Atmospheric Teleconnection over Eurasia Induced by Aerosol Radiative Forcingduring Boreal Spring

MAENG-KI KIM,* WILLIAM K. M. LAU,� MIAN CHIN,# KYU-MYONG KIM,@ Y. C. SUD,& AND

GREG K. WALKER**

*Department of Atmospheric Science, Kongju National University, Gongju, South Korea�Laboratory for Atmospheres, NASA Goddard Space Flight Center, Greenbelt, Maryland

#Atmospheric Chemistry and Dynamics Branch, Laboratory for Atmospheres, NASA Goddard Space Flight Center,Greenbelt, Maryland

@Science Systems and Applications, Inc., Lanham, Maryland&Climate and Radiation Branch, Laboratory for Atmospheres, NASA Goddard Space Flight Center, Greenbelt, Maryland

**SAIC/General Sciences Operation, Beltsville, Maryland

(Manuscript received 18 March 2005, in final form 22 November 2005)

ABSTRACT

The direct effects of aerosols on global and regional climate during boreal spring are investigated basedon numerical simulations with the NASA Global Modeling and Assimilation Office finite-volume generalcirculation model (fvGCM) with Microphyics of Clouds with the Relaxed–Arakawa Schubert Scheme(McRAS), using aerosol forcing functions derived from the Goddard Ozone Chemistry Aerosol Radiationand Transport model (GOCART).

The authors find that anomalous atmospheric heat sources induced by absorbing aerosols (dust and blackcarbon) excite a planetary-scale teleconnection pattern in sea level pressure, temperature, and geopotentialheight spanning North Africa through Eurasia to the North Pacific. Surface cooling due to direct effects ofaerosols is found in the vicinity and downstream of the aerosol source regions, that is, South Asia, East Asia,and northern and western Africa. Significant atmospheric heating is found in regions with large loading ofdust (over northern Africa and the Middle East) and black carbon (over Southeast Asia). Paradoxically, themost pronounced feature in aerosol-induced surface temperature is an east–west dipole anomaly withstrong cooling over the Caspian Sea and warming over central and northeastern Asia, where aerosolconcentrations are low. Analyses of circulation anomalies show that the dipole anomaly is a part of anatmospheric teleconnection pattern driven by atmospheric heating anomalies induced by absorbing aerosolsin the source regions, but the influence was conveyed globally through barotropic energy dispersion andsustained by feedback processes associated with the regional circulations.

The surface temperature signature associated with the aerosol-induced teleconnection bears strikingresemblance to the spatial pattern of observed long-term trend in surface temperature over Eurasia.Additionally, the boreal spring wave train pattern is similar to that reported by Fukutomi et al. associatedwith the boreal summer precipitation seesaw between eastern and western Siberia. The results of this studyraise the possibility that global aerosol forcing during boreal spring may play an important role in spawningatmospheric teleconnections that affect regional and global climates.

1. Introduction

Recent studies have shown that aerosols may play animportant role in climate change through their interac-tion with the global water and energy cycles (e.g., Ja-cobson 2001a,b, 2002; Menon et al. 2002; Lohmann and

Lesins 2002; Roberts and Jones 2004; Ramanathan etal. 2001). The effect of aerosols on the earth’s radiativebudget is not only limited to cooling by scattering butalso heating by absorption of solar radiation, dependingon aerosol types. Because of their ability to depletesurface insolation from either scattering or absorption,all aerosols cause cooling at the earth surface—the so-called “solar dimming” effect (Ramanathan et al. 2005).However, the sign of atmospheric temperature changeinduced by aerosol forcing can vary depending on theaerosol types, their elevation, and reflectivity of the

Corresponding author address: Dr. K. M. Lau, Laboratory forAtmospheres, NASA Goddard Space Flight Center, Code 613,Greenbelt, MD 20771.E-mail: lau@climate.gsfc.nasa.gov

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© 2006 American Meteorological Society

JCLI3871

underlying surfaces (Kaufman et al. 2002). Moreover,regional climate can also be influenced by circulationchanges induced by aerosol radiative forcing. Thusaerosol-induced local heating and cooling, as well as theeffects of induced circulation changes, must be includedin studies of climate responses to aerosol forcing.

For studies of global impact of aerosols on climateand the water cycle, including feedback between forc-ing and response, a global climate model is a very use-ful, albeit imperfect, tool. There have been numerousstudies to estimate the direct and indirect effects ofaerosols from optical and physical properties of clouds,and surface reflectivity, using global models with ideal-ized global aerosol forcings (Coakley and Cess 1985;Chylek et al. 1995, 1996; Hansen et al. 1997a; Quijano etal. 2000; Kay and Box 2000; Weaver et al. 2002; Nish-izawa et al. 2004; Cook and Highwood 2004; Chung etal. 2002). However, relatively less attention has beendevoted to studies of regional and global climate re-sponses by radiative heating and cooling induced bythree-dimensional aerosol loading, including feedbackfrom atmospheric water cycle dynamics. Recently, ageneral circulation model (GCM) simulation study byMenon et al. (2002) showed that there is a strong sur-face warming in remote regions of Eurasia arising fromregional sources of aerosols in China and India. How-ever, model aerosol forcing and impacts have large un-certainties, primarily because of the inadequate repre-sentation of aerosol–cloud interaction physics, com-pounded by the poorly defined aerosol opticalproperties and vertical distribution (Hansen et al.1997b). More recently, global chemical transport mod-els have simulated the three-dimensional climatologicalstructure of aerosol distributions as well as their opticalproperties for various aerosol types, which comparedreasonably well against available observations, in selectregions (Chin et al. 2002, 2003; Takemura et al.2002a,b). Outputs from global chemical transport mod-els can be translated into realistic, four-dimensional(time and space) radiative forcing effects for GCMs.This approach has been adopted in the present study.

In view of the inhomogeneous distribution of aerosolsource, transports, associated thermal forcing, and cir-culation changes, there can be no simple balance ofaerosol, greenhouse gas, and dynamical forcing vis-à-visregional climate change. For long-term climate change,almost all extratropical land regions show pronouncedwarming during boreal winter (Houghton et al. 2001).However during boreal spring [March–May (MAM)],both warming and cooling of extratropical land regionsare noted. It has been suggested that the spring surfacetemperature pattern may be related to changes in large-scale circulation, possibly related to aerosol effects

(M.-K. Kim et al. 2004). Figure 1 shows the observedlong-term mean boreal spring surface temperatureanomaly (1980–98 minus 1961–79) for Africa and Eur-asia from surface climate data (New et al. 2000) forAfrica and Eurasia. Large areas of cooling appear insouthern and eastern Asia, where aerosol loadings areknown to be high. However, the most pronounced sig-nal appears to be a dipolelike pattern with maximumsurface cooling around the Caspian Sea and warming inthe continental region around Lake Baikal. Yet, thesetwo regions are located outside the large aerosol sourceregions, that is, China, India, and the Sahara. Whatthen is the cause of this dipolelike temperatureanomaly? Are remote forcing and large-scale circula-tion changes induced by aerosols a contributing factor?

The aforementioned considerations have motivatedus to investigate regional climate impact of global aero-sol forcing on climate using a GCM combined with re-alistic global aerosol forcing derived from chemistrytransport models. In this paper, we will limit our studyto the “direct effect” of aerosols including the full dy-namical feedback involving large-scale circulation,clouds, and convection. Investigation of “indirect ef-fect” is an ongoing effort that is outside the scope ofthis work. Here, as a first step, we will focus on theboreal spring to identify possible signals of regional andglobal atmospheric response to aerosol forcing. Borealspring is chosen for this study because it is the time ofyear when a large number of land-based aerosol par-ticles are injected into the atmosphere from biomassburning in Southeast Asia and Africa and dust stormsin the deserts of Sahara and in northern Asia, as well asurban pollution in industrial regions of East Asia andEurope (Takemura et al. 2000; Zhang et al. 2002; Kauf-man et al. 2002; Novakov et al. 2003; Duncan et al.2003; Liu et al. 2005). These aerosols will have a longerresidence time because the Northern Hemisphere as awhole has less precipitation before the summer mon-

FIG. 1. Spatial pattern of long-term change in observed borealspring land surface temperature for the last 40 years. The changeis calculated as the difference in the mean between 1980–98 and1961–79. Positive values over 0.8°C per 20 yr are shaded.

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soon season so that wet deposition is at a minimum inboreal spring compared to the rest of the season. Inaddition, dynamical forcings are also relatively weakduring the spring season compared to winter. Hence,thermally induced atmospheric responses by aerosolsare more readily detectable. Moreover, the dynamicalresponses, in all likelihood, would precondition theevolving climate to influence the subsequent evolutionof the anomalous summer monsoon (Lau et al. 2006).

The model and simulation experiments are describedin section 2. In sections 3 and 4, we describe globalaerosol forcing, global response, and possible mecha-nisms of regional water cycle feedback. Section 5 dis-cusses the special role of absorbing aerosols, that is,black carbon and dust, on the regional climate changes.The summary and additional discussions are presentedin section 6.

2. Model experiments

The climate impacts of aerosol radiative forcing areinvestigated using the NASA Goddard Modeling andAssimilation Office finite-volume General CirculationModel (fvGCM) with Microphysics of Clouds with theRelaxed Arakawa–Schubert Scheme (McRAS). Thisversion of the fvGCM has a horizontal resolution of 2°by 2.5° and 55 vertical levels. The finite-volume dy-namical core has a horizontal discretization built uponthe flux-form semi-Lagrangian transport algorithms(Lin and Rood 1996, 1997). The vertical structure isbased on the Lagrangian control volume concept (Lin1997). This alleviates some of the problems associatedwith sigma, pressure, or isentropic coordinates, therebyincreasing the physical integrity and computational ef-ficiency of the model (Chang et al. 2001). The McRASmicrophysics cloud scheme takes into account moistprocesses, microphysics of clouds, and cloud–radiationinteraction in GCMs. It allows for three cloud types:convective, stratiform, and boundary layer (Sud andWalker 1999a,b, 2003). All clouds convect and advect,as well as diffuse both horizontally and vertically, withfully interactive cloud microphysics throughout the lifecycle of the cloud, while the optical properties of cloudsare derived from the statistical distribution of hydrom-eters and idealized cloud geometry. An evaluation ofMcRAS in a single-column model with the Global At-mospheric Tropical Experiment (GATE) Phase IIIdata has shown that, together with the rest of the modelphysics, McRAS can simulate the observed tempera-ture, humidity, and precipitation without discerniblesystematic errors (Sud and Walker 1999a). The fvGCM

is coupled to a land surface model (Bonan 1996) and aplanetary boundary model (Holtslag and Boville 1993).The radiative transfer model is described by Chou andSuarez (1999) and Chou et al. (2001).

Aerosol optical thicknesses of each of the five aero-sols, that is, black carbon, organic carbon, dust, sulfate,and sea salt, are obtained in the form of three-dimensional monthly values from emission inventoriesand global transport calculations from the GoddardChemistry Aerosol Radiation and Transport (GOCART)model (Chin et al. 2002, 2003) for two years (2000–2001). The extinction coefficient, single scattering al-bedo (SSA), and an asymmetric factor of each aerosoltype are computed at 11 broad wavelength bands fromMie theory as a function of the ambient relative humid-ity. The optical parameters of dust are assumed to be afunction of wavelength only. The SSA used for dust isin the range 0.90–0.92 at 0.55 �m averaged over thedust size bins (Carlson and Benjamin 1980; Tegen andLacis 1996; Quijano et al. 2000; Takemura et al. 2002a;Nakajima et al. 2003; D.-H. Kim et al. 2004; Miller et al.2004), corresponding to moderate absorption. Morediscussion regarding uncertainties in estimation of theSSA of dust is included in section 6.

We have performed two 10-yr-long baseline modelexperiments, with all aerosols (AA) from the GOCARTmodel climatology, and one with no aerosols (NA). InAA, the monthly aerosol radiative forcing functions foreach aerosol type, derived from the 2-yr climatology,are interpolated into daily values, which are prescribedevery year during the 10-yr integration. In both experi-ments, the sea surface temperature (SST) is prescribedusing the weekly SST from September 1987 to Decem-ber 1996 interpolated to daily values. All greenhousegases and interactive land–atmosphere forcing are keptthe same in both simulations. To assess the contribu-tions of individual aerosol types and to delineate thenonlinear nature of the interaction between forcing andresponses (Kondratyev 1999), we carry out three addi-tional model simulations (which are otherwise identicalto NA and AA): one without black carbon (NB), onewithout dust (ND), and one with dust only (OD).

3. Global-scale response

a. Aerosol-induced radiation anomalies

Figure 2 shows MAM mean distribution of aerosoloptical thickness of all aerosols, dust, and black carbon(BC), respectively, as provided by the GOCART cli-matology. The optical thickness of all aerosols (Fig. 2a)shows two major maxima located over northern Africa

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and East Asia and two secondary maxima over Europeand northeastern India. The elongated spatial patternthat extends from China to Japan reflects aerosol trans-port by the western Pacific subtropical high, a persis-tent circulation in the Northern Hemisphere during bo-real spring and summer (Lau et al. 2000). Although theseasonal-mean aerosol optical thickness is relativelylow over North America, it can be influenced by aero-sols transported episodically from Asia and Africa.Dust from the African continent is transported by up-per-level tropical easterlies to Central America andnorthern South America along the southern flank of ahigh pressure system located in the Atlantic Ocean(Moulin et al. 1997; Kaufman et al. 2002). Even thoughdust lifted by wind and dust storms in arid regions is amajor source of global dust loading, anthropogenic dust(disturbed soils) may also contribute up to 10% to the

total atmospheric dust load (Tegen et al. 2004). Theprimary source of aerosols in central China is sulfatesfrom industrial pollution, while in northern China it isdust from the Taklamakan and Gobi Deserts. In south-ern China, aerosols are made up of carbonaceous par-ticles from biomass burning and industrial pollution, aswell as dust transported from deserts in northern China(Figs. 2b,c). Aerosol optical thickness (AOT) in north-eastern India/ Bay of Bengal is mainly due to blackcarbon and organic carbon from industrial sources (Fig.2c). In northern and northwestern India and Pakistan,aerosols are lofted into the troposphere by wind stormsover local deserts, as well as transported from Afghani-stan, the Middle East, and North Africa. Conversely, inthe European and East Asian region, most carbon-aceous particles are from industrial sources. The trendin global BC emission is quite nonlinear. In both theUnited States and China, BC emissions have increasedover the entire period from 1950 to 1990 (Tegen et al.2000). The spatial pattern of organic carbon (notshown) is similar to, but its optical thickness is twicethat of, black carbon.

Figure 3 shows the monthly evolution of the AOT ofdust, carbonaceous, and sulfate aerosols in Marchthrough May. Dust aerosols over northern Africa ap-pear to be most pronounced, with increasing concen-tration from March through May. Other regions thathave high dust concentration include the northern In-dian Ocean and northern China–Mongolia regions,where it contributes 60%–90% of the total AOT (Figs.3a–c). The most important source of carbonaceousaerosols is biomass burning over northeastern Indiaand southeastern Asia during March through April(Figs. 3d–f). The carbon aerosols are reduced substan-tially, as a result of wet scavenging following the onsetof the South China Sea monsoon in mid-May (Fig. 3f).Carbonaceous aerosols are also emitted from industrialregions of the Northern Hemisphere, which have asteady rate of consumption of fossil and domestic fuelsthrough the year (Takemura et al. 2000), but theamounts are much smaller compared to those in Southand Southeast Asia. A large concentration of sulfateaerosols is found in East Asia and, to a lesser extent, inEurope (Figs. 3g–i). Sulfate is mainly formed in theatmosphere from the oxidation of SO2, which is emittedpredominantly from fuel combustion and industrial ac-tivities (Chin et al. 2002).

The effectiveness of aerosol forcing in producing afar-field, global response is very much dependent onthe height of the aerosol layer. A deep heating layer,with vast (continental scale) horizontal extent, is likelyto excite barotropic Rossby wave energy dispersion ona global scale, with fast response time (Hoskins and

FIG. 2. Distribution of GOCART aerosol optical thickness inMAM of (a) all aerosols at wavelength of 0.55 �m, (b) dust, and(c) black carbon. AOT of black carbon is scaled by a factor of 10,regions over 0.4 are shaded, and contour intervals are 0.1.

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Karoly 1981; Lau and Lim 1984). Moreover, absorbingaerosols lofted above cloud layer will have an enhancedheating effect due to multiply reflected solar radiationfrom the underlying brighter clouds (Quijano et al.2000; Takemura et al. 2002a). Figure 4 shows the lon-gitudinal average (0°–150°E) latitude–height distribu-tion of AOT of all aerosols, dust, black carbon, andsulfate, respectively. Two major regions exhibiting highaerosol loading are found over latitude 25°–30°N andaround 40°N, associated primarily with dust over NorthAfrica and inland deserts, and sulfate and BC over Eu-rope and China (Fig. 4a). Even at a mid-to high-tropospheric level, the aerosol concentration remainsconsiderable owing to dynamic uplifting and transport.Figure 4b clearly shows that most aerosols at higherelevation are due to dust (mostly fine dust) that can betransported up and away from the source regions by thelarge-scale circulation. In contrast, BC is most pro-nounced in the lower troposphere and within theboundary layer (�900 hPa), resulting from biomassburning and air pollution. The vertical distribution ofsulfate aerosols (Fig. 4d) shows heavy concentrationconfined near or below the boundary layer, with a

maximum located at 30°N and two secondary maximaat 50°–55°N.

In the following, we define the aerosol-induced ra-diative flux anomaly (AIRA) as the difference in netradiation between the aerosol simulation (AA) and no-aerosol simulation (NA). The AIRA is not the same asthe conventional definition of aerosol radiative forcing,which is the difference between the radiative fluxes be-tween an ambient aerosol-laden atmosphere and that ofa clean atmosphere under the same atmospheric con-ditions. The AIRA includes not only the direct radia-tive forcing (both shortwave and longwave) by aerosolsbut also anomalies induced by changes in atmosphericstates through feedback in land surface temperature,atmospheric temperature and water vapor profiles, at-mospheric stability, and clouds. Figure 5 shows the dis-tribution of AIRA at the top of the atmosphere (TOA),within the atmosphere (top minus surface), and at thesurface, respectively. The magnitude of AIRA at theTOA is much smaller than that of the atmosphere orthe surface. This is because the negative forcing (cool-ing) due to aerosol scattering is compensated by aerosolabsorption (warming). As a result, the pattern of AIRA

FIG. 3. Spatial distribution of monthly AOT of (top) dust, (middle) carbonaceous, and (bottom) sulfate aerosolsat wavelength of 0.55 �m in (left) March, (center) April, and (right) May. Regions over 0.3 are shaded, and contourintervals are 0.1.

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of the atmosphere is similar to that at the surface butwith opposite sign, implying that the net effect of globalaerosol is to redistribute heat between the surface andthe interior of the atmosphere. Table 1 shows that theglobally averaged net AIRA (shortwave and longwavecombined) at the surface (�5.0 W m�2) is about threeand a half times larger than that at the top of the at-mosphere (�1.4 W m�2). The anomalous heating of theatmosphere is 3.6 W m�2. Notice also that the depletionin surface shortwave absorption (5.9 W m�2) in springis approximately balanced by a reduction in longwaveemission at the surface (0.9 W m�2) and a gain in sen-sible (1.8 W m�2) and latent heat flux (1.5 W m�2),resulting in a smaller net energy loss at the surface (1.7W m�2). Since globally evaporation must balance pre-cipitation, the energy balance implies that there is aglobal reduction in precipitation, equivalent to a loss ofatmospheric heat of condensation of 1.5 W m�2, as aresult of the adjustment of the atmospheric circulationand associated changes in cloudiness and water vapor.Roeckner et al. (1999) show that the annual-mean sur-face net radiation anomaly induced by the direct forc-

ing of anthropogenic sulfate is about �1.28 W m�2,calculated as the difference between the experimentwith greenhouse gas plus sulfate aerosols and the ex-periment only with greenhouse gases in their study.This anomaly was almost balanced by the reduction insensible (0.17 W m�2) and latent (0.95 W m�2) heatflux. Overall these values are smaller than our results(�4.8 W m�2 for net surface radiation, 1.5 W m�2 forsensible, and 1.6 W m�2 for latent heat flux) becauseour results are obtained from all aerosol effects includ-ing anthropogenic sulfate. Compared to ours, their re-sults show that when dust is excluded the role of latentheat flux in balancing the reduction of net radiation atthe surface is more important than that of sensible heatflux. Since our results include dust aerosol effects overdeserts, it is reasonable that sensible heat flux has amore important role compared to Roekner et al.(1999). It should also be noted that the period for com-parison is different.

The spatial distribution of AIRA indicates that thedistribution of radiative flux anomalies exhibits pat-terns similar to the total aerosol optical thickness, ex-

FIG. 4. Latitude–height distributions of aerosol optical thickness at wavelength of 0.55 �m averaged between 0°and 150°E for (a) all aerosols, (b) dust, (c) black carbon, and (d) sulfate. Contour intervals in (a), (b), (c), and (d)are 10, 5, 1, and 5 � 10�4 hPa�1, repspectively.

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cept over China where sulfate aerosols are abundant.The atmospheric heating is mainly due to the contribu-tion of black carbon and dust. Even though the opticalthickness of black carbon is almost 10 times smallerthan dust, the former plays a very important role in

absorbing solar energy in the atmosphere, especiallyover East and South Asia (see section 4 for more dis-cussions). As suggested by Jacobson (2001a), the ab-sorption power of solar radiation by black carbon maybe potentially larger than that of the model simulation.

FIG. 5. MAM aerosol-induced radiation anomaly (a) at the top of the atmosphere, (b) in theatmosphere, and (c) at the surface in W m�2, computed as the difference between the all-aerosol experiment and the control (AA minus NA). Contour intervals are 5 W m�2, andregions larger than 15 or smaller than �15 W m�2 are shaded.

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It is because black carbon can exist in one of severalpossible internally or externally mixing states that havedifferent and stronger absorption compared to purecarbon (Haywood et al. 1997; Jacobson 2001a).

Table 2 shows estimates of the local response in sur-face air temperature, and tropospheric column meantemperature anomalies associated with the AIRA,computed as the composite mean at grid points overland with AOT between the peak and its thresholdvalue for total aerosol, sulfate, BC, and dust, respec-tively. The threshold value is approximately equal to40% of the peak value, covering 5% of the land area,for each aerosol type. Because of the large concentra-tion of sulfate compared to BC and the large overlapbetween their source regions, the BC effect may bemasked by the sulfate effect. To minimize the sulfateinfluence, the estimates for BC are taken only over landgrid points where sulfate aerosol concentrations are low(less than its threshold value). The overall surface cool-ing by all aerosol types due to solar dimming is evident,with sulfate having the strongest cooling effect(�0.53°C) compared to BC (�0.42°C) and dust(�0.36°C). The major differences among the aerosolsare in their atmospheric heating effects. Sulfate has theleast atmospheric heating effect because it is nonab-sorbing. While sulfate does not heat the atmosphere, itstill shows weak atmospheric warming in Table 2 be-cause of partial overlap between sulfate source and BCsource over the region where AOT is larger than 0.15,especially in China. Thus atmospheric heating by BCabsorption can mask cooling by sulfate scattering overthis region, resulting in weak warming. Even though itsheating efficiency, that is, heating per unit AOT, ishighest, the atmospheric temperature anomaly due toBC is moderate (�0.11°C) because of its small-scaleheight and the smaller AOT (see Fig. 4c) compared todust and sulfate. The atmosphere over the heavy dustregions (mostly deserts) shows the strongest heatingeffect (�0.41°C), due to the combined effects of en-hanced absorption over a bright surface, vertical heattransport through turbulent heat flux from the surface,and dry convection.

b. Dynamical responses

Figure 6 shows the aerosol-induced (AA minus NA)temperature anomalies at 850 and 500 hPa during bo-real spring. The most prominent feature is a wave trainpattern in the lower and middle layer of the tropo-sphere emanating from the Sahara region, through Eu-rope and Siberia, into northeastern Asia. A large posi-tive anomaly can be seen over northern Africa and theMediterranean, accompanied by a negative anomalyover Europe. Over western and northern Asia, eventhough aerosol loading is relatively low, the magnitudeof temperature changes are larger than those found atthe source regions, for example, central China, whichshows strong low-tropospheric cooling caused by sul-fate aerosols. At 500 hPa, the pattern (Fig. 6b) shows apronounced elongated “warm bridge” extending from

TABLE 1. Global mean aerosol-induced radiation anomaly dur-ing boreal spring. Positive (negative) values indicate the gain(loss) of energy (W m�2), respectively.

Energy flux

Shortwave LongwaveSensible

heatLatentheat Net

TOA �1.1 �0.3 — — �1.4Atmosphere 4.8 �1.2 �1.8 �1.5 0.3Surface �5.9 0.9 1.8 1.5 �1.7

TABLE 2. Estimates of column-mean tropospheric temperature(Ttrop), and surface temperature (Tsfc) anomalies (°C) induced bylocal effects of aerosols for all aerosols, sulfate, black carbon, anddust, respectively. Temperature anomalies are computed overland grid points with AOT greater than its threshold value (seetext).

All aerosols(AOT�0.25)

Sulfate(AOT�0.15)

Black carbon(AOT�0.03)

Dust(AOT�0.23)

Ttrop 0.28 0.09 0.11 0.41Tsfc �0.39 �0.53 �0.42 �0.36

FIG. 6. MAM temperature anomaly induced by all aerosols at(a) 850 hPa and (b) 500 hPa. Significance levels of 5% are shadedand contour intervals are 0.3°C.

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central Asia to North America coupled with weak cool-ing to the south. While portions of the warm bridgeover North Africa, South Asia, and East Asia lie withinthe aerosol source regions, a large signal-to-noise ratio(greater than 5% confidence level shown as shaded)appears mainly outside these regions. The wave trainresembles the steady-state response of the atmosphereto thermal and orographic forcing via Rossby wave en-ergy dispersion (Hoskins and Karoly 1981). The aboveresults suggest that the wave train pattern is induced bythe changes in circulation in response to aerosol forcingrather than by direct aerosol radiative forcing alone.

The large-scale circulation changes induced by aero-sol direct forcing can be seen clearly in the sea levelpressure (SLP) and 500-hPa geopotential heightanomalies, as shown in Fig. 7. A large positive SLPanomaly is found over the East Asian monsoon region,centered around the Korean peninsula, while a nega-tive anomaly is found around Tibetan Plateau, formingan east–west dipole pattern. The latter is mainly due tosurface warming induced by surface albedo–aerosol in-teraction associated with aerosols absorbing multiplyscattered solar radiation over the high surface albedo ofthe Tibetan Plateau (Bakan et al. 1991; Quijano et al.2000; Takemura et al. 2002a). The geostrophic south-westerlies associated with the SLP anomaly over theNorth Pacific will cause warm advection toward thecentral and northern part of the Asian continent, whichmay be responsible for the surface warming shown in

Fig. 6. This pattern also implies a strengthening of thewestern Pacific subtropical high, which governs the on-set and evolution of the East Asian summer monsoon.Large SLP anomalies are also found in the African–European region with the positive SLP north of theMediterranean Sea and negative SLP over North Af-rica. The previously described cooling around the Cas-pian Sea (Fig. 6) may be explained in terms of thesouthward cold advection induced by anticyclonic flowaround the European high. The circulation change inthe midtroposphere (Fig. 7b) is in agreement with thetemperature change shown in Fig. 6b. In the midtropo-sphere, the cold trough appears to be more significantlyaround the Caspian Sea than in the lower troposphere.It is conceivable that the aforementioned teleconnec-tion may also be affected by tropical and extratropicalsea surface temperature anomalies induced by aerosols(i.e., Rotstayn and Lohmann 2002), as well as El Niñoeffects. Effects of coupled air–sea interaction on atmo-spheric teleconnections will be reported in a separatepaper.

4. Regional responses

In this section, we discuss the regional circulationassociated with the aforementioned teleconnection in-duced by aerosol forcing, focusing on the East Asianand the North Africa/Europe sectors.

a. East Asian region (100°–120°E)

In boreal spring, the climatological convergence zonein East Asia is located around latitude 30°N (Fig. 8a),coinciding with a region of high aerosol concentration,primarily sulfate but with a significant contributionfrom dust and BC (Fig. 8c). The convergence zone isassociated with weak-to-moderate premonsoon frontalsystems that prevail over southern and central EastAsia in boreal spring. In the southern part of this re-gion, carbonaceous aerosols are abundant while, in thenorthern part, dust aerosols dominate (Fig. 8c). Thedirect surface cooling due to aerosols is obvious in thelarge negative surface temperature anomaly in the re-gion 20°–40°N. In spite of the complex aerosol loadingpattern in this region, the circulation change by themeridional wind (Fig. 8b) shows clearly a reversed pat-tern associated with the suppression of convection inthe mean convergence zone, compared to the control(Fig. 8a). The suppression is due to the increase of staticstability associated with warming of the lower tropo-sphere and cooling of the surface due to aerosols (Fig.8c), that is, the semidirect effect (Hansen et al. 1997a).The reduced latent heating, associated with suppressed

FIG. 7. Spatial distribution of (a) sea level pressure and (b)geopotential height anomaly at 500 hPa induced by all aerosols.Contour intervals in (a) and (b) are 0.5 hPa and 8 m.

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convection, coupled with surface cooling and turbulentmixing within the boundary layer produce a negativetemperature anomaly that extends from the surface tothe entire troposphere (Fig. 9a). The suppressed con-vection, in addition, induces a complex local circula-tion, as shown in the anomalous vertical motion field(Fig. 9d). In the convergence zone around 30°N wheresulfate aerosol loading is significantly large, anomalousdescent is induced, that is, the climatological upwardmotion is weakened. However, ascending motion is in-duced north and south of this zone where dust and BCconcentrations, respectively, are relatively high. Asshown in Fig. 9b, the shortwave heating due to aerosolsis slightly larger in the outer part of the convergencezone than within. Hence, this heating will provide apositive feedback due to the reinforced rising motions

outside the convergence zone. As shown in Fig. 9a,large atmospheric warming is found in the latitude belt40°–60°N, where shortwave heating by aerosol is rela-tively weak. Figure 9c shows that diabatic cooling stem-ming from reduced condensation heating is dominant inthis region. However, at 40°–60°N, adiabatic warmingdue to downward motion is more important in main-taining the large tropospheric warming there. Overall,it is clear that the response is complex and far from asimple pattern linked to local direct aerosol forcing inthe source regions. Rather, the structure and magnitudeof the regional response is determined by the balancebetween shortwave heating and induced condensationheating/cooling, as well as adiabatic effects, and turbu-lent diffusion in the boundary layer.

b. Africa–Europe region (0°–30°E)

Similar to the East Asian region, the initial aerosoldirect forcing induces a complex regional response in-volving feedback of the atmospheric water and energycycles in the North Africa–Europe sector. Here, theclimatological meridional convergence zone is locatedat the southern boundary of the desert region of NorthAfrica, near 10°N (Fig. 10a), as an extension of theAtlantic intertropical convergence zone (ITCZ) overthe African continent. Figure 10c shows that dust aero-sol dominates North Africa (5°–35°N), while sulfate isrelatively abundant over Europe (40°–60°N). The BCconcentration is generally low, with a slight maximumnear 5° and 50°N. The convergence is weakened byaerosol effects, as seen in the induced anomalous me-ridional circulation (Fig. 10b) and the surface coolingdue to aerosol direct effect near 10°N (Fig. 10c). Thesurface cooling extends into the lower troposphere upto 750 hPa (Fig. 11a), but the shortwave heating by dustextends up to 500 hPa, thus giving rise to a positivetemperature anomaly above 700 hPa. The semidirecteffect is evident in the differential vertical heating,which increases the vertical stability of the atmosphere,leading to reduced convection and anomalous sinkingmotion (Fig. 11d). Over the desert region (20°–30°N),slight surface warming is found (Fig. 10c). This is be-cause the absorbing aerosol over the high-albedo desertsurface can produce multiply reflected solar radiationbetween the dust layer and the surface, producingwarming that offsets most of surface cooling induced byaerosol scattering and absorption in the atmosphere.As a result, anomalous upward motion is induced byshortwave heating by the abundant dust aerosol abovethe boundary layer, reinforced by positive feedbackfrom the anomalous meridional circulation. Conse-quently, the latitudinal distribution of surface air tem-

FIG. 8. Latitude–height distributions of meridional wind in (a)NA and (b) the meridional wind anomaly (AA minus NA) aver-aged over the 100°–120°E longitude sector. Contour intervals are1.0 for (a) and 0.3 m s�1 for (b), respectively. Latitude distributionof surface air temperature anomaly (solid line) and AOT of sul-fate (dotted line), dust (long dashed line), and carbonaceous aero-sol (short dashed line) are indicated in (c).

15 SEPTEMBER 2006 K I M E T A L . 4709

perature does not correspond to that of the AOT dis-tribution (Fig. 10c).

Shortwave heating by aerosols is large over the Sa-haran region (10°–30°N) and small over the Europeanregion, 40°–60°N (Fig. 11b). The depth of the aerosolloading in the Saharan and European regions reaches atleast to the middle troposphere, due mainly to dustaerosols that are lifted above the planetary boundarylayer (PBL) reaching a height of 800 to 1500 m in theSaharan region (�850-hPa level). On the other hand, inthe Mediterranean belt, 30°–40°N, shortwave heating isvery weak because this region is not located in the maintrack of aerosol dust transport in the spring. The tem-perature anomaly pattern shown in Fig. 11a does notcorrespond directly to shortwave radiative heating butreflects the role of induced diabatic heating (Fig. 11c)associated with the dynamic responses. As shown inFig. 11c, diabatic cooling is almost �0.5 K day�1, sig-nificant in the lower troposphere at 10°N where a nega-tive temperature anomaly is found. This is becauseaerosols reduce the net solar flux reaching the surface

and, consequently, reduce the PBL height. In this wayturbulent mixing causes a reduced PBL height andnegative temperature anomaly in the lower tropo-sphere. When absorbing aerosols such as smoke andmineral dust are lifted over the PBL by penetrativeconvection and long-range transport, they can increasethe strength of capping inversion, which in turn coolsPBL and reduces the PBL height. Our analysis (notshown) indicates that the maximum cooling rate by ver-tical diffusion is about �1.0 K day�1 in this region.

As shown in Fig. 11d, the induced meridional circu-lation features enhanced upward motion coupled withlow SLP (see Fig. 7a) and increased tropospheric tem-perature (Fig. 11a) in the northern part of the Sahara(20°–30°N). In the upper troposphere between 30° and40°N, air mass is transported northward (Fig. 10b). Asthe air mass goes northward, it cools and sinks at thelatitude of 50°N (Fig. 11d), consistent with the devel-opment of the SLP high over Europe (see Fig. 7a). Theresults indicate that the anomalous meridional over-turning associated with interaction of atmospheric heat-

FIG. 9. Latitude–height distribution of (a) temperature, (b) shortwave heating, (c) total diabatic heating, and (d)vertical pressure velocity anomaly induced by all aerosols averaged over the East Asian domain [100°�120°E].Contour intervals for (a), (b), (c), and (d) are 0.2°C, 0.1°C day�1, 0.1°C day�1, and 0.3 � 10�4 hPa s�1, respectively.In (d) a negative anomaly denotes downward motion.

4710 J O U R N A L O F C L I M A T E VOLUME 19

ing and surface cooling near the aerosol source region iscritical in setting up the downstream teleconnectionpattern outside the region of aerosol forcing.

5. Relative roles of dust versus black carbon

In this section, we present results from additionalexperiments to demonstrate the relative importance ofthe two dominant absorbing aerosols, that is, dust andblack carbon, in contributing to the boreal spring tele-connection. Figure 12 shows the MAM atmosphericgain in radiative flux (top minus bottom), based on anexperiment where the forcing from dust aerosol is with-held (ND), and another one where BC forcing is with-held (NB). In ND, almost all the atmospheric radiativeheat gain is concentrated over the South Asia andSoutheast Asia monsoon region, and to a smaller extentin northeastern Europe where BC is dominant, asshown in Fig. 2c. On the other hand, in NB, the atmo-spheric radiative heat gain is confined over northernAfrica and the Middle East where the subtropicaldeserts are located.

To show the relative impacts of dust and BC, we haveconstructed the difference maps of AA minus ND (fordust) and AA minus NB (for BC) and compared themto the full aerosol impact (AA minus NA). The tacitassumption here is that the present-day climate is asso-ciated with AA, and we examine the impacts of dust,and BC, by asking the question: what if dust or BC weremissing?

Figure 13 shows the surface air temperature anoma-lies for the three cases. The statistical confidence here iscomputed as a two-tailed Student’s t test (Wilks 1995),based on the 10 independent samples, that is, 10 sea-sons, subject to the same aerosol forcings. With all-aerosol forcing (Fig. 13a), the most pronounced (ex-ceeding 1% confidence level) surface cooling is foundin a broad region over the Caspian Sea, which is not aregion of high aerosol loading. Hence the strong cool-ing here is not due to direct radiative effect of aerosols,but is a response to the large-scale circulation (see dis-cussion in section 3b). In contrast, the negative anoma-lies found over the aerosol source region of the Bay ofBengal (BC), East Asia (BC and sulfate), and westernAfrica (dust) are consistent with the direct forcing byaerosols. The strongest surface warming is found overan extensive region in northeastern Asia, far from theaerosol source regions and opposite in sign to that ex-pected from aerosol forcing. The surface warming overnorthern India is apparently due to the BC and dustaerosol, accumulating against the southern slope of theHimalayas, acting as an elevated heat source. Here thesurface rapidly rises to 3–5 km above sea level. Theelevated aerosol heat source leads to anomalous warm-ing of the upper troposphere in late spring and conse-quently causes enhanced increased rainfall over north-ern India and the Bay of Bengal during the followingsummer. The impacts of aerosols on the South Asianmonsoon is a subject of a separate paper (Lau et al.2006).

Figure 13a shows a broad region of surface warmingover southern Europe that, as shown below, is found tobe a part of a wave train signal forced by dust aerosolover northern Africa. From Fig. 13b, it is obvious thatthe pronounced cooling over eastern Europe and theCaspian Sea region is associated with an atmosphericteleconnection forced by dust aerosol heating overnorthern and western Africa. This wave train produceswarming over southern Europe and northern Siberia.As shown in Fig. 13c, BC produces surface cooling overeastern India and East Asia and, in doing so, induces ameridional circulation over the East Asian sector (seediscussion in section 4a) with strong subsidence innorthern latitudes (40°–60°N). The most pronounced

FIG. 10. As in Fig. 8 but induced by all aerosols at 0°–30°E.

15 SEPTEMBER 2006 K I M E T A L . 4711

effect is to enhance the surface warming over centraland eastern Siberia.

The signature of the atmospheric teleconnection in-duced by dust and BC from the source region is alsoevident in the 500-hPa geopotential height anomaliesfor the NA, ND, and NB experiments shown in Fig. 14.From Fig. 14b, it is clear that dust over northern Africaand the Middle East induces a transcontinental wavetrain spanning northwestern Europe across Siberia tonortheastern East Asia. This wave train appears to re-produce the polar component (north of 40°N) of thewave train due to radiative forcing by all aerosols (Fig.13a), suggesting the primary importance of dust aerosolin forcing the teleconnection. The negative anomalyassociated with tropospheric cooling north of the Cas-pian Sea is very prominent. In contrast, the contribu-tion of BC forcing to the atmospheric teleconnection ismanifested in increasing the magnitude and area oflarge positive height signal over eastern Siberia (Fig.14c). Comparing Figs. 13 and 14, and 200-hPa anoma-lies (not shown), the wave train signal is quasi-barotropic, with a vertical structure that extends fromthe surface to the upper troposphere.

Because of nonlinear interactions among the changesin diabatic heating and circulation, the sum of the in-fluences of all aerosol types is not equal to the effects ofthe sum of the individual contributions (Kondratyev1999). These results reaffirm that the two key absorbingaerosols (dust and BC) can contribute to regional cli-mate anomalies not only through direct radiative forc-ing but also through interaction of diabatic heating (ra-diative and latent heat of condensation) associated withthe anomalous circulation.

6. Discussion and conclusions

Experiments with the NASA fvGCM show that glob-al aerosol radiative forcing in boreal spring can inducean atmospheric teleconnection pattern with pro-nounced signatures in sea level pressure, surface tem-perature, and geopotential height that span northernAfrica, southern Europe, Siberia, and northeasternAsia. Surface cooling is found in the vicinity and down-wind of source regions over South and East Asia, WestAfrica, and Europe, consistent with the solar dimmingeffect of aerosols. However, the most pronounced sur-

FIG. 11. As in Fig. 9 but induced by all aerosols at the 0°–30°E longitude band. In (d), negative values denoteanomalous sinking motion.

4712 J O U R N A L O F C L I M A T E VOLUME 19

face temperature feature is a dipolelike anomaly withcooling found near the Caspian Sea and warming northof Lake Baikal. The dipole anomalies are located overregions where aerosol forcing is weak and arise from aremote dynamical response induced by aerosol forcingenhanced by local circulation feedback.

In the African–European region, a low pressureanomaly develops as a result of direct atmosphericheating by Saharan dust in the lower and middle tro-posphere. The heating enhances upward motion in theSaharan region and drives an anomalous meridionalcell with downward motion, high pressure, and surfacewarming over southern Europe. Through barotropicdispersion (Hoskins and Karoly 1981), atmospheric en-ergy propagates away from the source region eastwardover Europe, across Siberia, to the North Pacific. Thecooling at the Caspian Sea is not due to direct aerosolheating but, rather, is associated with the southwardadvection of cold air from high latitudes on the easternside of the surface pressure anomaly over Europe.Similarly, over the East Asian region, BC aerosol in-duces a multicellular meridional overturning with sink-ing motion over central and northeastern Asia (50°–70°N), which leads to adiabatic warming. The warmingis enhanced by advection from the western side of theanomalous West Pacific high, which is a part of theatmospheric teleconnection pattern.

We find that the springtime atmospheric teleconnec-tion is primarily caused by dust aerosol over North Af-rica with reinforcement due to BC aerosol forcing overSouth and East Asia. The result also implies that strongatmospheric heating by absorbing aerosols can redis-tribute energy over broad regions far away from thesource regions. The springtime atmospheric telecon-nection pattern is excited due to the interaction of at-mospheric dynamics and diabatic heating processes inthe atmosphere, driven by atmospheric heat sourcesassociated with dust and BC aerosol loading.

As already discussed in the main text, we need tocaution that AIRA should not be literally comparedwith previous estimates of aerosol radiative forcing be-cause the quantity depends on column properties thatare modified by the forcing as well as the responses inour experiments. As a qualitative comparison, our es-timate of the annual-mean surface global AIRA is �4.8W m�2 (not shown), consistent with that (� �4.0 Wm�2) of Jacobson (2001b). Ramanathan et al. (2001)estimated the change in aerosol radiative forcing at thesurface from the preindustrial period to the present tobe �2 to �4 (W m�2). Our estimate of AIRA is some-what larger than this range. However, if we considerthat, even in the preindustrial period, there will bebackground natural aerosols present, our AIRA esti-mate clearly would have overestimated the aerosolforcing. Ramanathan et al. also found that the annual-mean aerosol forcings by indirect and greenhouse ef-fects are �1 � 0.5 and 1 W m�2, respectively, suggest-ing that the indirect effect may cancel out the green-house effect. This means that the total effect (direct �indirect � greenhouse effects) would have the samerange as the direct effect. This notion is supported byrecent results (e.g., Liepert et al. 2004) showing that theobserved reduction in shortwave radiation and surfacetemperature increase over land during the last four de-cades is reasonably well simulated by considering thecombined effects of aerosol direct, indirect, and green-house gas increases. We also note that the generallylarger surface AIRA estimate compared with the afore-mentioned estimates may be due to the inclusion ofatmospheric dynamical feedbacks, which tend to am-plify the aerosol radiative forcing.

Because this is a first numerical attempt to elucidatea mechanism of atmospheric teleconnections inducedby aerosol forcing, there are obvious limitations andcaveats. First, aerosol indirect effects, two-way coupledaerosol–atmosphere dynamics, and interactive sea sur-face temperature effects have not been included in thisstudy. These effects may further modulate the structureof the teleconnection signal and need to be pursued infuture studies. Until now, studies of aerosol effects have

FIG. 12. MAM aerosol-induced radiation anomaly in the atmo-sphere by (a) all aerosols without dust (ND minus NA) and (b) allaerosols without BC (NB minus NA). Regions over 10 W m�2 areshaded, and contour intervals are 5 W m�2.

15 SEPTEMBER 2006 K I M E T A L . 4713

mostly been focused on estimates of direct and indirectforcing and local responses, but not on the impacts onatmospheric dynamics, as in the present study. Clearly,further sensitivity experiments to aerosol variabilityneed to be carried out to determine the signal-to-noiseratio with respect to the model as well as to observa-tions.

Second, even with aerosol direct forcing, there areuncertainties regarding the radiative properties of aero-

sols, particularly with dust, which we have found to bemost important as a natural source of aerosol thatspawns the atmospheric teleconnection. The SSA val-ues of 0.92–0.9 at 550 nm used in our experiments aresomewhat lower, implying stronger absorption thanvalues (�0.97–0.95) estimated by Kaufman et al. (2001)for Saharan dust. However, we note that absorptionproperties of airborne dust in different parts of theworld are likely to vary substantially, depending on the

FIG. 13. MAM mean surface air temperature anomaly induced by (a) the effects of allaerosols (AA minus NA), (b) dust aerosol with the background of four aerosols (AA minusND), and (c) BC aerosol with the background of four aerosols (AA minus NB). Contourintervals are 0.5°C. Significance levels are shown in the half-tone scale on the right.

4714 J O U R N A L O F C L I M A T E VOLUME 19

physical and chemical properties of the dust. For ex-ample, fine dust over Asia was found to be stronglyabsorbing, with SSA � 0.88 (Anderson et al. 2003).Moreover, in Asia, and in some regions of Africa, air-borne dust particles are likely to be externally mixedwith carbonaceous aerosols from industrial pollutionand biomass burning. Such a mixture is likely to havehigher absorption (lower SSA) than pristine dust fromthe Sahara Desert. For example, D.-H. Kim et al.(2004) found that SSA is approximately 0.8 at Anmyonand Gosan, South Korea, and Amami-Shima, Japan,

downstream of the dust source in western China, muchsmaller than the value of around 0.9 found in Duhuangand Yinchuan, China, during major dust events in bo-real spring, indicating strong mixing of dust and pol-luted air masses over Asia. Our results call for bettermeasurements of aerosol optical properties, especiallydust, globally. An important result, which our studyemphasizes, is that it is not so much the exact magni-tude of the SSA but rather the multiple reflections/absorption between the aerosol and underlying brightsurface, together with the vast horizontal extent and the

FIG. 14. As in Fig. 13 but for the MAM temperature anomaly at 500 hPa.

15 SEPTEMBER 2006 K I M E T A L . 4715

high elevation of the dust lifted by the large-scale cir-culation, that contribute to the efficacy of dust as aprimary energy source for atmospheric teleconnections.To better understand aerosol–atmospheric dynamicsinteraction, besides anthropogenic aerosol sources, cli-mate models should incorporate dust effects as a natu-ral heat source/sink in the atmosphere, and observa-tional emphasis should be on the global distribution,both vertical and horizontal, and differentiation ofaerosol types over different surfaces.

Our results should also provide useful clues for ex-tracting signals of climate response to aerosol forcingfrom observations. In the real world, the atmosphericteleconnection, described here, may be interpreted as asignal of interannual climate anomalies generated byepisodes of strong dust storms over North Africa andthe Middle East region during boreal spring. It is alsointeresting to note that the surface air temperatureanomalies induced by global aerosols (Fig. 13a) is quitesimilar to the long-term trend pattern shown in Fig. 1,except for the West Africa sector, which has oppositesign. This may lead to the hypothesis that direct aerosolforcing, perhaps from an increasing BC emission trendin the last several decades, may have played a role inthe observed long-term temperature dipole over Eur-asia. Similarly, the 500-hPa and 850-hPa temperatureanomaly patterns shown in Figs. 14a,b bear strong re-semblance to the teleconnections associated withanomalous storm track activity over Eurasia and a sum-mertime precipitation seesaw between western andeastern Siberia (Fukutomi et al. 2004), suggesting thatthe summertime rainfall seesaw may have been precon-ditioned by enhanced atmospheric-heating dust aero-sols over northern Africa and the Middle East. Clearlyfurther observation and modeling studies are needed tovalidate these hypotheses.

Acknowledgments. This work is supported by theNASA Modeling and Analysis Program. Most of thework was carried out while M.-K. Kim was visiting theLaboratory for Atmospheres, on a joint GSFC andGoddard Earth Sciences and Technology Center(GEST) fellowship. This work was partially supportedby the Climate Environmental System Research Center(CES) sponsored by the Korea Science and Engineer-ing Foundation.

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