3Aerosols in the Atmosphere

Andreas Petzold and Bernd Karcher

Abstract

This overview introduces atmospheric aerosol, including basic definitions andproperties. Types of aerosols, sources, transformation processes and sinks aredescribed. The spatial distribution of atmospheric aerosol is discussed on thebasis of field observations conducted at the DLR Institute of AtmosphericPhysics. The chapter concludes with a brief overview of aerosol impacts on theglobal climate.

3.1 Introduction

Airborne particles, or aerosols, form an intrinsic constituent of the global atmo-sphere. They affect global climate via their interaction with sunlight, influence theformation of clouds, and participate in many atmospheric chemical reactions.Figure 3.1 provides an overview on the various processes and interactions. Theseparticles can travel over distances of several thousands of kilometers and transportparticulate material such as dust, sea salt, spores, pollen and bacteria. Most of usthink of atmospheric aerosols in terms of fine dust, soot, air pollution, reduced

A. Petzold (&) � B. KärcherDLR, Institute of Atmospheric Physics (IPA), Münchner Straße 20,82234 Oberpfaffenhofen, Germanye-mail: a.petzold@fz-juelich.de

B. Kärchere-mail: Bernd.Kaercher@dlr.de

U. Schumann (ed.), Atmospheric Physics, Research Topics in Aerospace,DOI: 10.1007/978-3-642-30183-4_3, � Springer-Verlag Berlin Heidelberg 2012

37

visibility, and adverse health effect. However, spectacular atmospheric effects suchas colored sunsets are also associated with aerosol particles; oceanic algae arefertilized by dust-bound iron transported from the great deserts of Earth, and thepolar stratospheric ozone holes are also triggered by aerosol particles.

The ocean is one major source of natural aerosols with an estimated annualemission of 1 000 to 10 000 Tg (= 1012 g = 1 megaton) of material (Raes et al.2000; Seinfeld and Pandis 2006). The exchange of particulate matter between airand sea contributes to the global cycles of carbon, nitrogen, and sulfur particlessuch as dimethyl sulfide (DMS), produced by phytoplankton. Ocean water and seasalt are transferred to the atmosphere through sea spray and air bubbles at the seasurface. As a droplet of water evaporates, the salt is left suspended in the atmo-sphere, forming maritime aerosols, or sea salt. Another major natural source ofparticulate mass is windblown mineral dust from dry continental regions likedeserts or semi-arid areas. Estimates of the global mass of dust particles releasedper year into the atmosphere vary between 1 000 and 5 000 Tg (Raes et al. 2000).Other significant sources of airborne particles are biomass burning from forest orsavannah fires, volcanic eruptions, and anthropogenic pollution from industrialemissions and traffic.

Corresponding to the major sources, the global distribution of airborne particles interms of mass has its hot spots over or near the world’s large deserts, along the stormtracks of the southern hemispheric oceans south of the Americas and Africa, overtropical rainforests and over the highly industrialized regions of North America,Europe and South East Asia. Figure 3.2 shows the global distribution of aerosoloptical depth AOD (see below for a definition) as a measure of the atmosphericaerosol load, as recorded by the NASA satellite instrument MODIS (ModerateResolution Imaging Spectroradiometer, see http://modis.gsfc.nasa.gov/).

Fig. 3.1 Major sources of atmospheric aerosols and their impacts on the global climate and airquality; reprinted by permission from Macmillan Publishers Ltd:[NATURE] (Kolb 2002),copyright (2002)

38 A. Petzold and B. Karcher

3.2 Definitions and Properties

Atmospheric aerosols are defined as a suspension of fine particles consisting ofsolid or liquid compounds of various origins in air. Water-based airborne particleslike cloud drops or ice crystals, however, are not considered as aerosols but ashydrometeors. While gaseous atmospheric constituents are usually characterizedby their mixing ratio compared to air, a description of aerosols requires the con-sideration of a set of physical and chemical properties.

One key property is the size of the particles, given as particle diameter D or astypical dimension in the case of irregularly shaped particles. Particle size spansover multiple orders of magnitude from a few nanometers (1 nm = 10-9 m) forparticles freshly produced from gaseous precursors by gas-to-particle conversionup to almost 1 mm for large dust particles. For comparison, a typical ‘‘air’’molecule has a diameter of *0.4 nm. Atmospheric aerosol always containsparticles of different sizes and is therefore classified as polydisperse. In contrast, anaerosol composed of particles of a single size would be classified as monodisperse.Generally, the distribution of particle sizes present in an aerosol is given as theprobability of occurrence of particles with diameters at certain size intervals. Thisproperty is called the particle size distribution function.

In general terms, key properties required for a full characterization of anaerosol are:• particle size distribution• total number of particles per reference volume (number concentration N given

in cm-3)• total mass of suspended particulate matter per reference volume (mass con-

centration M given in lg m-3 or ng m-3

Fig. 3.2 Global distribution of aerosol optical depth (AOD) for the years 2005 to 2010 takenfrom the MODIS satellite instrument; since the MODIS instrument cannot resolve AOD overhighly reflecting surfaces like deserts or snow, no MODIS data are available for these regions.Adapted from http://en.wikipedia.org/wiki/File:Modis_aerosol_optical_depth.png

3 Aerosols in the Atmosphere 39

• particle density, ranging from 1.0 g cm-3 for organic aerosol to 2.6 g cm-3 formineral dust

• particle shape (spherical, irregularly shaped, agglomerates)• particle phase (liquid, solid)• chemical composition

Clean continental air often contains less than a few thousand particles per cubiccentimeter (cm-3), while polluted continental air may reach concentration valuesas high as 50 000 cm-3, and in urban air, particle number concentrations mayexceed 105 cm-3. Natural sources of combustion particles like forest fires may alsoexceed particle number concentrations of 105 cm-3 (Andreae and Merlet 2001).Desert air has about 1 000 cm-3 on average. Clean marine air generally has a few100 s cm-3 but may reach values well above 10 000 cm-3 in cases of new particleformation events. These are all near-surface concentration values. In the freetroposphere a typical background number concentration is 300 cm-3. Polaraerosols usually do not exceed number concentrations of 200 cm-3 on averageabove 1 km altitude, strongly decreasing to values of several 10 cm-3 near thesurface (Seinfeld and Pandis 2006). The background stratospheric aerosol layer ischaracterized by a typical number concentration and mean diameter of accumu-lation mode particles of 1 000 cm-3 and 0.14 lm, respectively. In honor of its firstinvestigator (C. Junge, 1912–1996), the layer of stratospheric aerosol is alsonamed the Junge layer, see Fig. 3.1. For comparison, 1 cm3 of air at sea levelcontains about 3 9 1019 gas molecules.

A common representation of the particle size distribution function is the so-called log-normal distribution (Hinds 1999)

n ðlog DÞ ¼ Nffiffiffiffiffiffi

2pp

log r g

exp � 12ðlog D � log Dmed Þ 2

log2 rg

!

which assumes a Gaussian (bell-curve shaped) or normal distribution of the log-arithm of the particle diameter, log D, around the logarithm of its count median (ornumber median) diameter, log Dmed, with a geometric standard deviation of thedistribution of rg. The distribution function can also be expressed on the basis ofthe natural logarithm. It is characterized by three parameters: total number con-centration N, Dmed and rg. For its simple parameterization and its straightforwardcharacter this distribution is widely used in Earth system models.

Atmospheric aerosol can be regarded as a superposition of four particle modes,each described by a single log-normal size distribution. Figure 3.3 is a schematicrepresentation of the modes of a generic atmospheric aerosol of the atmosphericboundary layer, the size ranges covered by the different aerosol modes, the majorchemical compounds and the respective size distribution.

The nucleation mode (D B 20 nm) contains particles freshly formed fromgaseous precursors by gas-to-particle conversion, or homogeneous particlenucleation; most aerosol particles in the nucleation mode are composed of sulfuriccompounds and hydrocarbon compounds.

40 A. Petzold and B. Karcher

The Aitken mode (kursiv setzen = 20–100 nm) includes particles formed bycoagulation of nucleated particles and condensation of vapors on already existingparticles, or emitted directly into the atmosphere; it was named after the Scottishmeteorologist and physicist John Aitken (1839–1919), who was one of the foundersof aerosol science. Nucleation mode and Aitken mode as defined here are combinedby some authors into a single Aitken mode which then covers all particles smallerthan 100 nm in diameter.

The accumulation mode (D = 0.1–1.0 lm) consists of particles formed fromthe Aitken mode by particle coagulation and particles emitted directly from pri-mary sources like combustion of vegetation or fossil fuels; for atmosphericaerosols the accumulation mode forms the sink of particles growing from nucle-ation via the Aitken mode into the accumulation mode.

The coarse mode contains all particles larger than 1 lm in diameter; theseparticles are generated mainly by mechanical processes like wind-blown dust, seaspray, or plant debris, or are emitted from volcanoes or large fires. Because of thedifferent formation pathways, coarse mode particles are separated from smallerparticles with respect to their chemical composition. Bioaerosols like spores,bacteria or pollen are also mainly part of the coarse mode, but may be part of theaccumulation mode as well.

The major chemical constituents of sub-lm aerosol in the lower troposphere aresulfates (SO4

2-), nitrates (NO3-), ammonium (NH4

+), black carbon (BC), andparticulate organic matter (POM) partially emitted directly from the source or

Fig. 3.3 Schematic representation of the modes of a typical atmospheric aerosol of the planetaryboundary layer for different particle diameters D Left key chemical compounds and particlegeneration processes; right size ranges

3 Aerosols in the Atmosphere 41

formed as secondary organic aerosol (SOA). Super-lm particles in the coarse modeare mainly associated with mineral dust, sea salt, volcanic ash and biogenic particleslike spores, bacteria or plant debris. In the free troposphere and above, the concen-tration of coarse mode particles in general and of source-related compounds likePOM and BC is much smaller than in the boundary layer. Stratospheric aerosol ispredominantly composed of sulfuric acid droplets and volcanic ash particles.

3.3 Sources, Atmospheric Processing, and Sinks

Atmospheric aerosol particles arise from a variety of sources and formation pro-cesses, as discussed in the previous section. They can be classified as being ofprimary or secondary origin according to their history of formation. Primaryparticles are emitted directly from the source while secondary particles form fromgaseous precursors which may have been emitted from ground-based sources orformed in the atmosphere by chemical reactions. Primary particles include soilmaterial and minerals, sea salt, biogenic particles such as pollen or bacteria, andsoot (elemental carbon including organic material). Examples for secondary par-ticles are sulfates originating from DMS (dimethylsulfide) or SO2, nitrates formedfrom NOx, and secondary organic aerosol (SOA) particles formed from biogenic oranthropogenic gas-phase hydrocarbons.

Source processes for primary particles are manifold (e.g., anthropogenic fossilfuel combustion, biomass burning, volcanic eruptions, wind-blown mobilization ofsoil material, desert dust storms, wind-driven sea spray production) but are mostlyrelated to particle emissions from Earth’s surface, with the exception of emissionsfrom cruising aircraft, which is the only source of primary aerosol particles in theupper troposphere (Petzold et al. 1999). The formation of secondary particles bygas-to-particle conversion is solely associated to nucleation. Secondary particleformation can occur in the marine boundary layer, over rural continental sites, inthe polluted urban atmosphere, and in the free troposphere (Raes et al. 2000;Kulmala et al. 2001).

Secondary particle formation by homogeneous nucleation always includesH2SO4, organic vapors and H2O (Kulmala et al. 2001). New particle formationtakes place in the form of bursts in clean air characterized by low particle numberconcentrations. Correspondingly, the aerosol surface area provided to condensablegases by these few airborne particles is low and supersaturation with respect tothose gases can develop. In contrast, at high levels of available aerosol, surfacearea condensation of molecules on pre-existing particles is favored over newparticle production, and this quenching of nucleation is proportional to the aerosolsurface area in most cases. Recent speculations about a link between galacticcosmic rays associated to the solar cycle and the formation of new particles in theupper troposphere could not be validated by observations performed during theEuropean field study EUCAARI, to which the DLR’s Falcon aircraft significantlycontributed (Kulmala et al. 2010).

42 A. Petzold and B. Karcher

Once released into the atmosphere or formed by gas-to-particle conversion,small particles further grow by condensation of molecules from the gas phase or bythe collision of small and large particles (so-called coagulation), which leads to theevolution of larger but fewer aerosol particles of mixed chemical composition.Figure 3.4 shows a map of states of atmospheric aerosol as proposed by Schröderet al. (2002). Particle nucleation is favored by low number concentrations of agedparticles in the accumulation mode, i.e., low surface area of already existingparticles. Aging by coagulation/accumulation drives the aerosol to a well-agedstate characterized by a high number concentration of accumulation mode particlesand a medium number concentration of fresh nucleation mode particles. If theseaged aerosol particles interact with clouds, preferably larger particles from theaccumulation mode with D [ 0.1 lm act as nuclei for cloud drop or ice formation.These particles are then removed from the atmosphere via scavenging by clouddroplets and wet deposition, leaving behind an aerosol characterized by lownumber concentrations of both fresh and aged particles, which then favors theformation of new particles by nucleation. The scheme shown in Fig. 3.5 sum-marizes how aerosol particles are processed in the atmosphere (Hinds 1999).

One specific mechanism of reversible particle growth in the atmosphere isassociated with the uptake of water vapor by water-soluble particle constituents ina humid atmosphere. This so-called hygroscopic growth of aerosol particles isattracting increasing attention because it governs the activation of aerosol particlesinto cloud condensation nuclei (CCN), initiating cloud formation. It also enhancesthe climate impact (see Sect. 3.6) of aerosol particles by significantly increasingthe particle size and thus the scattering of sunlight. To give an example, sulfuricacid and ammonium sulfate particles, which are quite common constituents ofatmospheric aerosol, more than double their diameter at 95 % relative humiditycompared to the dry state simply by taking up water from the surrounding air(Seinfeld and Pandis 2006). Contemporary concepts for understanding hygro-scopic growth and cloud droplet formation are sufficiently explained by the Köhlertheory (Pruppacher and Klett 1997).

Fig. 3.4 Schematicillustrating the map ofprincipal states oftropospheric aerosols: NUCfresh, nucleated state, ACCaged, accumulated state, SCAcloud-processed state. TheMED state represents typicalbackground aerosol; adaptedfrom Schröder et al. (2002)

3 Aerosols in the Atmosphere 43

3.4 Spatial Distribution of Aerosols

Driven by the spatial distribution of sources, sinks and transport, atmosphericaerosol exhibits a certain vertical number concentration profile. Over more than adecade of intense field study work at the DLR Institute of Atmospheric Physics, alarge set of data on the vertical distribution of aerosol particles was collected.Figure 3.6 shows examples of measurements taken at mid-latitudes (a) in the cleanbackground atmosphere over the southern tip of South America near Punta Arenas,Chile, (b) in the polluted background aerosol over the northern hemisphere nearPrestwick, Scotland, with both data sets emerged from the European field studyINCA (Minikin et al. 2003), and (c) over continental Europe related to the Germanfield study LACE (Petzold et al. 2002) near Berlin, Germany. Number concen-tration data are given for standard conditions (STP: 273.15 K, 1013.25 hPa) sincein this case number concentrations are conserved for the vertical motion of therespective air parcel.

In all cases, total particle number concentrations are highest near Earth’s sur-face, in accordance with the dominating aerosol sources there. Through the lowestpart of the tropospheric column, the planetary boundary layer, the vertical profileis usually characterized by a strong negative gradient. Across the transition fromthe boundary layer into the free troposphere, the particle number concentrationdevelops into a well-mixed free tropospheric aerosol with only weak verticalstructure. In the polluted northern hemisphere these profiles show minimum values

Fig. 3.5 Scheme of the microphysical processes that influence the size distribution and chemicalcomposition of atmospheric aerosol; adapted from Hinds (1999)

44 A. Petzold and B. Karcher

in the middle free troposphere at pressure levels around 800 to 500 hPa, which isthe altitude range for intense aerosol-cloud interaction. Near the tropopause theparticle number concentration increases again, which can be attributed to particlenucleation in the upper troposphere and the lack of efficient aerosol removalprocesses. As can be taken from Fig. 3.6, global aerosol models, see, e.g., Aquilaet al. (2011), successfully reproduce aerosol vertical profiles in the pollutednorthern hemisphere, while they produce less good agreement for the less pollutedsouthern hemisphere.

Specific primary aerosol particles emitted from surface-bound sources andtransported upward by lifted air parcels show even more pronounced vertical profilesthan total aerosol, as shown in Fig. 3.6. For example, black carbon mass concen-trations observed on average in the free troposphere above the planetary boundarylayer are lower by two orders of magnitude and beyond compared to the respectivesurface data (Petzold et al. 2007; Aquila et al. 2011; Hamburger et al. 2011).

Another aerosol type of particular interest is mineral dust, which was exten-sively studied during the German Saharan Mineral Dust Experiment SAMUMwhere DLR was responsible for the airborne observations using aircraft. Figure 3.7shows an illustrative example of the vertical distribution of mineral dust trans-ported from the African mainland to the Cape Verde islands where the measure-ments were performed. Data shown here originate from a high spectral resolutionlidar which is capable of actively sounding aerosol extinction from an airborneplatform. The plot shows the effect of an island on vertical dust distribution.

observation, median

EMAC/MADE-IN, 25%-75% perc.

EMAC/MADE-IN, median

observation, 10%-90%perc.

(a) (b) (c)

Fig. 3.6 Vertical profiles of the number concentration (converted to standard temperature andpressure (STP) conditions: 273 K, 1 013 hPa) of aerosol particles larger than 5 nm as measuredby Minikin et al. (2003) during INCA over Punta Arenas (Chile) in March–April 2000 and overPrestwick (Scotland) in October 2000 (a and b, respectively) by Petzold et al. (2002) duringLACE over central Europe in August 1998 (c), and from corresponding climate modelsimulations (Aquila et al. 2011)

3 Aerosols in the Atmosphere 45

The mineral dust layer extends from sea level to about 2.5 km and is topped by athin layer of biomass burning aerosol.

At the upwind side (the main wind direction is from left to right) the island liftsthe dense part of the mineral dust layer over the top of the island’s mountainsorographically, while at the downwind side a clear three-layer structure develops.SAMUM also demonstrated that over the desert mineral dust may extend up to analtitude of 4 km above ground (Weinzierl et al. 2009) with a sharp upper boundwhere total particle number drops off by more than one order of magnitudeover *100 m altitude. If dust particles are lifted to higher altitudes they serve asefficient ice nuclei initiating the formation of mixed-phase clouds containing bothliquid drops and ice crystals (Ansmann et al. 2008). SAMUM was among the firstintense field studies which provided this kind of information on the vertical dis-tribution of mineral dust (Ansmann et al. 2011).

3.5 Atmospheric Lifetime and Long-Range Transport

The lifetime of atmospheric aerosol particles depends on their chemical nature andsize and on the altitude range. Airborne particles in the troposphere have typicallifetimes of 3–10 days on average (Andreae 2007). After 30 days, 95 % of theparticles are removed by coagulation with other particles or cloud drops (so-calledscavenging), and wet deposition via precipitation. Aerosol particles in the lowerstratosphere have a much longer lifetime of up to 1 year before they penetrate thetropopause and undergo tropospheric loss processes. Figure 3.8 shows the

Fig. 3.7 Illustration of island effects on mineral dust layers over the ocean. The left plot showsthe backscatter ratio at 532 nm measured by the high spectral resolution lidar (HSRL) during anoverpass over Santiago Island (Cape Verde Islands) between 16:05 and 16:22 UTC 25 January2008. MULIS refers to a ground-based lidar operated at the airport of Praia. The inset shows theflight track during the sounding; the flight direction was from SE to NW (Weinzierl et al. 2011)

46 A. Petzold and B. Karcher

atmospheric lifetime of individual particles for the different altitude ranges of theatmosphere (Jaenicke 1980). As can be deduced from this figure, smaller particlesare efficiently removed by coagulation and large particles by sedimentation, whileparticles of the accumulation mode size range are efficiently removed only by wetdeposition via aerosol-cloud interaction processes.

During its atmospheric lifetime, aerosol can easily be transported on intra-hemispheric scales during its residence time of more than one week in the freetroposphere. Intercontinental transport of particles is observed frequently asso-ciated with lifting in warm conveyor belts associated with low pressure systemsand long-range transport in the free troposphere. In 10 years of regular mea-surements of aerosols over Leipzig, Germany, by means of Raman lidar (Mattiset al. 2008), the authors observed pollution layers above the continental boundarylayer in 43 % of the cases. The lofted layers largely resulted from intercontinentallong-range transport. In 19 % of all regular measurements, free-troposphericpollution was advected from North America. Forest-fire smoke from Canada andanthropogenic pollution from urban areas of the USA and Canada were the sourcesof the particle layers.

From a European perspective, extreme cases of lifted aerosol layers in the freetroposphere are strong fire plumes transported from North American sources(Petzold et al. 2007), Saharan dust layers advected from North Africa (Mattis et al.2008), and volcanic ash plumes (Schumann et al. 2011). Even air pollutiontransportation over the Mediterranean Sea (Lelieveld et al. 2002) and pollutionplumes from southern and eastern Asia, including regions in India and China

Fig. 3.8 Atmosphericlifetime for individualparticles with indicateddominant loss processes inthe troposphere; adapted fromJournal of Aerosol Science,11, Jaenicke, R., Atmosphericaerosols and global climate,577–588, copyright (1980)with permission fromElsevier

3 Aerosols in the Atmosphere 47

(Stohl et al. 2007), were observed over Europe in measurements of the DLRFalcon research aircraft after more than one week of atmospheric transport.

We observed a particularly spectacular example of aerosol transport and mixingduring the SAMUM project when investigating the mixing of the urban pollutionplume over the city of Dakar (Senegal) with mineral dust. Figure 3.9 shows across-section of the Dakar plume with clean dust moving in from the east (left partof the plot) and mixing with urban pollution during westerly transport from thecontinent out to the ocean. The origin of the distance scale is arbitrarily set to apoint west of Dakar. The shoreline is situated at distance 60 km. Signatures ofurban pollution were still found 100 km offshore after several hours of atmo-spheric transport. However, the data analysis clearly showed that the impact ofurban emissions on mineral dust via mixing with black carbon (soot) is limited toless than one day of the plume’s lifetime (Petzold et al. 2011).

3.6 Aerosols in the Climate System

Anthropogenic aerosol particles influence the radiative budget of the Earth-atmosphere system in two different ways. The first is the direct effect, wherebyparticles scatter and absorb solar and thermal infrared radiation and thereby alterthe radiative balance of the Earth-atmosphere system or, equivalently, the plane-tary albedo (= reflectivity of the Earth-atmosphere system). The second are calledthe indirect effects, since particles modify the microphysical and hence the radi-ative properties and lifetime of clouds (Haywood and Boucher 2000). Figure 3.10

Fig. 3.9 High Spectral Resolution Lidar (HSRL) cross-section of the Dakar plume mixing intothe mineral dust layer; the city of Dakar is situated at 50 km distance from the arbitrarily chosenpoint of origin of the distance scale, whereas the plume appears to originate from an area 30 kminland of Dakar; this local shift is caused by sea-breeze circulation (see text for details); atdistance 150 km a residual plume is observed. Viewing direction is south, direction to west isindicated (Petzold et al. 2011)

48 A. Petzold and B. Karcher

schematically summarizes the effects of aerosols and clouds on Earth’s radiativebalance; see also Fig. 3.1 for illustration.

Radiative effects are evaluated in terms of radiative fluxes in the solar (solidarrows in Fig. 3.10) and thermal infrared (broken arrows in Fig. 3.10) spectralregions at the top of the atmosphere (TOA). If the anthropogenic aerosol or theclouds modified by the anthropogenic aerosol increase the net (i.e., the solar plusthermal infrared) radiative flux at TOA, more energy is reflected back into spacethan in the unperturbed case without anthropogenic aerosol in the atmosphere, andthe overall effect is a cooling. If the net radiative flux at TOA is reduced by light-absorbing aerosol particles or by clouds, the resulting effect is a heating of theEarth-Atmosphere system because more radiative energy remains in the systemcompared to the unperturbed case.

According to Fig. 3.10 the effects of atmospheric aerosol particles on planetaryalbedo can be sketched as follows: Incoming sunlight is partially scattered (solidarrows) and partially absorbed (broken arrows) and subsequently transferred asheat to the surrounding air. These processes are described by the respectivecoefficients rsp for particle scattering and rap for particle absorption, which arereported as the removed fraction of light by the respective processes per unitlength of the light path. The sum of both is the extinction coefficient rep, whichexhibits a distinct vertical profile with its major contribution arising from atmo-spheric boundary layer aerosol; see also the discussion on the vertical distributionof aerosols in Sect. 3.4. The free troposphere contributes significantly only if loftedaerosol layers from long-range transport events are embedded.

The fundamental aerosol parameters governing the aerosol impact on climateforcing are the aerosol optical depth and the ratio of particle scattering toextinction at a wavelength of 550 nm, the so-called single-scattering albedo x0

(Haywood and Shine 1995). Aerosol optical depth (AOD) is defined as the

Fig. 3.10 Left schematic of direct and indirect radiative effects of aerosols and clouds on theglobal radiative balance in terms of coefficients r for particle scattering (sp), absorption (ap), andextinction (ep). S surface, BL planetary boundary layer, FT free troposphere, TOA top ofatmosphere. Right satellite image of a mineral dust outbreak from the northern Sahara towards theMediterranean Sea; seen on October 26, 2007 by the MODIS instrument on NASA’s Terrasatellite (http://earthobservatory.nasa.gov/NaturalHazards/view.php?id=19264)

3 Aerosols in the Atmosphere 49

extinction coefficient integrated from the surface to TOA for a unit cross-sectionfootprint. AOD is a common product delivered by space-borne sensors for theglobal mapping of atmospheric aerosol; see also Fig. 3.2 for a global distributionof AOD.

Exclusively light-scattering particles are characterized by x0 = 1.0, whilepartially light-absorbing aerosol particles are associated with values of x0 \ 1. Asa rule-of-thumb, the critical value x0

* which separates net heating from net coolingaerosol is x0

* % 0.87 (Hansen et al. 1997). Strongly light-absorbing aerosol par-ticles with x0 \ x0

* have a net-heating effect, while weakly absorbing aerosolswith x0 [ x0

* still have a net cooling effect dominated by their contribution tolight-scattering. Tropospheric aerosols that have a substantial anthropogeniccomponent include sulfate particles, black carbon, organic carbon, mineral dust,and nitrate aerosol. From these aerosol components black carbon and mineral dustare the only constituents which absorb solar radiation in the visible spectrum(Petzold et al. 2011), while sulfate and nitrate particles are purely light-scattering.Since, in particular, the efficiency of particles for scattering light strongly increaseswith increasing particle size (Hinds 1999), the above-mentioned humidity-relatedgrowth of particles considerably enhances the net cooling effect of water-solubleaerosols in a humid atmosphere (Seinfeld and Pandis 2006). Furthermore, the roleof organic aerosols, also named ‘‘brown carbon,’’ in the climate system is stillwidely unknown (Andreae and Gelencsér 2006).

Fig. 3.11 Schematic diagram showing the various radiative mechanisms associated with cloudeffects that have been identified as significant in relation to aerosols (modified from Haywood andBoucher, 2000). The small black dots represent aerosol particles; the larger open circles clouddroplets. Straight lines represent the incident and reflected solar radiation, and wavy linesrepresent terrestrial radiation. The filled white circles indicate cloud droplet number concentra-tion (CDNC). The unperturbed cloud contains larger cloud drops as only natural aerosols areavailable as cloud condensation nuclei, while the perturbed cloud contains a greater number ofsmaller cloud drops as both natural and anthropogenic aerosols are available as cloudcondensation nuclei (CCN). The vertical grey dashes represent rainfall, and LWC refers to theliquid water content; Fig. 2.10 of IPCC report 2007 (IPCC 2007)

50 A. Petzold and B. Karcher

Figure 3.11 shows a schematic diagram of the various radiative mechanismsassociated with aerosol and cloud effects that have been identified as significant(Haywood and Boucher 2000). Clouds interact with radiation both in the solarspectrum and in the thermal infrared spectrum. In the solar spectrum clouds increaseplanetary reflectivity and reduce the fraction of radiation reaching the surface. Theunperturbed cloud contains larger cloud drops as only natural aerosols are availableas cloud condensation nuclei. The perturbed cloud contains a greater number ofsmaller cloud drops as both natural and anthropogenic aerosols are available as cloudcondensation nuclei. In the infrared spectrum clouds efficiently absorb thermalradiation emitted from the surface and re-emit thermal radiation corresponding totheir actual temperature. Therefore, thermal effects are largest for cold high-altitudecirrus clouds, while warm low-level clouds mainly contribute to an increase inplanetary albedo. Investigation of aerosol-cloud interactions and resulting impactson the global climate system is one of the hot spots of current research.

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