How Aerosols Affect the Global Climate and the Relation to Climate Sensitivity
Gunnar Myhre, Cathrine E. Lund Myhre, Bjørn H. Samset, Trude Storelvmo
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Atmospheric aerosols are part of the human influence on climate, and limited understanding of their
effects is a major source of uncertainty in our knowledge about climate change. Aerosols enter the
atmosphere from both anthropogenic and natural sources and affect the global radiation balance
depending on their sizes and chemical composition. We describe the complexity associated with
aerosols, from emissions to climatic effects, and illustrate their importance for the global climate
sensitivity. We also show how the aerosol size distribution and chemical composition vary globally
due to the multitude of available sources. Finally we describe recent progress in the understanding of
aerosol climatic effects and insights into the prevailing uncertainties.
What is the source of the aerosols?
Atmospheric aerosols are suspensions of liquid, solid, or mixed particles with highly variable chemical
composition and size. The variability is due to the numerous sources and varying formation
mechanisms (Figure 1). Aerosols are mainly emitted directly to the atmosphere ( primary aerosols) or
produced in the atmosphere from precursor gases (secondary aerosols), but may also be influenced
by evaporation of clouds (cloud processing).
Primary aerosols have both inorganic and organic sources. Inorganic primary aerosols are relatively
large (often larger than 1 m) and result from sea spray, mineral dust, and volcanoes. These coarse
aerosols have short atmospheric life times, typically less than a few days. Primary aerosols with
carbonaceous constituents stem from plant/microbial materials, combustion processes, and biomass
burning. The carbonaceous primary aerosols include both organic carbon (OC) in aqueous phase and
solid black carbon (BC). BC is the light-absorbing carbonaceous fraction present in aerosols. This
compound is of particular importance for climate. The main sources are combustion processes of
organic material and fires, but domestic sources (cooking) are also very important in certain regions.
Primary BC and OC containing aerosols are generally smaller than 1 m.
Secondary aerosols are produced in the atmosphere from precursor gases by condensation of
vapours from the gas phase on pre-existing particles or by nucleation of new particles. Secondary
aerosols are small; they range in size from a few nanometres up to 1 m and have lifetimes of days
to weeks. Secondary aerosols consist of mixtures of compounds; the main components are sulphate,
nitrates, and OC. The main precursor gases are anthropogenic emissions from fossil fuel combustion,
but fires and biogenic emissions of volatile organic compounds (VOCs) are also important.
Occasionally volcanic eruptions result in huge amounts of aerosols both at the ground and in the
stratosphere: the volcanic eruption of Mt Pinatubo in 1991 resulted in high concentrations of
secondary sulphate aerosols in the stratosphere during the following years.
The size and chemical composition of the aerosols evolve with time through coagulation and
condensation. Mixing of different aerosol types is of importance for their climate effect. The aerosols
may grow by uptake of water, a process which depends on chemical composition, aerosol size, and
ambient relative humidity.
Figure 1: Atmospheric aerosols: their dominating sources, typical structure, and appearance in the
atmosphere. A mixture of aerosol pollution is shown at the top: the picture to the left shows urban
aerosol pollution and the picture to the right shows aerosol pollution from the India and Bangladesh
region. The main natural and anthropogenic sources are, from the top left corner and
counterclockwise: The volcanic eruption of Grimsvotn, 21 May 2011, emitting primary volcanic ash
and SO2 resulting in sulphate aerosols. Sea spray, illustrating sea salt aerosols and secondary
sulphate aerosols from dimethyl sulphide. Desert storm in Iraq, 26 April 2005 showing mineral dust
aerosols. Savannah biomass burning, illustrating both anthropogenic and natural fires as sources of
OC and BC containing aerosols. Coal power plants, illustrating secondary sulphate aerosols and other
industrial and fossil aerosol sources. Ship in a Norwegian fjord resulting in aerosols containing OC, BC,
sulphates, and nitrates. Cooking, representing a large domestic source of BC and OC. A truck,
representing the transport sector and emissions of BC and VOCs yielding OC in aerosols. The core of
the figure shows examples of aerosol structures illustrated by an electron microscope image of black
carbon attached to sulphate particles. The spherical structures in image A are sulphates; the arrows
point to smaller chains of black carbon. Black carbon is shown in detail in image B. Image C shows fly
ash, a product of coal-combustion,. Credit for microscope picture: Peter Buseck, Arizona State
University.
How are aerosols distributed globally?
Figure 2 shows the aerosol optical depth (AOD) determined by remote sensing from space by the
MODIS instrument. The AOD pattern is highly inhomogeneous, with the largest values in the tropical
regions in Africa and Asia. The pie charts show the contributions from different aerosol types in
selected regions. In general there is large variability in the global aerosol composition. The different
aerosol species have varying impacts in the atmosphere, and the wide spread of both sources and
sizes further complicates a quantification of their effects. Remote sensing from both space and the
ground has substantially advanced the understanding of the total aerosol distribution, but there are
still large uncertainties in the chemical composition and the anthropogenic contribution to the
observed AOD.
Figure 2: MODIS aerosol optical depth (AOD) averaged over the 10 year period 2001–2010 (Remer et
al., 2008). AOD is the column integrated aerosol extinction, where extinction is the attenuation of
solar light given in terms of length (m-1).The contribution to the total AOD from different aerosol
types in selected regions, as indicated on the figure (pie charts), is estimated by use of a global
aerosol model based on a simulation for one year (Myhre et al., 2009). The aerosol types are Sul
(sulphate), BC and OC from use of fossil fuel, Bio (OC and BC from biomass burning), Nitrate, Sea (sea
salt), and Min (mineral dust). Water uptake of the aerosols is taken into account and enhances the
radiative properties of some of the aerosol types substantially, e.g. sulphate aerosols, which increase
in size due to hygroscopic growth. In the grey areas there is no data available from MODIS.
How do aerosols affect the climate?
All atmospheric aerosols scatter incoming solar radiation, and a few aerosol types can also absorb
solar radiation. BC is the most important of the latter, but mineral dust and some OC types are also
solar absorbers. Aerosols which mainly scatter solar radiation have a cooling effect, by enhancing the
total reflected solar radiation from the Earth. Strongly absorbing aerosols have a warming effect. In
the atmosphere there is a mixture of scattering and absorbing aerosols, and the degree of scattering
is quite variable. It is quantified through the single scattering albedo  = s/, where s is scattering
AOD and  is total AOD. The net effect of the scattering and absorbing aerosols is dependent on
surface and cloud characteristics. Scattering aerosols above a dark surface and absorbing aerosols
above a bright surface are most efficient (see Figure 3a). Scattering aerosol above a bright surface is
less efficient since the solar radiation is reflected anyway; similarly for absorbing aerosol and a dark
surface, the solar radiation will be to a large extent absorbed. Since clouds are the main contributors
to the total reflection of solar radiation, absorbing aerosols are particularly efficient when positioned
above clouds.
Aerosols are vital for cloud formation, acting as sources of cloud condensation nuclei (CCN) and ice
nuclei (IN). An increased amount of aerosols may increase the CCN number concentration and lead
to more, but smaller, cloud droplets. This increases the albedo of the cloud, resulting in enhanced
reflection and a cooling effect, termed the cloud albedo or first indirect effect (Twomey, 1977); see
Figure 3b). Smaller and more numerous cloud droplets may also inhibit the precipitation process,
since many similarly sized cloud droplets reduce the growth to cloud droplets necessary for initiation
of precipitation. This effect, called the cloud lifetime or second indirect effect, may enhance the
cloud cover and also has a cooling effect (Albrecht, 1989).
Absorbing aerosols also have the potential to modify clouds properties, without directly acting as
CCN and IN, by: i) reducing the solar radiation reaching the ground, which diminishes the convection
and thus the potential for cloud formation, ii) changing the temperature profile, which will in most
cases lead to reduced instability and weakened cloud formation, iii) increasing the temperature,
which reduces the relative humidity, inhibits cloud formation, and enhances evaporation of existing
clouds. This is termed the semi-direct aerosol effect (Hansen et al., 1997). The net effect is uncertain
and depends highly on the vertical profile of BC (Koch and Del Genio, 2010).
In addition BC (and other absorbing aerosols) deposited on snow or ice surfaces may reduce the
albedo, leading to reduced reflectance of solar radiation, and hence a heating effect (Hansen and
Nazarenko, 2004).
Figure 3: a) An illustration of the direct aerosol effect for low surface albedo (in this example
zero surface albedo) and high surface albedo (in this example all solar radiation is reflected
and surface albedo equals one) for scattering and absorbing aerosols. A dark surface (i.e. low
surface albedo) will absorb a large portion of the solar radiation, and absorbing aerosols will
thus have a small effect since solar radiation is absorbed anyway. On the other hand,
scattering aerosols may substantially amplify the total reflectance of solar radiation since the
solar radiation will otherwise be absorbed at the surface. Similarly over a bright surface (high
surface albedo), scattering aerosols have a small effect since the solar radiation is scattered
anyway. Absorbing aerosol may however substantially reduce the outgoing radiation under
such conditions and has a warming effect. b) Illustration of the first indirect aerosol effect
(cloud albedo effect), second indirect aerosol effect (cloud lifetime effect), and semi-direct
effect. Different sizes of cloud droplets, cloud fraction, and initiation of the precipitation
process are shown.
Why is the uncertainty in the aerosols important for predictions of the climate sensitivity?
The change in temperature (T) as a result of radiative forcing (RF, a change in the Earth’s radiation
balance due to a perturbation of anthropogenic or natural origin) can be expressed by the following
simple heat balance equation:
c d(T)/dt = RF -T
(1)
Here c is the heat capacity and  is the climate sensitivity. At radiative equilibrium (d(T)/dt = 0),
Equation 1 reduces to T = RF/. However the Earth is not in radiative equilibrium, since less thermal
radiation is currently emitted to space compared to what is absorbed of solar radiation (Hansen et
al., 2005). This radiative imbalance causes the Earth to gradually warm, with global warming as a
result (Trenberth and Fasullo, 2010). The simple equation above has two key uncertainties. The
observed surface temperature change is rather well determined, but the climate sensitivity and the
total RF are both highly uncertain. Quantifying the climate sensitivity , an essential parameter for
prediction of future climate change, has long been attempted by using global climate models or
temperature records, but it still has a wide range of reported values (IPCC, 2007; Knutti and Hegerl,
2008). The total RF through the industrial era is also uncertain, mainly due to lack of quantification of
the aerosol effects discussed above. The implication of this uncertainty in the aerosol RF for the
quantification of the climate sensitivity can be illustrated as follows:
Figure 4a shows the climate sensitivity as a function of the total aerosol RF (shown for both the most
certain aerosol effects and a more comprehensive set of aerosol effects), where uncertainties in the
radiative imbalance are included. A similar figure has previously been presented in Andreae et al.
(2005). The figure shows that using the known industrial age warming of around 0.8K and the
present best knowledge on RF from non-aerosol components, radiative imbalance and the 90%
confidence interval of the total aerosol RF for the most certain effect, we can get a climate sensitivity
range from about 2 to 8 K for a doubling of CO2. Figure 4b breaks down the total aerosol forcing
probability density function (PDF) into individual components, indicating both the magnitudes and
uncertainties of the effects.
Figure 4: a) Climate sensitivity for a doubling of CO2 as a function of the total aerosol RF. The curve is
shown for a best estimate of the radiative imbalance of 0.85 Wm-2 (Hansen et al., 2005) with grey
shading for imbalances between 0.7 and 1.0 Wm-2 representing the uncertainty. A thin dotted line for
radiative imbalance of zero (radiative equilibrium) is added. The observed temperature change is
taken as 0.8 K and radiative forcing of non-aerosol components is 2.9 Wm-2 over the industrial era.
The red line shows the probability density function (PDF) for the total of the direct aerosol, cloud
albedo effect, and the BC on snow and ice, whereas the dotted red line shows the total of all aerosol
effects. b) PDFs of various aerosol effects taken from Isaksen et al. (2009), except a small update for
the cloud albedo effect; the range for the cloud lifetime effect is here estimated to be somewhat
weaker, and for the mixed phase and ice clouds recent studies are taken into account to narrow the
range somewhat. Monte Carlo simulations are performed to derive total aerosol RF, shown as the
red curves (Boucher and Haywood, 2001).
Has there been any progress in the understanding of the climate effect of aerosol?
There has been a tremendous improvement in the understanding of atmospheric aerosols and their
climate effect over the last decades, with some important observational and modelling
breakthroughs. Long term measurements of aerosols (e.g.Putaud et al., 2004), observational
campaigns (e.g.Quinn and Bates, 2005), and remote sensing from space and ground (Holben et al.,
1998; Remer et al., 2008) have remarkably increased the knowledge about the atmospheric aerosol
composition and characteristics. However, the understanding of the greater complexity of
atmospheric aerosols has at the same time limited more robust quantification of their climate effect.
The first estimate of the direct aerosol effect in the early 1990s was limited to sulphate aerosols
(Charlson et al., 1991), with estimates for BC coming a few years later (Haywood and Shine, 1995).
Observations have shown that OC is an important aerosol component (Novakov et al., 1997;
Ramanathan et al., 2001) and substantial investigations have later explored the complex composition
and optical characteristics of this compound (e.g.Graber and Rudich, 2006; Kanakidou et al., 2005).
Global aerosol models today provide RF estimates for a large set of aerosol components such as
sulphate, BC (from fossil fuel and biomass burning), OC (primary and secondary from fossil fuel and
biomass burning), nitrate (Jacobson, 2001; Koch et al., 2009; Liao and Seinfeld, 2005). In addition
multi-model studies are performed to understand and reduce uncertainties due to model differences
(Schulz et al., 2006).
An example of recent progress is reduced uncertainty in the estimate of the total direct aerosol
effect given by IPCC AR4 (Forster et al., 2007). This estimate was made possible by the advances that
have occurred on both the modelling and the observational side, and was based on a combination of
global aerosol models and observation based methods (mostly sensed data). Initially, the two
approaches gave conflicting results, the observation based methods yielding a stronger RF than the
models. Consistency between these two different approaches has later been reached, and was found
to arise from necessary and simplified assumptions of the pre-industrial aerosol composition in the
observation based method (Myhre, 2009). Although the uncertainty in the total direct aerosol effect
is reduced, the uncertainty for several of the aerosol components such as BC, OC, and nitrate is still
large.
Similar to the early estimates of the direct aerosol effect, many of the first model estimates of the
aerosol indirect effect only accounted for the effect of sulphate particles acting as CCN (Jones et al.,
1994; Kaufman and Chou, 1993). Furthermore, they only included the influence of sulphate aerosols
on cloud albedo, disregarding any effects on cloud lifetime and extent. With the realization that
other aerosol species of anthropogenic origin could also form cloud droplets and that effects on
cloud lifetime and extent were also possible, global climate models estimated the aerosol indirect
effect to be stronger (e.g.Lohmann and Feichter, 1997; Menon et al., 2002). Some even predicted
this cooling effect to be comparable in magnitude to the warming greenhouse effect. Recent
publications have later pointed to oversimplifications in the model representation of clouds and how
their lifetimes are affected by aerosols (e.g.Stevens and Feingold, 2009). It is now acknowledged that
aerosol effects on cloud lifetime will vary with the cloud type in question. For example, a few recent
model studies have found that by forming ice in super-cooled liquid clouds, aerosols may in fact
shorten cloud lifetime, because of the more efficient precipitation formation when cloud ice is
present (e.g.Lohmann and Hoose, 2009; Storelvmo et al., 2011). In summary, whether aerosols are
acting as CCN or IN or are simply modifying atmospheric stability by absorbing solar radiation, there
is still high uncertainty associated with their effect on cloud lifetime. In contrast, model and satellite
estimates of the cloud albedo effect seem to agree on a negative RF that has about half the
magnitude of the positive RF attributed to increasing CO2 concentrations.
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