Radiative effects of natural aerosols: A review

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Atmospheric Environment 39 (2005) 2089–2110

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Radiative effects of natural aerosols: A review

S.K. Satheesha,�, K. Krishna Moorthyb

aCentre for Atmospheric and Oceanic Sciences, Indian Institute of Science, Bangalore-560 012, IndiabSpace Physics Laboratory, Indian Space Research Organisation (ISRO), Vikram Sarabhai Space Centre, Trivandrum-695 022, India

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

e front matter r 2005 Elsevier Ltd. All rights reserved.

mosenv.2004.12.029

ing author. Tel.: +9180 22933070/22932505; fax: +9180 23600865.

ess: [email protected] (S.K. Satheesh).

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ARTICLE IN PRESSS.K. Satheesh, K. Krishna Moorthy / Atmospheric Environment 39 (2005) 2089–21102090

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 de-

signed to increase the understanding of how atmo-

spheric 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, respec-

tively, 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

ARTICLE IN PRESS

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

aerosols.

S.K. Satheesh, K. Krishna Moorthy / Atmospheric Environment 39 (2005) 2089–2110 2091

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,

H ¼S0

4ð1� aÞ � sT4

e ¼ 0, (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)

(1365–1372Wm�2) (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 343Wm�2 of the mean solar

radiation incident at the top of the atmosphere,

�103Wm�2 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

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 theirchemical 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 propor-

tion 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

ARTICLE IN PRESS

Table 1

Source strength (data from d’Almeida et al., 1991; Andreae,

1995)

Source Emission,

Tg yr�1Column

burden,

mgm�2

Optical

depth

Natural

Primary

Soil dust 1500 32.2 0.023

Sea-salt 1300 7.0 0.003

Volcanic dust 33 0.7 0.001

Biological debris 50 1.1 0.002

Secondary

Sulphates 102 2.7 0.014

Organic matter 55 2.1 0.011

Nitrates 22 0.5 0.001

Total Natural 3060 46 0.055

Anthropogenic

Primary

Industrial dust 100 2.1 0.004

Black carbon 20 0.6 0.006

Secondary

Sulphates 140 3.8 0.019

Biomass burning (w/o BC) 80 3.4 0.017

Nitrates 36 0.8 0.002

Organic matter 10 0.4 0.002

Total Anthropogenic 390 11.1 0.050

Total 3450 57 0.105

Anthropogenic fraction 11% 19% 48%

Iron fertilisation dueto transported dust

Surface Winds

sea-salt ProductionRate

Reduction in SurfaceSolar Flux over

Ocean

Cloud Coverover Ocean

DMS Emissionover Ocean

Dust ProductionRate

BoundaryLayer

Propertiesover Land

Reduction inSurfaceSolar Flux & LowerAtmosphereHeating

Cloud Formation

SSTCloud

Formation

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

aerosols.

S.K. Satheesh, K. Krishna Moorthy / Atmospheric Environment 39 (2005) 2089–21102092

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

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 trans-

port 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 sul-

phate 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.

ARTICLE IN PRESS

0.00

0.20

0.40

0.60

0.80

1.00

0 5 10 15 20 25 30 35 40

Wavelength (µm)

Sin

gle

Sca

tter

ing

Alb

edo

Dust Soot Sea-salt

Fig. 3. Single scattering albedo for major aerosol species.Fig. 4. The effect of clouds on aerosol radiative forcing.


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 Saharandust by Haywood et al. (2003) have shown that the SSA

of dust at 0.55mm 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 �0.4Wm�2 (Tanre et al.,

2003). In the terrestrial region, Saharan dust decreased

the upwelling radiation at the top of the atmosphere by

6.5Wm�2 and increased the surface radiation by

11.5Wm�2 (Highwood et al., 2003). The SSA and

phase function of African mineral dust were retrieved at

14 wavelengths across the visible spectrum from ground-

based 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 1000–10,000Tg 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

ARTICLE IN PRESS

-8.5

-7.5

-6.5

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

0.5

1.5

2.5

2 4 6 8 10 12

Wind Speed (m s-1)

Aer

oso

l TO

A F

orc

ing

(W

m-2

)

SW LW Net

Fig. 5. Aerosol forcing as a function of wind speed (based on

data reported in Satheesh and Lubin, 2003).


responsible for the airborne salt particles. The high-

speed 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 C0 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 character-

istics 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 �2Wm�2 and at higher wind speeds this can be

as high as �1.5 to �4Wm�2. 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 +4Wm�2. The

forcing caused by the increase in CO2 since the advent of

the industrial era is about +1.46Wm�2 (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

15m s�1, the magnitude of aerosol forcing at the TOA is

enhanced by �6Wm�2 (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 �10Wm�2 (Satheesh and Rama-

nathan, 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

ARTICLE IN PRESSS.K. Satheesh, K. Krishna Moorthy / Atmospheric Environment 39 (2005) 2089–2110 2095

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 coagula-

tion to larger sizes (radius, R40.1mm), play a dominantrole 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 o0.25mm 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 atmo-

sphere 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 0.2mm radius behaved like

sulphuric acid or ammonium sulphate/bisulphate and

particles with r40.25mm behaved like sea salt. Hoppel

et al. (1990) measured the volatility of sub-micrometre

particles (ro0.3 mm) over remote parts of the PacificOcean and found that most of the particles were non-sea

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 oxida-

tion 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 phyto-

plankton 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), H2SO4and numerous other compounds (Russell et al., 1994;

Fitzgerald, 1991). The photo-oxidation products of

DMS are converted to non-sea salt sulphate by gas-to-

particle conversion processes (Fitzgerald, 1991). These

non-sea salt sulphate particles grow by acid condensa-

tion to a radius of�0.04mm in about two days where theparticles 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 cm�3 (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 concentra-

tion 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


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 counter-

balance the temperature increase due to doubling of CO2concentration.

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 concentra-

tion, 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 concentra-

tion 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 long-

range 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 inthe size range r45 mm are present only in the source

regions but in general particles in the radius range

0.1–5mm are transported long distances (�5000 km) intothe 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 kilo-

metres 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


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; Pros-

pero, 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 60–200Tg Saharan

dust is generated over the Sahara and is transported

each year, whereas Duce et al. (1991) estimated that

49

71

5

32

40 39

20

30

40

50

60

70

80

F M A M J

% C

on

trib

utio

n in

Fo

rcin

g

30N

50E 55E 60E 65E 70E 75E

0 0.1 0.2 0.3 0.4 0.5

(July 2Aerosol Optic

27N

24N

21N

18N

15N

12N

9N

6N

3N

E0

Fig. 6. (a) Aerosol optical depths over Arabian Sea demonstrating th

region. (b) Contribution of natural aerosols to optical depth at India

�220Tg 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

761

51

34

47

66

J A S O N D

Natural Anthropogenic

80E 85E 90E 95E 100E

0.6 0.7 0.8 0.9

003)al Thickness

1

e transport of dust aerosols from Arabian Peninsula to Indian

n Ocean.


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 500mb 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 investiga-

tions 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 (�0.2Kday�1) 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.,


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 �2Wm�2 at low surface reflection (�0.1) and

+2Wm�2 at high surface reflection (�0.5). Under

cloudy skies these values are in the range of +2 to

+3Wm�2. 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 +0.2–1.0Wm�2 and corresponding values

for cloudy skies were 0.0 and +0.6Wm�2. 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 year�1. 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

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 produc-

tion, transport and removal processes are poorly

understood.

Recent studies have demonstrated that a fraction of

the atmospheric dust load originates from anthropo-

genically 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 atmo-

spheric mineral aerosol distributions, with some dis-

crepancies (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 distribu-

tion 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,

14907160Tg yr�1; burden, 1772Tg; and optical depthat 0.63 mm, 0.03070.004. These values for emission,burden, and optical depth are significantly lower than

some recent estimates. The model underestimates trans-

port 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 regio-

nal climate model (NARCM), simulated the production


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

�1.3670.46Wm�2 and that due to dust and soot are,

respectively �0.7270.3 and +0.6470.38Wm�2. The

radiative forcing due to sulphate (natural and anthro-

pogenic) aerosol was �6.4Wm�2. 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 anthropo-

genic sulphate is �0.6Wm�2 for the northern hemi-

sphere and �0.15Wm�2 for the southern hemisphere.

Similar results have been reported by Haywood and

Shine (1995), who report radiative forcing of

�0.55Wm�2 for the northern hemisphere and

�0.13Wm�2 for the southern hemisphere. For clear

skies, Haywood et al. (1997) reported a radiative forcing

of �0.59Wm�2 for northern hemisphere and

�0.14Wm�2 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 +0.35Wm�2 for the northern hemisphere

and +0.06Wm�2 for the southern hemisphere under

cloudy skies. The corresponding values under clear skies

were +0.11 and +0.02Wm�2.

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 �1.51Wm�2 (for low

sea salt; �5.03Wm�2 for high sea salt), respectively.

This means that radiative forcing due to natural aerosols

is �3.02Wm�2 (for low sea salt; �6.54Wm�2 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,

�0.54Wm�2, respectively. Thus radiative forcing due

to anthropogenic aerosols is �2.11Wm�2. 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 +0.36Wm�2 and that due to anthropogenic

sulphate was �0.32Wm�2. The corresponding values

for clear sky conditions were +0.21 and �0.72Wm�2,

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 �0.31Wm�2

under cloudy skies and +0.26 and �0.59Wm�2 under

clear sky conditions. The value of dust radiative forcing

is higher than the value of +0.14Wm�2 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 �0.96Wm�2) 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

ARTICLE IN PRESS

-6.5

-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

0.5

1.5

2.5

3.5

GHGs Sea-salt Dust BC OC Sulphate(N)

Sulphate(A)

TO

A D

irec

t F

orc

ing

(W

m-2

)

Global Average

Indian Ocean : Regional

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

Anthropogenic Sea-salt Direct Sea-salt Indirect

TO

A F

orc

ing

(W

m-2

)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

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].


value of natural sulphate forcing from Haywood et al.

(1999), radiative forcing due to natural aerosols is

�1.26Wm�2.

An atmospheric general circulation model is coupled

to an atmospheric chemistry model to calculate the

radiative forcing by anthropogenic sulphate and carbo-

naceous aerosols (Penner et al., 1998). They estimated

that the direct forcing by anthropogenic sulphate

aerosols is in the range of �0.55 to �0.81Wm�2. The

climate forcing associated with fossil fuel emissions of

carbonaceous aerosols is calculated to range from

+0.16 to +0.20Wm�2. The direct forcing of carbonac-

eous aerosols associated with biomass burning is

calculated to range from �0.23 to �0.16Wm�2. Myhre

et al. (1998) estimated that the direct radiative forcing

due to sulphate and soot is �0.32 and +0.16Wm�2,

respectively.

The above discussion shows that the radiative forcing

due to sea salt aerosols ranges from �0.5 to �6.0Wm�2

while that of natural dust aerosols ranges from �2 to

+0.5Wm�2. Now, we discuss the IPCC (2001) esti-

mates 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 �0.82Wm�2 based on several

studies (Kiehl and Briegleb, 1993; Boucher and Ander-

son, 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 +0.54Wm�2, and the

corresponding estimate for organic carbon (OC) is in the

range �0.04 to �0.41Wm�2 (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-indus-

trial time, and found that the radiative forcing due to all

the well-mixed GHGs is +2.25Wm�2. IPCC reports

that radiative forcing due to major GHGs such as CO2,

CH4, N2O is +1.46, +0.48 and +0.15Wm�2, respec-

tively. The total radiative forcing due to well-mixed

GHGs is 2.43Wm�2. 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 +4Wm�2 (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.

4.2. Indirect effect

Sea salt aerosols and natural sulphates are hygro-

scopic 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 concentra-

tion 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 precipita-

tion. This also leads to an increase in fractional cloud

coverage and influences both short-wave and long-wave

ARTICLE IN PRESS

Table 2

Comparison of direct radiative forcing (Wm�2) by various species

Location (regional/global) Species Radiative forcing (Wm�2) Reference

Global Sea salt �0.6 to 2.0 (low) Winter and Chylek (1997)

�1.5 to 4.0 (high)

Global Sea salt �1.51 (low) Haywood et al. (1999)

�5.03 (high)

Deserts Dust �2 to +2 Liao and Seinfeld (1998)

Indian Ocean Sea salt �1.3670.5 Podgorny et al. (2000)

Indian Ocean Sea salt �1.5 to �6.0 Satheesh and Lubin (2003)

Indian Ocean Dust �0.7270.3 Podgorny et al. (2000)

Indian Ocean Soot (BC) +0.6470.4 Podgorny et al. (2000)

Indian Ocean Sulphate (natural and anthropogenic) �6.470.5 Podgorny et al. (2000)

Northern Hemisphere Sulphate (anthropogenic) �0.55 to �0.6 Haywood and Shine (1995)

Haywood et al. (1997)

Southern Hemisphere Sulphate (anthropogenic) �0.13 to �0.15 Haywood and Shine (1995)

Haywood et al. (1997)

Northern Hemisphere Soot (BC) +0.11 Haywood et al. (1997)

Southern Hemisphere Soot (BC) +0.02 Haywood et al. (1997)

Global Sulphate (anthropogenic) �0.72 Haywood et al. (1999)

Global Soot (BC) +0.17 Haywood et al. (1999)

Global Sulphate (natural) �0.58 Haywood et al. (1999)

Global Dust �0.93 Haywood et al. (1999)

Global Sulphate (anthropogenic) �0.72 Tekemura et al. (2002)

Global Soot (BC) +0.21 Tekemura et al. (2002)

Global Dust +0.14 Tegen et al. (1996)

Global Sulphate (anthropogenic) �0.26 to �0.82 IPCC (2001)

Global BC (FFB) +0.27 IPCC (2001)

Global BC (BB) +0.57 IPCC (2001)


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

basic equation is of the form,

NCCN ¼ 0:12N1:25ðS=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 cm�3 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 cm�3 (Ramanathan et al.,

2001). The number of cloud droplets increases from �75

to �300 cm�3 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


et al., 1997). O’Dowd and Smith (1993) reported that sea

salt number increased by 100 cm�3 when wind speed

increased from 3 to 15m s�1. This observation, when

combined with observations of Ramanathan et al.

(2001), shows that a change in wind speed from 3 to

15m s�1 can change the cloud droplet number by

�30 cm�3 and increase the cloud optical depth by �3.

The estimates of sea salt direct and indirect effects over

the Indian Ocean were �271 and �774Wm�2,

respectively (Vinoj and Satheesh, 2004). This is quite

large compared to anthropogenic aerosol forcing

reported over this region (�572.5Wm�2) (Rama-

nathan 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

�1.8Wm�2 based on various studies (Chuang et al.,

1997; Boucher and Lohman, 1995; Jones and Slingo,

1996, 1997). Chuang et al. (2002) obtained an indirect

radiative forcing due to black carbon and organic

carbon aerosols of �1.51Wm�2. Kaufman and Naka-

jima (1993) have estimated the indirect radiative forcing

by smoke to be �2Wm�2 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 investiga-

tions 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


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 prop-
erties 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 �2Wm�2

and at higher wind speeds this can be as high as in the

range �1 to �6Wm�2. 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 +1.46Wm�2, and forcing caused

by projected doubling of CO2 is about +4Wm�2.

Similarly detailed estimates of the dust radiative forcing

shows values in the range of �2 to +0.5Wm�2. 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 mitiga-

tion), 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

aerosol properties not only over oceans but also over

a significant part of the continents along the vast

coastal areas of the globe.

C.
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?

D.
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.

E.
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 func-

tion is poorly understood due to lack of an adequate

database.

F.
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.

G.
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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Radiative effects of natural aerosols: A review

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