobe consideredto consistof two or more suchunit structuresheld togetherby interparticleadhesiveforcessuchthat it behavesas a singleunit in suspensionor upondeposit. A term derivedfrom smokeand fog, appliedto extensivecontamination by aerosols.Now sometimesusedloosely for any contaminationof the air. Small gas-borneparticlesresultingfrom incompletecombustion,consistingpredominantlyof carbonandother combustiblematerial,and presentin sufficientquantity to be observableindependentlyof the presenceof other solids.Dp 2: 0.01 JLm. Agglomerationsof particlesof carbonimpregnatedwith "tar," formed in the incompletecombustionof carbonaceous material.

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