Radiative effects of natural aerosols: A review

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Atmospheric Environment 39 (2005) 2089–2110 www.elsevier.com/locate/atmosenv

Radiative effects of natural aerosols: A review S.K. Satheesha,, K. Krishna Moorthyb a Centre for Atmospheric and Oceanic Sciences, Indian Institute of Science, Bangalore-560 012, India Space Physics Laboratory, Indian Space Research Organisation (ISRO), Vikram Sarabhai Space Centre, Trivandrum-695 022, India

b

Received 22 March 2004; received in revised form 30 September 2004; accepted 9 December 2004

Abstract In recent years, there has been a substantial increase in interest in the influence of anthropogenic aerosols on climate through both direct and indirect effects. Several extensive investigations and coordinated field campaigns have been carried out to assess the impact of anthropogenic aerosols on climate. However, there are far fewer studies on natural aerosols than on anthropogenic aerosols, despite their importance. Natural aerosols are particularly important because they provide a kind of base level to aerosol impact, and there is no effective control on them, unlike their anthropogenic counterparts. Besides, on a global scale the abundance of natural aerosols is several times greater than that of the major anthropogenic aerosols (sulphate, soot and organics). The major natural aerosol components are sea salt, soil dust, natural sulphates, volcanic aerosols, and those generated by natural forest fires. As with anthropogenic aerosols, the abundance of natural aerosols such as soil dust is also increasing, due to processes such as deforestation, which exposes more land areas which may then interact directly with the atmosphere, and due to other human activities. Since a major fraction of the natural aerosol (sea salt and natural sulphate) is of the non-absorbing type (and hygroscopic), it partly offsets the warming due to greenhouse gases as well as that due to absorbing aerosols (e.g., soot). The mineral dust transported over land and ocean causes surface cooling (due to scattering and absorption) simultaneously with lower atmospheric heating (due to absorption); this could in turn intensify a low-level inversion and increase atmospheric stability and reduce convection. To accurately predict the impact of dust aerosols on climate, the spatial and temporal distribution of dust is essential. The regional characteristics of dust source function are poorly understood due to the lack of an adequate database. The reduction of solar radiation at the surface would lead to a reduction in the sensible heat flux and all these will lead to perturbations in the regional and global climate. Enhanced concentration of sea salt aerosols at high wind speed would lead to more condensation nuclei, increase in the cloud droplet concentration and hence cloud albedo. Even though direct radiative impacts due to sea salt and natural sulphate are small compared to those due to anthropogenic counterparts, their indirect effects (and the uncertainties) are much larger. There is a considerable uncertainty in sea salt aerosol radiative forcing due to an inadequate database over oceans. The presence of natural aerosols may influence the radiative impact of anthropogenic aerosols, and it is difficult to separate the natural and anthropogenic aerosol contributions to radiative forcing when they are in a mixed state. Hence it is necessary to document the radiative effects of natural aerosols, especially in the tropics where the natural sources are strong. This is the subject matter of this review. r 2005 Elsevier Ltd. All rights reserved. Keywords: Aerosols; Climate change; Radiative forcing; Radiation budget

Corresponding author. Tel.: +91 80 22933070/22932505; fax: +91 80 23600865.

E-mail address: [email protected] (S.K. Satheesh). 1352-2310/$ - see front matter r 2005 Elsevier Ltd. All rights reserved. doi:10.1016/j.atmosenv.2004.12.029

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1. Introduction The global climate system is a consequence of interactions between its sub-components (Shaw, 1983; Murthy, 1988; Charlson et al., 1992; Andreae, 1995; Hansen et al., 1997, 1998; Clarke, 1998; Haywood et al., 1999; Prospero et al., 2002; Seinfeld et al., 2004). The main processes that determine the overall state of the climate system are heating by incoming solar radiation and cooling by outgoing long-wave (infrared) terrestrial radiation (Coakley et al., 1983; Ramanathan et al., 1989; Charlson et al., 1991; Kiehl and Briegleb, 1993; Hansen et al., 1998; Seinfeld et al., 2004). Any process that can disturb the overall energy balance can cause climate change or perturbation (Kaufman and Fraser, 1997; Kaufman et al., 1997; Seinfeld and Pandis, 1998). A process that alters the radiative balance of the climate system is known as radiative forcing (Coakley et al., 1983; Coakley and Cess, 1985; Ramanathan et al., 1989; Charlson et al., 1991, 1992; Hansen et al., 1997, 1998; Russell et al., 1999; Bates, 1999; Raes et al., 2000). Radiative forcing can be internal or external. External forcing operates from outside the Earth’s climate system and includes orbital variations and changes in incident solar flux. Volcanic activity is an example of an internal forcing mechanism (Hoffmann et al., 1987; Moorthy et al., 1996; Soden et al., 2002). Similarly, changes in the composition of the atmosphere constitute another major internal forcing mechanism, and the best examples are the greenhouse gases and aerosols (Shaw, 1983; Crutzen and Andreae, 1990; Charlson et al., 1992; Clarke, 1993; Kaufman et al., 1997; Bates, 1999; Bates et al., 2000; Rodhe, 2000; Prospero et al., 2002). Changes in the greenhouse gas or aerosol content of the atmosphere affects the radiative balance of the climate system (Haywood and Ramaswamy, 1998; Myhre et al., 1998; Haywood et al., 2003). The Earth’s climate is strongly influenced by the manner in which solar radiation is absorbed and reflected in the atmosphere (Chylek and Wong, 1995; Schwartz et al., 1995). During the past 100 years the amount of carbon dioxide in the atmosphere has increased by about 25% on account of human activities (fossil fuel/biomass burning) (Meehl et al., 1996; Le Treut et al., 1998; IPCC, 2001). This has caused the surface temperature of the Earth to increase globally by about one kelvin (Meehl et al., 1996; Le Treut et al., 1998). In recent years, there has been a substantial increase in interest in the influence of anthropogenic aerosols on the climate through both direct and indirect radiative effects. Several extensive investigations and coordinated field campaigns have been carried out to assess the impact of anthropogenic aerosols on climate. However, studies of natural aerosols are few compared to those of anthropogenic aerosols, despite the importance of the

former. Among these studies, Aerosol characterization experiment-1 (ACE-1) focussed on natural aerosols. Aerosol characterization experiments (ACE) were designed to increase the understanding of how atmospheric aerosol particles affect the Earth’s climate system (Bates, 1999; Russell and Heintzenberg, 2000; Seinfeld et al., 2004). ACE-1 was conducted over southern hemispheric mid-latitudes with a specific goal of understanding the properties and controlling factors of aerosols in the remote marine atmosphere that are relevant to radiation balance and climate (Bates, 1999; Hainsworth et al., 1998; Griffiths et al., 1999). This environment provided an opportunity to establish the chemical, physical and radiative properties of a natural aerosol system. ACE-2 was conducted during July 1997 to study the radiative effects of anthropogenic aerosols from Europe and desert dust from Africa as they are transported over the North Atlantic Ocean (Russell and Heintzenberg, 2000). While ACE-1 and ACE-2 focussed mostly on natural and anthropogenic aerosols, respectively, ACE-Asia focussed on a complex mix of anthropogenic and natural aerosols over the Asian region (Huebert et al., 2003; Seinfeld et al., 2004). In this paper, we review the role of natural aerosols in modifying the Earth’s radiation budget and demonstrate its importance in the climate change debate. Throughout this paper we use the term ‘aerosols’ to address to the particulate phase of the atmospheric aerosol system.

2. Earth’s radiation balance: role of aerosols The Sun’s radiation, much of it in the visible region of the spectrum, warms our planet. On average, the Earth must radiate back to space the same amount of energy that it gets from the Sun (Seinfeld et al., 2004). Greenhouse gases (GHGs) in the Earth’s atmosphere, while largely transparent to incoming solar radiation, absorb most of the infrared (IR) radiation emitted by the Earth’s surface. Clouds also absorb in the IR. Thus, part of the IR emitted by the surface gets trapped (and this is the natural greenhouse effect). Under a clear sky, about 60–70% of the natural greenhouse effect is due to atmospheric water vapour (Seinfeld and Pandis, 1998). The next most important GHG is carbon dioxide, followed by methane, ozone, and nitrous oxide. If we represent solar radiation incident at the top of the atmosphere (global) as 100 units, then a net amount of 51 units reaches the surface (Fig. 1). Of the remaining 49 units, 3 units are absorbed by clouds and 16 units by aerosols, water vapour and CO2 together. The clouds, surface and atmosphere (which include aerosols as well) reflect 17 units, 6 units and 7 units, respectively. Of the 51 units absorbed by the Earth’s surface, 23 units are released as latent heat, 7 units as sensible heat and 21 units as infrared. About 15 units of infrared are

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Fig. 1. Earth’s radiation budget demonstrating the role of aerosols.

absorbed by aerosols, water vapour and CO2. Thus, aerosols have an important influence on the Earth’s radiation balance. It is widely known that warming, which tends to enhance evaporation, will increase the water vapour content of the troposphere (Seinfeld et al., 2004). This further amplifies the warming, as water vapour is a dominant GHG. Snow and ice reflect much of the incident sunlight back to space; thus a reduction of snow and ice cover would also lead to enhanced warming. Clouds are generally good absorbers of infrared, but high clouds have cooler tops than low clouds, so they emit less infrared spaceward. The interplay between atmosphere (GHGs, molecular absorbers and aerosols), ocean, clouds, and ice is poorly understood. The mean temperature of the Earth (Te) is given by the balance between the absorbed solar energy and emitted terrestrial energy given by the steady state condition, H¼

S0 ð1  aÞ  sT 4e ¼ 0, 4

(1)

where H is the net energy input to the climate system and S0 is the solar power per unit area intercepted at the mean Sun–Earth distance (solar constant) (1365–1372 W m2) (Seinfeld and Pandis, 1998). The factor 4 is the ratio of the Earth’s surface area to the cross-sectional area. The quantity a is the albedo (reflectance) of the Earth, which is the fraction of the incident solar radiation reflected by the Earth’s surface and atmosphere and has a mean value of 0.3. Consequently, of the 343 W m2 of the mean solar radiation incident at the top of the atmosphere, 103 W m2 is reflected back to space by the Earth’s surface and atmosphere. Aerosols can influence the albedo and thus can have an impact on the climate

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system. The energy balance equation implies that a change of 0.01 in the value of a results in about 1% change in the global temperature. The question of whether aerosols increase or decrease the value of a (warm or cool the planet) depends on their chemical composition. Completely scattering aerosols will increase a (which means a decrease in temperature) whereas absorbing aerosols (e.g., soot) would lead to a decrease in a (Coakley and Cess, 1985; Hansen et al., 1998; Russell et al., 1999) and hence an increase in Te. This means the warming or cooling effect can change from region to region depending on many factors such as the relative strengths of various aerosol sources and sinks (Kaufman et al., 1997; Clarke, 1998; Russell et al., 1999; Ginoux et al., 2001, 2004; Ramanathan et al., 2001; Luo et al., 2003). When the net effect of aerosols is cooling, they partly offset the greenhouse warming, while if the net effect is warming, they complement the greenhouse warming (IPCC, 2001). Since aerosol properties show large regional variations, the regional impact can be very different, and this is the main reason why the importance of aerosols is poorly characterised in climate models. This is especially true for natural aerosols, because of the lack of a comprehensive database.

3. Radiative effects of natural aerosols It is well known that aerosols are of natural or anthropogenic origin. The source strengths of various natural and anthropogenic species are given in Table 1 (data from Andreae, 1995). It can be seen that, in terms of emission, natural aerosols contribute 89%. In terms of column mass and optical depth, natural aerosols contribute 81 and 52%, respectively. Thus there exists no direct relationship between aerosol mass, optical depth and its radiative impact. Out of the major natural and anthropogenic aerosol types (sulphate, nitrate, sea salt, carbonaceous matter (organic carbon and black carbon), mineral dust, oceanic sulphate and so on) sea salt, soil dust and oceanic sulphate constitute a major portion of the global natural aerosol abundance (during volcanically quiescent periods) even though a proportion of the dust could also be due to anthropogenic activities (Tegen and Fung, 1994; Sokolik and Toon, 1999; Sokolik et al., 1998; Tanre´ et al., 2003; Haywood et al., 2003; Highwood et al., 2003). So the accurate estimate of natural/anthropogenic fraction is difficult to determine. Another natural component of aerosol is naturally occurring soot (smoke from natural burning such as forest fires). Natural and anthropogenic soot is the main absorbing fraction of aerosol (Crutzen and Andreae, 1990; Chylek and Wong, 1995; Kaufman et al., 1998; Jacobson, 2001; Babu and Moorthy, 2002; Sato et al., 2003) and is among the most complex aerosol

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Table 1 Source strength (data from d’Almeida et al., 1991; Andreae, 1995) Source

Emission, Tg yr1

Column burden, mg m2

Optical depth

Natural Primary Soil dust Sea-salt Volcanic dust Biological debris Secondary Sulphates Organic matter Nitrates Total Natural Anthropogenic

1500 1300 33 50

32.2 7.0 0.7 1.1

0.023 0.003 0.001 0.002

102 55 22

2.7 2.1 0.5

0.014 0.011 0.001

3060

46

0.055

Primary Industrial dust Black carbon

100 20

2.1 0.6

0.004 0.006

Secondary Sulphates Biomass burning (w/o BC) Nitrates Organic matter

140 80 36 10

3.8 3.4 0.8 0.4

0.019 0.017 0.002 0.002

390 3450 11%

11.1 57 19%

0.050 0.105 48%

Total Anthropogenic Total Anthropogenic fraction

types. It is produced both by natural and anthropogenic processes such as forest fires, man-made burning or combustion and transportation (Schwartz et al., 1995). Its radiative effects vary depending on the production mechanism. Soot has a significant role in climate modification because of its absorption characteristics (Haywood and Shine, 1995; Kaufman et al., 1997; Haywood and Ramaswamy, 1998; Haywood and Boucher, 2000; Babu and Moorthy, 2002). Even though laboratory analysis can distinguish soot from biomass burning from that of fossil fuel origin, in a global scenario it is not possible to quantify the natural fraction of soot, and it is generally believed that a major fraction of soot is produced by anthropogenic activities. Thus, we focus more on sea salt, dust and oceanic sulphates. However, for the purpose of comparison, we discuss anthropogenic counterparts as well. A simplified block diagram in Fig. 2 shows the radiative effects of the three major natural aerosols considered here. Although, the generation of sea salt and dust depends primarily on the surface wind speed, their subsequent upward transport depends on the

Cloud Cover over Ocean sea-salt Production Rate Surface Winds

Reduction in Surface Solar Flux over Ocean

Dust Production Rate

Boundary Layer Properties over Land

DMS Emission over Ocean

Iron fertilisation due to transported dust

Reduction in Surface Solar Flux & Lower Atmosphere Heating SST

Cloud Formation

Cloud Formation

Fig. 2. Block diagram showing the climate impact of natural aerosols.

boundary layer characteristics, including mixing height, vertical winds and so on. These would be different over land and sea. After production, dust aerosols are often transported long distances from their sources (Arimoto et al., 2001). Examples are dust transport from the Sahara across the Atlantic Ocean, Arabian dust transport across the Arabian Sea and dust from China across the Pacific. Mineral dust is believed to play an important role in marine biological processes (Falkowski et al., 1998). For example, dust is a source of iron, which acts as a nutrient for phytoplankton (Falkowski et al., 1998; Fung et al., 2000). This, in turn, would influence dimethyl sulphide (DMS) emission from the oceanic phytoplankton and hence natural production of sulphate aerosols over the ocean. Natural sulphate aerosols over oceans are good condensation nuclei for formation of clouds. Charlson et al. (1987) hypothesised that there exists a negative feedback mechanism by which an increased number of natural sulphate aerosols over oceans increases the cloud albedo and hence causes a reduction of surface-reaching solar radiation. This, in turn, reduces the DMS emission leading to a reduction in the natural sulphate production rate (Fig. 2). This hypothesis was extensively studied in experiments such as ACE-1 and ACE-2. Similarly, sea salt aerosols are also hygroscopic in nature and act as condensation nuclei for the formation of clouds. Dust aerosols reduce the surface-reaching solar radiation (due to scattering and absorption) while heating the lower atmosphere (due to absorption). This modifies the atmospheric boundary layer characteristics over land and ocean (Fig. 2). Over the ocean an increased concentration of dust also contributes to a reduction of surface-reaching solar radiation. The combined effect of these three major natural aerosols may have an influence on sea surface temperature. Detailed discussions on the radiative effects of each of these aerosols are included in the following sections.

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Single Scattering Albedo

1.00

0.80

Dust

Soot

Sea-salt

0.60

0.40

0.20

0.00 0

5

10

15

20

25

30

35

40

Wavelength (µm)

Fig. 3. Single scattering albedo for major aerosol species.

The radiative effects of aerosols depend strongly on the single scattering albedo (SSA) (ratio of scattering to extinction), which in turn depends on the real and imaginary components of the aerosol refractive index. The single scattering albedos of three major aerosol species (sea salt, soot and dust) are shown in Fig. 3 (data from Hess et al., 1998). It may be noted that SSAs of dust reported by various studies show discrepancies. Despite these, it can be seen that in the visible region, sea salt and sulphate are non-absorbing (SSA is close to unity) and soot is highly absorbing (SSA is 0.23 at 500 nm). In the long-wave region, sea salt and sulphate are partly absorbing and soot is completely absorbing. Recent studies have shown that a proportion of the mineral dust in the atmosphere may be of anthropogenic origin and exert significant radiative forcing (Cattrall et al., 2003; Haywood et al., 2003; Highwood et al., 2003; Tanre´ et al., 2003). However, the optical and radiative properties of dust are not known precisely. The SSA of dust at 0.55 mm reported by Hess et al. (1998) based on observations in the past was 0.84 (Fig. 3). Recent studies have shown that the refractive indices used for dust aerosols in global models are in error (Kaufman et al., 2001; Haywood et al., 2003). Kaufman et al. (2001), using remote sensing, inferred the SSA of Saharan dust as 0.97 at 0.55 mm. The studies on Saharan dust by Haywood et al. (2003) have shown that the SSA of dust at 0.55 mm is in the range of 0.95–0.99. The significantly lower SSA reported in the past (Hess et al., 1998, for example) could be due to the possible mixing of Saharan dust with biomass (possibly soot) aerosols. The Saharan dust experiment (SHADE) was designed to better understand the controlling factors that determine radiative forcing of dust (Haywood et al., 2003; Tanre´ et al., 2003). These studies suggest that mineral dust has a cooling effect and the model estimate of direct radiative forcing of Saharan dust is 0.4 W m2 (Tanre´ et al., 2003). In the terrestrial region, Saharan dust decreased the upwelling radiation at the top of the atmosphere by 6.5 W m2 and increased the surface radiation by

Fig. 4. The effect of clouds on aerosol radiative forcing.

11.5 W m2 (Highwood et al., 2003). The SSA and phase function of African mineral dust were retrieved at 14 wavelengths across the visible spectrum from groundbased measurements (Cattrall et al., 2003). The SSA showed a spectral shape expected of iron-bearing minerals but is much higher than climate models have assumed, indicating that wind-blown mineral dust cools the Earth more than is generally believed (Haywood et al., 2001; Catrall et al., 2003). The radiative effects of aerosols depend on the type and altitude of clouds as well (Heintzenberg et al., 1997; Satheesh, 2002a, b). In Fig. 4, we show a representation of a cloudy atmosphere. In case (a) most of the aerosols are concentrated below clouds whereas in case (b) aerosols are mostly above clouds. The radiative impact of aerosols in case (a) and case (b) can be significantly different even when aerosol column properties are the same. When a cloud layer is present above aerosols, most of the incident radiation will be reflected back and a small fraction only will interact with aerosols. On the other hand when an elevated aerosol layer is present with a cloud below, the aerosols interact not only with radiation incident from the Sun, but also with that reflected from the cloud layer below. This would result in an enhanced aerosol radiative impact. 3.1. Sea salt aerosols The strongest natural aerosol production rate is that of sea salt, at an estimated 1000–10,000 Tg per year (Winter and Chylek, 1997). This is about 30–75% of all natural aerosols (Blanchard and Woodcock, 1980). The source of airborne salt particles is obviously the sea. But most of the early investigators did not concentrate on the exact mechanism of production of sea salt particles. In the light of laboratory experiments, Stuhlman (1932) reported that the bursting of bubbles in distilled water produced jets of water which broke into small droplets. Later, Kohler (1936, 1941) proposed that the formation of spray at the wave crest by strong winds was

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responsible for the airborne salt particles. The highspeed photographic study of Kientzler et al. (1954) of bursting bubbles confirmed the mechanism of ejection of droplets from a breaking bubble as suggested by Stuhlman. A number of investigators have studied the wind speed dependence of the concentration of sea salt particles in the marine boundary layer (Woodcock, 1953, 1957; Monahan, 1968; Blanchard and Syzdeck, 1988; Tsunogai et al., 1972; Lovett, 1978; Monahan et al., 1982, 1983; Hoppel et al., 1990; Parameswaran et al., 1995; Moorthy et al., 1997; Moorthy and Satheesh, 2000; Vinoj and Satheesh, 2003). These showed a clear dependence of sea salt aerosol mass concentration on wind speed. Many of these investigators have suggested an exponential relation of the form C ¼ C 0 expðbUÞ,

(2)

where C is the aerosol number or mass concentration at wind speed U, C0 that at U ¼ 0 and b is a ‘wind index’. There have been a few experiments to understand the effects of natural aerosols on climate. Among these, ACE-1 was one of the major experiments. ACE-1 focussed on remote marine aerosol, minimally perturbed by continental sources, whereas ACE-2 studied the outflow of European aerosol into the northeast Atlantic atmosphere (Bates et al., 2000; Quinn et al., 2000). During ACE-2 sub-micrometre aerosol dominated scattering by the whole aerosol in contrast to ACE-1 where super-micrometre aerosol was the dominant scatterer. During the first aerosol characterisation experiment (ACE-1), extensive studies were carried out on the influence of sea salt on aerosol radiative properties (Murphy et al., 1998). However, in both ACE-1 and ACE-2, there was poor correlation between local wind speed and sea salt mass concentration (Quinn et al., 2000). On the other hand, many investigations have observed a correlation between aerosol characteristics and averaged wind (O’Dowd and Smith, 1993; Parameswaran et al., 1995; Moorthy et al., 1997, 1998; Satheesh et al., 2002). It may be noted that there is a possibility of sea salt advection from regions of high wind to regions where wind speeds are low (Gong et al., 1997, 2002; Kinne et al., 2003) and this can result in a high aerosol load even over regions of low winds. This might possibly explain the observations of poor correlation between local wind speed and sea salt mass concentration during ACE (Quinn et al., 2000). Detailed estimates of sea salt aerosol radiative forcing (Winter and Chylek, 1997) showed that at low wind speed, the sea salt radiative forcing is in the range of 0.6 to 2 W m2 and at higher wind speeds this can be as high as 1.5 to 4 W m2. This negative forcing by naturally occurring sea salt aerosol is quite significant when we consider the fact that forcing caused by projected doubling of CO2 is about +4 W m2. The

forcing caused by the increase in CO2 since the advent of the industrial era is about +1.46 W m2 (Charlson et al., 1992; Winter and Chylek, 1997). It may be noted that there are very few data on sea salt aerosols where wind speeds are high. In such conditions the measurements are difficult. Thus there is a considerable uncertainty in sea salt aerosol radiative forcing (Gong et al., 2002; Kinne et al., 2003). Another recent study has demonstrated that as wind speed increases there are two competing effects which determine the aerosol forcing at the surface; they are the increase in the single scattering albedo (SSA) and the increase in the optical depth (Satheesh, 2002a, b; Satheesh and Lubin, 2003). An increase in single scattering albedo decreases the forcing efficiency at the surface whereas an increase in optical depth increases the forcing (Heintzenberg et al., 1997). But at the top of the atmosphere (TOA), increases in both SSA and optical depth increase the forcing. The study has shown that as the sea-surface wind speed increases from 0 to 15 m s1, the magnitude of aerosol forcing at the TOA is enhanced by 6 W m2 (i.e., larger negative value) (Satheesh, 2002a, b; Satheesh and Lubin, 2003). It may be noted that the magnitude of composite aerosol forcing at the TOA observed over the tropical Indian Ocean was only 10 W m2 (Satheesh and Ramanathan, 2000; Satheesh et al., 2002). This shows that modulation in forcing by sea salt aerosols (produced by sea-surface winds) is quite significant. It also demonstrates that surface wind has a significant role in changing the chemical composition of aerosols over the sea and hence the forcing (Satheesh and Lubin, 2003). Aerosol short-wave, long-wave and net forcing as a function of wind speed is shown in Fig. 5 (data from Satheesh and Lubin, 2003). These values are in agreement with those reported by Winter and Chylek (1997). Model estimates of aerosol forcing in clear and

2.5 1.5 Aerosol TOA Forcing (W m-2)

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0.5 -0.5

2

4

6

8

SW

LW

Net

10

12

-1.5 -2.5 -3.5 -4.5 -5.5 -6.5 -7.5 -8.5

Wind Speed (m s-1)

Fig. 5. Aerosol forcing as a function of wind speed (based on data reported in Satheesh and Lubin, 2003).

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cloudy skies (when absorbing aerosols are present) have shown that aerosol forcing at the TOA decreases as cloud cover increases and can be positive if cloud coverage exceeds 25% (Podgorny and Ramanthan, 2001; Satheesh, 2002a, b). When a reflecting cloud layer is present (see Fig. 4), both aerosol scattering and absorption effects are amplified due to the multiple interactions of the radiation reflected back by clouds or between clouds and surface (Heintzenberg et al., 1997; Satheesh, 2002a, b). The effect of the sea-surface winds is thus to offset part of the heating by absorbing aerosols (as the TOA forcing by sea salt aerosol is negative in both clear and cloudy skies). 3.2. Oceanic sulphate aerosols There is evidence that fine particles are produced over the sea (Clarke et al., 1987, 1997; Clarke, 1993; Hoppel et al., 1990; Fitzgerald, 1991; Hoppel et al., 1994; Pandis et al., 1994; Russell et al., 1994; Bates et al., 2000; Quinn et al., 2000; Johnson et al., 2000; Putaud et al., 2000; Clarke and Kapustin, 2002). The resulting particles, after subsequent growth by condensation and coagulation to larger sizes (radius, R40.1 mm), play a dominant role in producing the marine stratocumulus clouds by acting as cloud condensation nuclei (CCN) over remote oceanic regions (Hoppel et al., 1986; Clarke, 1993; Lawrence, 1993; Russell et al., 1994; Bates et al., 2000; Johnson et al., 2000; Putaud et al., 2000). Aerosol measurements made over the tropical oceans have shown that the sub-micrometre aerosol size distributions can be constant for a week or longer irrespective of the prevailing meteorological conditions (Clarke et al., 1987, 1997; Hoppel et al., 1986, 1990; Pandis et al., 1994). Several investigations in clean marine air have shown that most of the particles o0.25 mm are composed of non-sea salt sulphate. Aerosol volatility measurements are in good agreement with this fact (Clarke et al., 1987; Fitzgerald, 1991; Clarke, 1993; Pandis et al., 1994). The studies as part of ACE-1 have shown that new particles are not formed in abundance in the marine boundary layer, but rather in the relatively particle-free atmosphere of the upper troposphere (at least above the marine boundary layer) (Bates et al., 2000; Quinn et al., 2000). Particles from gas-to-particle conversion are more volatile than sea salt and can be distinguished from sea salt by measuring the temperature at which the particles decomposed (Fitzgerald, 1991). By measuring the volatility of particles, Clarke et al. (1987) have shown that approximately 99% of the particles smaller than 0.2 mm radius behaved like sulphuric acid or ammonium sulphate/bisulphate and particles with r40.25 mm behaved like sea salt. Hoppel et al. (1990) measured the volatility of sub-micrometre particles (ro0.3 mm) over remote parts of the Pacific Ocean and found that most of the particles were non-sea

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salt particles except during stormy periods, during which enough salt particles can be produced. It is believed that a significant fraction of the tropospheric aerosol mass over oceans in the sub-micrometre size range is principally derived from homogeneous in-cloud oxidation of gaseous sulphur compounds (Charlson et al., 1987; Langner et al., 1992; Clarke, 1993). The sulphur compounds present over remote oceans can be of marine or continental origin. Consideration of the source strengths of various organo-sulphur gases emitted by the ocean and their rate constants for oxidation by the hydroxyl ion have lead to the conclusion that DMS is the major source of non-sea salt sulphate over oceans (Andreae et al., 1983; Fitzgerald, 1991). Charlson et al. (1987) have also proposed that DMS is the major source of aerosol sulphate in the remote marine atmosphere. Natural emissions of sulphur represent a significant part of the total flux of gaseous sulphur to the atmosphere (Andreae, 1985). Almost all species of marine phytoplankton release DMS as DMS vapour, which gets oxidised by different radicals to form SO2 (Fitzgerald, 1991; Russell et al., 1994). In the atmosphere DMS is oxidised by the several radicals including OH, NO3 and IO (Fitzgerald, 1991), the OH radical being the major oxidant. Photo-oxidation of DMS (CH3–S–CH3) with OH yields SO2, methane sulphonic acid (MSA), H2SO4 and numerous other compounds (Russell et al., 1994; Fitzgerald, 1991). The photo-oxidation products of DMS are converted to non-sea salt sulphate by gas-toparticle conversion processes (Fitzgerald, 1991). These non-sea salt sulphate particles grow by acid condensation to a radius of 0.04 mm in about two days where the particles are large enough to act as CCN, and can grow further while cycling through non-precipitating clouds (Hoppel et al., 1994). The non-sea salt sulphate particles present in the marine atmospheric boundary layer (MABL) play an important role in acting as CCN (Charlson et al., 1987; Lawrence, 1993; Clarke, 1993). The number of these particles capable of acting as CCN varies from 30 to 200 cm3 (Pruppacher and Klett, 1980; Clarke et al., 1987; Hoppel et al., 1990, 1994). Sulphate aerosols present over oceans can be of natural or anthropogenic origin. Though anthropogenic sulphur emissions can influence the sulphate concentration over oceans, in most of the remote areas of oceans, natural emissions of sulphur can account for almost all non-sea salt sulphate (Savoie and Prospero, 1982). There are only a few studies to distinguish the proportions of natural and anthropogenic components (Savoie et al., 2002 is an example). New particle formation in the atmosphere is inversely related to available aerosol surface area (Clarke, 1993). So any sudden decrease in aerosol concentration due to various removal processes (especially precipitation) will result in the homogeneous nucleation of the sulphur compounds. This leads to new

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particle formation in the marine boundary layer (MBL) (Hoppel et al., 1994). An increase in the marine DMS emission increases the number density of sulphate aerosol over the marine atmosphere and consequently the number density of cloud droplets, which results in an increase in cloud albedo (Charlson et al., 1987; Hegg, 1990; Hegg et al., 1991; Covert et al., 1992, 1996). This enhancement in cloud albedo will act as negative forcing on global temperature; lower temperature in turn results in reduced productivity and emission of marine DMS. Charlson et al. (1987) estimated that a 40–100% increase in CCN concentration is sufficient to counterbalance the temperature increase due to doubling of CO2 concentration. The organic, inorganic, mineral content and mass concentration of the sub-micrometre aerosol were measured in June and July 1997 on Tenerife in the MABL and free troposphere (Putaud et al., 2000). They observed that in the unperturbed MABL the aerosol average composition was 37% non-sea salt sulphate, 21% sea salt and 20% organic carbon. In the unperturbed free troposphere, organic carbon and non-sea salt sulphate accounted for 43% and 32% of the sub-micrometre aerosol mass respectively (Putaud et al., 2000; Schmeling et al., 2000). Based on extensive observations at the MABL simultaneously with the free troposphere (FT), these studies have concluded that the source for the free troposphere could be transport from continents; in background conditions MABL aerosol is formed by dilution of continental aerosol by FT air modified by deposition and condensation of species of oceanic origin. However, the outbreaks in the MABL were due to transport of polluted air masses from Europe. The evolution of the aerosol characteristics in the marine atmosphere was thoroughly studied during Lagrangian experiments of ACE-2 (Johnson et al., 2000). Observations during the first ACE-2 Lagrangian experiment suggested that the important processes controlling the sub-micrometre mode aerosol concentration, which dominated the total aerosol concentration, included scavenging of interstitial aerosol by cloud droplets, enhanced coagulation of Aitken mode and accumulation mode aerosols due to increased sea salt surface area and the dilution of MBL by FT air (Raes, 1995; Johnson et al., 2000). Observations during the second ACE-2 Lagrangian experiment found evidence of processing of aerosol particles by stratocumulus cloud, in particular by aqueous phase reactions (Clarke, 1998; Osborne et al., 2000; Wood et al., 2000). Measurements indicate that the concentration of DMS is higher in summer than in winter and highest over low-latitude oceans (Andreae, 1985; Bates et al., 1987). These indicate that production of DMS increases with an increase in ocean temperature, which depends

on the duration of sunlight received by the ocean surface. The warmest, most saline and most intensely illuminated regions of oceans have the highest rate of DMS emissions to the atmosphere (Russell et al., 1994). The largest DMS flux comes from the tropical and equatorial oceans (Russell et al., 1994). The concentration of non-sea salt sulphates decreases from coastal regions of the continent to the remote ocean areas (Parungo et al., 1987; Fitzgerald, 1991). 3.3. Soil dust aerosols Particles originating from the soil are usually mineral aerosols and are produced by weathering of soil (Jaenicke, 1980, 1993; Prospero et al., 1983, 2002; d’Almeida, 1986; Zender et al., 2003; Ginoux et al., 2004; Miller et al., 2004; Tegen et al., 2004). Ultra-fine sand particles are formed by winds mostly in the arid regions of the world (Pye, 1987; Schwartz et al., 1995; Prospero et al., 2002; Ginoux et al., 2004). The longrange transport of continental derived particles by the combined action of convection currents and general circulation systems make these particles a significant constituent even at locations far from their sources (Delany et al., 1967; Prospero et al., 1970, 1981; Carlson and Prospero, 1972; Prospero, 1979; Shaw, 1980; Bergametti et al., 1989; Tegen and Fung, 1994; Arimoto et al., 1995, 1997; Moorthy and Satheesh, 2000; Arimoto et al., 2001; Zender et al., 2003; Ginoux et al., 2004). Soil derived particles are among the largest aerosols with radii ranging from below 0.1 mm to 100 mm. Particles in the size range r45 mm are present only in the source regions but in general particles in the radius range 0.1–5 mm are transported long distances (5000 km) into the marine atmosphere (Arimoto et al., 2001; Prospero et al., 2002; Gong et al., 2003; Maring et al., 2003; Reid et al., 2003a, b). The measurements of aerosol size distribution and analysis of chemical composition of aerosols over Antarctica have found mineral particles with radii greater than 2 mm of Australian origin (Shaw, 1980). The data from the TOMS satellite have been extensively used to study the global distribution of dust aerosols (Prospero et al., 2002). Global maps of TOMS absorbing aerosol index shows an example of a significant amount of dust aerosols over the Sahara during the month of May (Prospero et al., 2002). When the wind pattern is favourable these aerosols are transported over the Atlantic Ocean and the Arabian Sea to reach far ocean locations (thousands of kilometres away from source). There are a number of investigations available in the literature regarding the transport of aerosols from continents to ocean and vice versa (Eriksson, 1959, 1960; Toba, 1965a, b; Junge, 1972; Delany et al., 1973; Prospero, 1979; d’Almeida, 1986; Bergametti et al., 1989; Arimoto et al., 1995; Gong et al., 2003; Zender

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220 Tg mineral dust is transported to the North Atlantic each year. An example of dust transport over the Arabian Sea is shown in Fig. 6a (using data from the moderate resolution imaging spectro-radiometer (MODIS) on board the TERRA satellite). The Arabian Sea region has a unique weather pattern on account of the Indian monsoon and the associated winds that reverse direction seasonally. Chemical analysis of aerosols over the tropical Indian Ocean have shown that more than six months every year natural aerosols contribute more than 50% to composite aerosol optical depth (Fig. 6b) (Satheesh and Srinivasan, 2002; Satheesh et al., 2002). They have demonstrated that radiative forcing due to

et al., 2003). Some of these authors found the existence of Saharan dust even over the remote areas of the Atlantic and Pacific Oceans (Carlson and Prospero, 1972; Junge, 1972; Prospero and Carlson, 1972; Prospero, 1979; d’Almeida, 1986; Bergametti et al., 1989; d’Almeida et al., 1991). Prospero et al. (1970) traced the origin of a dust event at Barbados to West Africa with a transport time of 5 days. The chemical analysis of marine aerosol samples collected over the Atlantic Ocean revealed an African source (Bergametti et al., 1989). The major source of mineral dust in Africa is the Sahara. Junge (1972) estimated that 60–200 Tg Saharan dust is generated over the Sahara and is transported each year, whereas Duce et al. (1991) estimated that

(July 2003) Aerosol Optical Thickness 30N 27N 24N 21N 18N 15N 12N 9N 6N 3N E0

50E

55E

60E

0

0.1

70E

65E 0.2

0.3

75E

0.4

0.5

80E 0.6

85E 0.8

0.7

90E

95E

0.9

1

100E

80 Natural

% Contribution in Forcing

71 70

Anthropogenic

66 61 57

60

51

49

50 40 40

47

39 34

32 30 20 F

M

A

M

J

J

A

S

O

N

D

Fig. 6. (a) Aerosol optical depths over Arabian Sea demonstrating the transport of dust aerosols from Arabian Peninsula to Indian region. (b) Contribution of natural aerosols to optical depth at Indian Ocean.

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natural aerosols in this region is about 1.5 times larger compared to that due to anthropogenic aerosols. Most of the natural aerosol forcing was contributed by dust (from the Arabian Peninsula) and locally generated sea salt. These observations are inconsistent with that reported by Li and Ramanathan (2002). For absorbing aerosols like dust, radiative forcing at the surface differs substantially from the value at the TOA and the climate response depends not only upon the TOA forcing, but its difference with respect to the surface value, which represents radiative heating within the atmosphere (Miller et al., 2004). Surface forcing alters evaporation and the hydrologic cycle. Studies by Miller et al. (2004) have shown that while global evaporation and precipitation are reduced in response to surface radiative forcing by dust, precipitation increases locally over desert regions, so that dust emission can act as a negative feedback to desertification. Dust aerosols are significant contributors to radiative warming below 500 mb due to short-wave absorption but they have less effect on long-wave radiation (Mohalfi et al., 1998; Alpert et al., 1998; Miller and Tegen, 1999 and Fig. 7). Typically, dust approximately doubles the short-wave radiation absorption under clear-sky conditions (Tegen and Miller, 1998). Tegen and Fung (1994) has shown that dust from disturbed soil causes a net cooling at the surface, accompanied by an

Fig. 7. The radiative impact of dust aerosol in short wave and long wave regions. Dust aerosols larger in size and have absorbing property in infrared and hence unlike other aerosol species, dust aerosol influence infrared as well. The symbols SW and LW represent short wave and long wave radiation and the subscripts TD, TU, BD, BU represents top of the atmosphere down-welling, top of the atmosphere upwelling, bottom of the atmosphere (surface) down-welling, and surface upwelling, respectively.

increase in atmospheric heating. Such radiative effects are found to be most pronounced over the desert regions (Mohalfi et al., 1998). There have been several investigations to understand the characteristics of the dust layer and the radiative heat balance. There are only very few studies on the impact of dust on synoptic-scale systems. The reduction of solar radiation reaching the Earth’s surface as a result of scattering and absorption by dust aerosols reduces the sensible heat flux. This is balanced by the radiative heating of dust aerosols at low levels. The dust aerosols over the Arabian Sea warm the levels between 800 and 600 hPa (0.2 K day1) and cool the lower levels during daytime (Alpert et al., 1998; Mohalfi et al., 1998). Thus the presence of dust transported over oceans intensifies a low-level inversion, which in turn affects the stability of the atmosphere (Miller and Tegen, 1999; Mohalfi et al., 1998). Both land and sea are heated during daytime by radiation from the Sun. But since solar radiation only penetrates a few centimetres of soil so that only top layer heats up. The air above heats up much more rapidly because of the low heat capacity of air. On the other hand, the sea warms up much more slowly because of the large heat capacity as well as longer penetration of solar radiation. Warm air rises over land causing a low-pressure region compared to the ocean. To compensate for this, air flows from sea to land—the well-known sea breeze. When the winds are strong enough, the land areas (especially with low vegetation cover) produce soil dust aerosols. The presence of this dust reduces the surface-reaching solar radiation due to scattering and absorption, and heats the lower atmosphere due to absorption. This cooling from below and heating aloft creates low-level inversion. This reduces the intensity of convection currents, and thus increases the atmospheric stability. The reduction of solar radiation at the surface reduces the surface heating which in turn decreases the land–sea temperature contrast and consequently the intensity of the sea breeze. Thus, depending on the concentration of the dust layer, the impact can be different. There can be changes in sea-breeze onset time also. Since stable conditions resist upward movement, we might conclude that clouds would not form when stable conditions prevail in the atmosphere. Since the surface air is cooler and heavier than the air aloft, little vertical mixing occurs between layers. Since air pollutants are added from below, temperature inversion confines them to the lowermost layers where they continuously build in concentration (Mohalfi et al., 1998). The fact that atmosphere is either stable or not, determines whether clouds develop or not. The accumulation of aerosols at lower levels would increase the lower atmosphere heating further, which in turn would increase the stability (positive feedback). Dust aerosols absorb sunlight to a greater extent than industrial sulphate and sea salt aerosols (Tegen et al.,

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1997; Miller and Tegen, 1999; Haywood and Boucher, 2000). These authors also suggest that dust optical properties (to which the top of the atmosphere forcing is sensitive) in global models should be allowed to vary with the mineral composition of the source region in a computation of the climate response. More extensive measurements of the dust optical properties, along with the vertical distribution of the dust layer, are needed to reduce the uncertainty of the climate response to dust aerosols. Liao and Seinfeld (1998) examined radiative forcing by mineral dust aerosols in short-wave and long-wave regions using a one-dimensional column radiation model. They estimated clear sky TOA radiative forcing as 2 W m2 at low surface reflection (0.1) and +2 W m2 at high surface reflection (0.5). Under cloudy skies these values are in the range of +2 to +3 W m2. They also observed that unlike scattering aerosols such as sea salt, dust radiative forcing depends on the surface reflection, the altitude at which the dust layer is located and the relative altitude from the cloud layer. Clear sky TOA long-wave radiative forcing was in the range of +0.2–1.0 W m2 and corresponding values for cloudy skies were 0.0 and +0.6 W m2. These results are consistent with Tegen and Lacis (1996) and Tegen et al. (1997). Dust can serve as a reaction surface for reactive gas species in the atmosphere (Dentener et al., 1996; Huebert et al., 2003; Carmichael et al., 2003; Seinfeld et al., 2004). Mineral dust is believed to play an important role in marine biological processes (Maher and Dennis, 2001; Prospero et al., 2002). Trace metals on dust are essential to some marine biological processes; for example, dust is a source of iron, which acts as a nutrient for phytoplankton (Falkowski et al., 1998; Fung et al., 2000; Maher and Dennis, 2001; Prospero et al., 2002; Huebert et al., 2003; Carmichael et al., 2003; Seinfeld et al., 2004). 3.3.1. Dust transport models There is a substantial transport of mineral aerosol from Asia to wide areas of the North Pacific with an estimated total annual input in the range of 6–10 million tons year1. This atmospherically transported dust is a significant source of sedimentary material for the North Pacific. Global dust distributions are usually calculated with transport models. Measurements of dust at various locations alone cannot provide information on its transport and consequent impact over other regions. Mathematical models provide the necessary framework for the integration of our understanding of various atmospheric processes and to study their interactions (Luo et al., 2003; Gong et al., 2003; Zender et al., 2003; Ginoux et al., 2004; Tegen et al., 2004). Measurements and models together provide a powerful tool to study the dust aerosol transport. Many global

2099

models do not accurately simulate regional distribution of dust due to their low grid resolution and inaccuracy of dust source function. To accurately predict the impact of dust aerosols on climate the spatial and temporal distribution of dust is essential. The dust emission is calculated depending on soil moisture, surface wind speed and soil surface conditions. The major sink is gravity settling. The model simulations have shown that the contribution of dust to aerosol optical depth is 9–27% for 201S–201N, in general, 40–66% in the Sahel region and 30–54% in East Asia (Tegen, 1994). Over the Indian Ocean dust contributes 15% to total aerosol optical depth during winter (Satheesh et al., 1999). However, regional characteristics of soil dust production, transport and removal processes are poorly understood. Recent studies have demonstrated that a fraction of the atmospheric dust load originates from anthropogenically disturbed soils (Tegen et al., 2004). By calibrating a dust source model with emission indices derived from dust storm observations, Tegen et al. (2004) estimated the contribution to the atmospheric dust load from agricultural areas to be o10% of the global dust load. Comparisons between a 22-year simulation of mineral aerosols with satellite and in situ observations suggest that the model can predict atmospheric mineral aerosol distributions, with some discrepancies (Luo et al., 2003). In addition, there were differences between the model results and previously published results (e.g., Ginoux et al., 2001). The sensitivity analysis showed that differences between simulated dusts near Australia are likely due to differences in both source parameterisation and surface winds (Luo et al., 2003). Zender et al. (2003) described a model for predicting the size-resolved distribution of atmospheric dust for climate and chemistry-related studies. The dust distribution from 1990 to 1999 is simulated with our mineral aerosol entrainment and deposition model embedded in a chemical transport model (Zender et al., 2003). Without invoking anthropogenic mechanisms the model captures the seasonal migration of the transatlantic African dust plume, and it captures the spring maximum in Asian dust outflow and concentration over the Pacific. Zender et al. (2003) estimated the 1990s’ global annual mean and variability of dust (diameter, Do10 mm) to be the following: emissions, 14907160 Tg yr1; burden, 1772 Tg; and optical depth at 0.63 mm, 0.03070.004. These values for emission, burden, and optical depth are significantly lower than some recent estimates. The model underestimates transport and deposition of East Asian and Australian dust to some regions of the Pacific Ocean. Gong et al. (2003), using a size-segregated soil dust emission and transport model, Northern aerosol regional climate model (NARCM), simulated the production

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and transport of Asian soil dust during the aerosol characterization experiment-Asia (ACE-Asia) period from March to May 2001. The model was driven by the NCEP reanalysed meteorology and has all the atmospheric aerosol physical processes of soil dust: production, transport, growth, coagulation, and dry and wet deposition. Model simulations were compared with ground-based measurements in East Asia and North America and with satellite measurements for the same period of time. The model captured most of the dust mobilisation episodes during this period in China and reasonably simulated the concentrations in source regions and downwind areas from East China to western North America. About 252.8 megatonnes of soil dust below Do40 mm was estimated to be emitted in the East Asian deserts between 1 March and 31 May 2001. Ginoux et al. (2004) simulated the global distribution of aeolian dust from 1981 to 1996 with the global ozone chemistry aerosol radiation and transport (GOCART) model. The simulated annual emission varies from a minimum of 1950 Tg in 1996 to a maximum of 2400 Tg in 1988. Of these emissions, 65% are from North Africa and 25% from Asia. It was found that North America received twice as much dust from other continents than it emits per year. The inter-annual variability of dust distribution was analysed over the North Atlantic and Africa. It was found that in winter a large fraction of the North Atlantic and Africa dust loading correlates with the North Atlantic Oscillation (NAO) index. It is shown that a controlling factor of such correlation can be attributed to dust emission from the Sahel. However, the long record of dust concentration measured at Barbados indicates that there is no correlation with the NAO index and surface concentration in winter. Longer simulation should provide the information needed to understand whether the effects of the NAO on dust distribution are rather limited or whether Barbados is at the edge of the affected region.

4. Radiative impact: natural versus anthropogenic aerosols and GHGs In this section, we compare the radiative forcing due to various (natural and anthropogenic) aerosol species as well as that due to GHGs. 4.1. Direct effect Observations over the tropical Indian Ocean have shown that TOA forcing due to sea salt aerosol is 1.3670.46 W m2 and that due to dust and soot are, respectively 0.7270.3 and +0.6470.38 W m2. The radiative forcing due to sulphate (natural and anthropogenic) aerosol was 6.4 W m2. Haywood et al. (1997), using a radiation code within a GCM, assessed

the direct radiative forcing by two major anthropogenic aerosol components: anthropogenic sulphate and soot aerosols from fossil fuel burning. They estimated that under cloudy skies, radiative forcing due to anthropogenic sulphate is 0.6 W m2 for the northern hemisphere and 0.15 W m2 for the southern hemisphere. Similar results have been reported by Haywood and Shine (1995), who report radiative forcing of 0.55 W m2 for the northern hemisphere and 0.13 W m2 for the southern hemisphere. For clear skies, Haywood et al. (1997) reported a radiative forcing of 0.59 W m2 for northern hemisphere and 0.14 W m2 for the southern hemisphere, which are comparable with cloudy sky values. In the case of soot aerosols, Haywood et al. (1997) estimated a radiative forcing of +0.35 W m2 for the northern hemisphere and +0.06 W m2 for the southern hemisphere under cloudy skies. The corresponding values under clear skies were +0.11 and +0.02 W m2. Haywood et al. (1999) have estimated clear sky radiative forcing due to natural sulphate, natural dust and sea salt as 0.93, 0.58, and 1.51 W m2 (for low sea salt; 5.03 W m2 for high sea salt), respectively. This means that radiative forcing due to natural aerosols is 3.02 W m2 (for low sea salt; 6.54 W m2 for high sea salt). They estimated the corresponding values for anthropogenic sulphate, organic carbon, black carbon and anthropogenic dust as 0.72, 1.02, +0.17, 0.54 W m2, respectively. Thus radiative forcing due to anthropogenic aerosols is 2.11 W m2. These results clearly show the significant role natural aerosols have in determining the radiative forcing due to a composite aerosol system. Using an aerosol transport model coupled with a GCM, Tekemura et al. (2002) estimated radiative forcing due to various aerosol species. The global mean radiative forcing due to black carbon under cloudy skies was +0.36 W m2 and that due to anthropogenic sulphate was 0.32 W m2. The corresponding values for clear sky conditions were +0.21 and 0.72 W m2, respectively. These values are slightly smaller than those estimated by Penner et al. (1998) and Kiehl et al. (2000), but comparable with those estimated by Boucher and Anderson (1995) and Feichter et al. (1997). Sea salt and dust radiative forcing were +0.36 and 0.31 W m2 under cloudy skies and +0.26 and 0.59 W m2 under clear sky conditions. The value of dust radiative forcing is higher than the value of +0.14 W m2 reported by Tegen et al. (1996). The estimate of Tekemura et al. (2002) of radiative forcing due to organic carbon, black carbon and anthropogenic sulphate (total anthropogenic forcing of 0.96 W m2) is comparable with that of Haywood et al. (1999) when considering that anthropogenic dust was not included in Tekemura et al. (2002). They did not provide forcing due to natural sulphate. If we use the

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-2

TOA Direct Forcing (W m )

2.5

Global Average

1.5 0.5 -0.5 -1.5 -2.5 -3.5 -4.5 -5.5 -6.5 GHGs

Sea-salt

Dust

BC

OC

Sulphate Sulphate (N) (A)

Indian Ocean : Regional

TOA Forcing (W m-2 )

value of natural sulphate forcing from Haywood et al. (1999), radiative forcing due to natural aerosols is 1.26 W m2. An atmospheric general circulation model is coupled to an atmospheric chemistry model to calculate the radiative forcing by anthropogenic sulphate and carbonaceous aerosols (Penner et al., 1998). They estimated that the direct forcing by anthropogenic sulphate aerosols is in the range of 0.55 to 0.81 W m2. The climate forcing associated with fossil fuel emissions of carbonaceous aerosols is calculated to range from +0.16 to +0.20 W m2. The direct forcing of carbonaceous aerosols associated with biomass burning is calculated to range from 0.23 to 0.16 W m2. Myhre et al. (1998) estimated that the direct radiative forcing due to sulphate and soot is 0.32 and +0.16 W m2, respectively. The above discussion shows that the radiative forcing due to sea salt aerosols ranges from 0.5 to 6.0 W m2 while that of natural dust aerosols ranges from 2 to +0.5 W m2. Now, we discuss the IPCC (2001) estimates of the radiative forcing due to anthropogenic aerosols. The global mean direct radiative forcing due to anthropogenic sulphate aerosols reported by IPCC ranges from 0.26 to 0.82 W m2 based on several studies (Kiehl and Briegleb, 1993; Boucher and Anderson, 1995; Feichter et al., 1997; Graf et al., 1998; Haywood et al., 1997; Hansen et al., 1998; Haywood and Ramaswamy, 1998). The IPCC estimates of black carbon (BC) aerosols from fossil fuel and biomass burning is in the range +0.27 to +0.54 W m2, and the corresponding estimate for organic carbon (OC) is in the range 0.04 to 0.41 W m2 (Hansen et al., 1998; Jacobson, 2001). It should be noted that uncertainties in these estimates are large due to the limited number of studies available. Next, we come to radiative forcing due to GHGs. Myhre et al. (1998) have performed calculations of the radiative forcing due to changes in the concentrations of the most important well-mixed GHGs since pre-industrial time, and found that the radiative forcing due to all the well-mixed GHGs is +2.25 W m2. IPCC reports that radiative forcing due to major GHGs such as CO2, CH4, N2O is +1.46, +0.48 and +0.15 W m2, respectively. The total radiative forcing due to well-mixed GHGs is 2.43 W m2. Thus negative forcing by naturally occurring aerosols is quite significant when we consider the fact that forcing caused by projected doubling of CO2 is about +4 W m2 (Charlson et al., 1992; Winter and Chylek, 1997). A comparison of the radiative forcing due to various aerosol species with that of GHGs is shown in Fig. 8a (data obtained from the literature discussed in Sections 2 and 3 and summarised in Table 2). It can be seen that sea salt aerosol forcing (and its variability) is quite large compared to other species.

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0 -1 -2 -3 -4 -5 -6 -7 -8 -9 -10

Anthropogenic Forcing = -5 ± 2.5 W m-2 Sea-salt Direct Forcing = -2 ± 1 W m-2 Sea-salt Indirect Effect = -7 ± 4 W m-2 Anthropogenic

Sea-salt Direct

Sea-salt Indirect

Fig. 8. (a) Comparison of greenhouse gas forcing with that of aerosol forcing due to various species. (b) Natural vests anthropogenic forcing over tropical Indian Ocean [The data from the following sources: Kiehl and Briegleb, 1993; Boucher and Anderson, 1995; Tegen and Lacis, 1996; Feichter et al., 1997; Graf et al., 1998; Haywood et al., 1997; Tegen et al., 1997; Moorthy et al., 1997; Winter and Chylek, 1997; Alpert et al., 1998; Mohalfi et al., 1998; Miller and Tegen, 1999; Haywood and Ramaswamy, 1998; Penner et al., 1998; Haywood et al., 1999; Satheesh and Ramanathan, 2000; Podgorny et al., 2000; Jacobson et al., 2001; Ramanathan et al., 2001; Satheesh, 2002; Tekemura et al., 2002; Soden et al., 2002; Satheesh and Lubin, 2003; Vinoj and Satheesh, 2004].

4.2. Indirect effect Sea salt aerosols and natural sulphates are hygroscopic in nature and hence act as condensation nuclei for the formation of clouds (Fitzgerald, 1991). Cloud albedo has a significant role in determining the global energy balance (Chuang et al., 1997). An increased concentration of aerosols results in an enhanced concentration of cloud droplets, which in turn increases the albedo of clouds and this causes a decrease in the short-wave solar radiation reaching the Earth’s surface (Clarke, 1998). The increase in condensation nuclei (CN) also influences the cloud lifetime. An increase in CN increases the cloud droplet concentration and reduces the mean droplet size. This increases the cloud lifetime and inhibits precipitation. This also leads to an increase in fractional cloud coverage and influences both short-wave and long-wave

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Table 2 Comparison of direct radiative forcing (W m2) by various species Location (regional/global)

Species

Radiative forcing (W m2)

Reference

Global

Sea salt

Winter and Chylek (1997)

Global

Sea salt

Deserts Indian Ocean Indian Ocean Indian Ocean Indian Ocean Indian Ocean Northern Hemisphere

Dust Sea salt Sea salt Dust Soot (BC) Sulphate (natural and anthropogenic) Sulphate (anthropogenic)

0.6 to 2.0 (low) 1.5 to 4.0 (high) 1.51 (low) 5.03 (high) 2 to +2 1.3670.5 1.5 to 6.0 0.7270.3 +0.6470.4 6.470.5 0.55 to 0.6

Southern Hemisphere

Sulphate (anthropogenic)

0.13 to 0.15

Northern Hemisphere Southern Hemisphere Global Global Global Global Global Global Global Global Global Global

Soot (BC) Soot (BC) Sulphate (anthropogenic) Soot (BC) Sulphate (natural) Dust Sulphate (anthropogenic) Soot (BC) Dust Sulphate (anthropogenic) BC (FFB) BC (BB)

+0.11 +0.02 0.72 +0.17 0.58 0.93 0.72 +0.21 +0.14 0.26 to 0.82 +0.27 +0.57

radiation. Cloud albedo depends on the cloud droplet number. For a given water vapour content, the average cloud droplet size is larger for a lower number of aerosols and is smaller for a higher number of aerosols (Han et al., 1998). This is because the water vapour availability per CN is more in the former case compared to latter case. But the relation between aerosol number and number of cloud droplets is not simple and depends on a number of factors, including the aerosol chemical composition, size distribution, supersaturation of air and so on (Clarke, 1993; Ramanathan et al., 2001). Not all aerosols are capable of acting as CN. To be able to act as CN, the aerosol should be larger than a critical size (1 mm) and should be hygroscopic (water-soluble) (Hoppel et al., 1990, 1994). As the number of aerosols increases, the supersaturation (S) reduces. The inverse correlation is due to the fact that as more drops form, the water supply available will be less and as a result S is reduced (Ramanathan et al., 2001). Based on direct measurements of aerosols, cloud droplet concentration and supersaturation over the tropical Indian Ocean, Ramanathan et al. (2001) derived empirical relations between aerosol number and various parameters such as cloud drop number, cloud drop effective radius, cloud optical depth and so on. Their

Haywood et al. (1999) Liao and Seinfeld (1998) Podgorny et al. (2000) Satheesh and Lubin (2003) Podgorny et al. (2000) Podgorny et al. (2000) Podgorny et al. (2000) Haywood and Shine (1995) Haywood et al. (1997) Haywood and Shine (1995) Haywood et al. (1997) Haywood et al. (1997) Haywood et al. (1997) Haywood et al. (1999) Haywood et al. (1999) Haywood et al. (1999) Haywood et al. (1999) Tekemura et al. (2002) Tekemura et al. (2002) Tegen et al. (1996) IPCC (2001) IPCC (2001) IPCC (2001)

basic equation is of the form, N CCN ¼ 0:12N 1:25 ðS=3Þ0:76 ,

(3)

where NCCN is the number of aerosols which are activated, N is the total number of particles, and S is the supersaturation in percentage. The equation is valid for values of N ranging from 300 to 2000 cm3 and So0.3%. Here S is a function of N as the amount of water vapour available per nuclei depends on the total N for a given water vapour amount. They also found from observed data that not all CCN becomes cloud droplets. When the total aerosol number is low almost all CCN becomes cloud droplets, whereas at high aerosol concentrations, only about 80% of the CCN becomes cloud droplets. The effective radius of cloud droplets decreases from 8.0 to 5.5 when aerosol numbers change from 300 to 2000 cm3 (Ramanathan et al., 2001). The number of cloud droplets increases from 75 to 300 cm3 and the corresponding cloud optical depth increases from 3 to 14 for the same change in aerosol number (Ramanathan et al., 2001). Investigations have revealed that sea salt number concentration over the ocean is a function of wind speed (Lovett, 1978; Blanchard and Woodcock, 1980; O’Dowd and Smith, 1993; Parameswaran et al., 1995; Moorthy

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et al., 1997). O’Dowd and Smith (1993) reported that sea salt number increased by 100 cm3 when wind speed increased from 3 to 15 m s1. This observation, when combined with observations of Ramanathan et al. (2001), shows that a change in wind speed from 3 to 15 m s1 can change the cloud droplet number by 30 cm3 and increase the cloud optical depth by 3. The estimates of sea salt direct and indirect effects over the Indian Ocean were 271 and 774 W m2, respectively (Vinoj and Satheesh, 2004). This is quite large compared to anthropogenic aerosol forcing reported over this region (572.5 W m2) (Ramanathan et al., 2001). Thus, clearly the direct and indirect effects of sea salt aerosols have a significant role in offsetting the positive forcing by absorbing aerosols and GHGs. The direct and indirect forcing due to sea salt aerosols compared with anthropogenic forcing over the Indian Ocean is shown in Fig. 8b. The magnitude of indirect radiative forcing (and uncertainty) due to sea salt aerosols is several-fold more than the direct radiative forcing of sea salt aerosols. The large magnitude and variability in both direct and indirect forcing due to sea salt aerosols emphasises the importance of natural aerosols. Soil dust is not hygroscopic and as such does not participate as CCN. There are two extremes of insoluble nuclei: nuclei which are activated (wetted) easily, and nuclei which are not easily activated. Nuclei which are easily activated rapidly, get coated with liquid and subsequently behave like droplets and further grow in size by condensation (Levin et al., 1996; Wurzler et al., 2000; IPCC, 2001). The droplet growth thereafter can be predicted by using Kelvin’s equation. In cases where nuclei surfaces are not wettable, condensation proceeds with much more difficulty. The surfaces of the nuclei try to make the condensing liquid into small spheres. When the entire surfaces are covered with these small spheres, liquid coatings can form. Hereafter the nuclei behave like normal droplets and grow in size by condensation of vapour. Soil dust is often internally mixed with other species and thus can be hygroscopic (Prospero et al., 2002). Levin et al. (1996) observed that desert dust was coated with sulphate, which probably originated from in-cloud scavenging of interstitial dust particles followed by evaporation of the cloud droplets. The presence of soluble materials over dust makes them into large and effective CCN, which may affect cloud microphysics (Levin et al., 1996; IPCC, 2001). The role of insoluble nuclei in condensation is still a question to be answered (Levin et al., 1996; Wurzler et al., 2000). The IPCC estimates of the indirect radiative effect due to anthropogenic sulphate ranges from 0.3 to 1.8 W m2 based on various studies (Chuang et al., 1997; Boucher and Lohman, 1995; Jones and Slingo, 1996, 1997). Chuang et al. (2002) obtained an indirect

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radiative forcing due to black carbon and organic carbon aerosols of 1.51 W m2. Kaufman and Nakajima (1993) have estimated the indirect radiative forcing by smoke to be 2 W m2 using satellite data over Brazil.

5. Summary and conclusions Aerosols are of natural or anthropogenic origin. Natural aerosols account for 70% of the global aerosol loading and of this the main contributors are sea salt, dust and natural sulphates. Nevertheless, the abundance of these shows significant variability from region to region and season to season. Recent investigations have shown that a proportion of dust is due to anthropogenic activities. Similarly, a proportion of anthropogenic soot originates from natural forest fires. It is difficult to separate the anthropogenic components of dust from natural, or natural components of soot from anthropogenic. Besides, away from the source regions, both natural and anthropogenic components mix together and on a global scale it is almost impossible to exactly apportion the natural and anthropogenic shares of the total aerosol. Nevertheless, several investigations and coordinated field campaigns have been carried out to assess the impact of anthropogenic aerosols on climate (particularly because they are amenable to mitigation). The ACE-2, Tropospheric Aerosol Radiative Forcing Experiment (TARFOX), Indian Ocean Experiment (INDOEX) are examples. The ACE-1 and ACE-Asia, however, have provided valuable information on natural aerosols. Even so, there are still far fewer studies on natural aerosols compared with anthropogenic aerosols, despite their importance. To accurately predict the impact of dust aerosols on climate, the spatial and temporal distribution of dust is essential. However, regional characteristics of soil dust production, transport and removal processes are poorly understood. Many global models do not accurately simulate regional distribution of dust due to their low grid resolution and inaccuracy of dust source function. To accurately predict the impact of dust aerosols on climate the spatial and temporal distribution of dust is essential. More extensive measurements of the dust optical properties, along with the vertical distribution of the dust layer, are needed to reduce the uncertainty of the climate response to dust aerosols. Similarly, there are very few data on sea salt aerosols where wind speeds are high. In such conditions accurate measurements are extremely difficult. Thus the data on the global distributions of two major natural aerosol types (sea salt and mineral dust) are not adequate. Several experiments and simulations have attempted to quantify the radiative impacts of natural aerosols, particularly sea salt, dust and oceanic sulphate, yet large

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uncertainties persist in these estimates especially due to the following:

 Fewer 

data over oceans, especially on sea salt aerosols, and limitations of available observations over oceans. At high wind speeds, it is difficult to make accurate measurements over oceans. Inadequate understanding of optical/radiative properties of dust aerosols, their large regional differences (depending on the soil type at the source region), the transport processes and sinks.

Thus, more data using well-focussed experiments are needed to reduce the uncertainties in the characteristics of these two major natural aerosol species from a global perspective. Notwithstanding the above, estimates have shown that even at low wind speeds, radiative forcing due to sea salt aerosols can be in the range from 0.5 to 2 W m2 and at higher wind speeds this can be as high as in the range 1 to 6 W m2. This negative forcing (cooling) by naturally occurring sea salt aerosols is quite significant when we consider the fact that the forcing caused by the increase in CO2 since the advent of the industrial era is about +1.46 W m2, and forcing caused by projected doubling of CO2 is about +4 W m2. Similarly detailed estimates of the dust radiative forcing shows values in the range of 2 to +0.5 W m2. It may be noted that due to the poor data on regional characteristics of soil dust source function, we do not even know whether dust radiative forcing is positive or negative. Most of the recent investigations, however, indicate that dust radiative forcing is negative. Thus, natural aerosols contribute quite significantly to global radiative forcing. This contribution has large seasonal and spatial variability and is comparable to or even larger than anthropogenic forcing. Though no steps can be taken to reduce these effects (unlike anthropogenic effects, which are amenable to mitigation), a clear understanding is needed to appreciate the climate impact of aerosols on the one hand and the extent of perturbation caused by human activities on the other hand. Thus to assess the climate impact of aerosols, while separating out the human factor, we need to address the following issues. A. To what extent can wind-generated sea salt aerosols offset the atmospheric heating due to absorbing aerosols such as soot transported over oceans? B. At high wind speeds, newly produced sea salt droplets may coat over pre-existing absorbing soot aerosols, thus significantly altering their absorbing efficiency. What is the consequent impact on radiative forcing? This could significantly alter the

C.

D.

E.

F.

G.

aerosol properties not only over oceans but also over a significant part of the continents along the vast coastal areas of the globe. The reduction in solar radiation at the surface simultaneous with lower atmospheric heating by dust aerosols could intensify a low-level inversion and reduce the sensible heat flux. How does this impact the formation of clouds? The dust containing iron transported over the ocean serves as nutrients to marine phytoplankton. The consequent enhancement in DMS emission (due to iron fertilisation) will increase the natural sulphate aerosols over the ocean, which may have an influence on cloud droplet concentration, cloud albedo and hence alter the radiation balance as much as, or at times even more than, the changes brought about by anthropogenic sulphates over oceans. Dust optical properties vary from region to region. Recent investigations have shown that the dust absorption is lower than that assumed in global models. Regional distribution of dust source function is poorly understood due to lack of an adequate database. Prior to 2001, international panels such as the Intergovernmental Panel on Climate Change (IPCC) focussed mainly on the anthropogenic aerosol components. In IPCC (2001), great effort was made to assess the impact of natural aerosols. It is, however, true that we lack sufficient information on natural aerosols over large areas of the world, especially the oceans. The presence of natural aerosols influences the anthropogenic aerosol forcing (either directly or indirectly). I t is difficult to separate the natural and anthropogenic aerosol contributions on radiative forcing when they are in a mixed state.

Hence there is an urgent need to focus attention on radiative effects of natural aerosols, especially in the tropics where data on aerosols are sparse.

Acknowledgements The authors thank ISRO for supporting this study through the ISRO-Geosphere Biosphere Programme. We thank Prof. J. Srinivasan, CAOS, Indian Institute of Science, for valuable suggestions.

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