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Aerosol & Climate

Direct and Indirect Effects

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Many Sources / Lifetimes

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Aerosols are liquid or solid particles suspended in the air. They can scatter and absorb both solar and terrestrial radiation. This is called “direct radiative forcing”. Clouds particles (both liquid and ice) rarely form directly from homogeneous nucleation (the direct formation of clouds from water vapor), but rather are formed on seed aerosol particles (cloud condensation nuclei). Because the chemical and microphysical properties of the aerosol influence both how/when clouds form and the radiative properties of the clouds, aerosols are said to exert an “indirect effect”. As we have discussed earlier, clouds play a critical role in Earth’s climate. The climate science community has identified the forcing by both the direct and indirect effects of aerosol as both important and poorly understood (IPCC 2000, 2007).

Aerosols range in size from very small clusters only a few nanometers (10-9 m) in diameter to several µm. Aerosol number density range from 10 cm-3 in the lower stratosphere during volcanically quiescent times to 103 cm-3 in clean tropospheric air (particularly in the southern hemisphere) to 106 cm-3 or more in polluted urban environments. In LA basin, aerosols limit the visibility significantly. Aerosol larger than ~20 µm sediment quickly and thus have a relatively short lifetime. The radiatively important properties of aerosols (both direct and indirect) are determined at the most fundamental level by the composition and size distribution.

Figure 11.4 of Hartmann illustrates the size distribution typical of atmospheric aerosol. The number density is dominated by very small particles (10 nm). These very small particles have little mass or surface area and in general do not directly influence climate. The largest contribution to surface area (and thus scattering) come from aerosols with radii between 0.1 and 1 µm. These particles are formed by the coagulation of the smallest particles (this mode is often called the “accumulation” mode). The mass is often dominated by the largest particles (so-called course mode) with radii near 10 µm.

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Extinction, Optical Depth, and all that Extinction is the interaction of light with aerosol (and clouds). It is

(somewhat) analogous to absorption by gases. The extinction coefficient, σext or kext (mass units) represents the

efficiency of removing photons from the direct beam:

F(x) F(x + dx) Absorbed flux dFabs

Scattered flux dFscat

x x + dx Recall that the relationship between F(x+dx) and F(x) is given by the Beer-Lambert extinction equation:

F(x+dx) = F(x)exp[-n (σabs + σscat)L] where L is the pathlength.

Optical depth (τ) is given by: n (σabs + σscat)L. where L is generally the vertical column of the atmosphere. If we measure the direct solar beam (with a “sun photometer” at the ground, the radiance will be attenuated by exp(-τ/cos(SZA)) (assuming a plane-parallel atmosphere). Note that this doesn’t tell us what happened to the photons, only that they are no longer in the direct solar beam. This is different than absorption by gases where the photon no longer exists

Extinction, Optical Depth, and all that (cont) Where did the photons go? This is very important for climate.

One of the most important properties of aerosol with respect to the direct forcing is the ratio of scattering to total extinction – known as the single scattering albedo, ωo :

ωo ≡ Qsca / Qext.

Where Q is the ratio of the absorption (kabs) or extinction coefficient (kabs+kscat) to the geometric size of the aerosol (or shadow area). Q and ωo are strongly wavelength dependent. Q can exceed 1! A change in ωo (the fraction scattered to the total extinction) from 0.9 to 0.8 can, depending on the nature of the underlying surface, change the sign of the direct effect. This has led to major efforts to try to understand the amount of soot or “black carbon” in aerosol globally – something that is quite difficult to measure.

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Another property of aerosol critical to direct radiative forcing is the scattering phase function, P, which is normalized :

where dω is the increment of solid angle (steradians). The phase function can be characterized by the single scatter asymmetry factor cosθ defined as:

where θ is the angle between the direction of the incident beam and the scattered beam. The single scatter asymmetry factor varies between 1.0 (complete forward scattering and thus minimal climate influence) and

-1.0 complete backscattering.

An asymmetry factor of 0 implies isotropic scattering. For typical accumulation mode aerosol, the scattering of solar radiation will follow the aerosol surface area while absorption of terrestrial radiation will follow the aerosol mass (and thus increase linearly with effective radius for a given visible optical depth). Most aerosol of size between 0.2 and 2 µm have asymmetry factors close to 1 (mostly forward scattering). This is quite apparent by observation of the sky. Note the halo that is often seen around the sun. This brightness reflects the highly forward scattering of atmospheric aerosol. As the size of the scatterer becomes much smaller than the wavelength of light, the asymmetry factor becomes close to 0. Rayleigh scattering (by molecules) is a good example. Note the sky again - the blue sky is not substantially brighter near the sun than far away from the sun.

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The composition of atmospheric aerosol So what are these aerosols made of? It is a virtual cornucopia of the

periodic table (see "In situ measurements of organics, meteoritic material, mercury, and other elements in aerosols at 5 to 19 kilometers", Murphy DM, et al., Science, 282, 1664, 1998). Nevertheless it is useful to discuss a few classes of aerosol.

Stratosphere: While searching for debris from nuclear bomb tests, Christian Junge

discovered in 1960 a layer of microscopic aerosol particles between the tropopause and about 30 km altitude. This layer is called the Junge Layer or the Stratospheric Aerosol Layer. These particles of mean size 0.1 µm diameter are produced from the condensation of sulfuric acid with a co-condensation of small amounts of water (25-50% by weight). The sulfuric acid is produced from the oxidation of SO2, OCS (produced in the surface ocean), and in fact transport of aerosol from the troposphere. In the absence of volcanic emissions, this aerosol has negligible optical depth (though they greatly influence the ozone chemistry). Eruption of large volcanoes can, however, increase the stratospheric aerosol optical depths by orders of magnitude.

During the last twenty years, the effects of stratospheric aerosols has become much better understood by studies of the eruption of El Chichon in southern Mexico (1982, 17.3 oN) and Mt. Pinatubo in the Philippine Islands (1991, 15.1 oN). Peak, globally-averaged aerosol optical depths were 0.07 and 0.15 following these two volcanic eruptions. These volcanoes inject large amounts of SO2 directly into the stratosphere. Most of the ash is significantly large that it falls quickly. The SO2 is oxidized in the stratosphere over a period of ~ 3-6 months by the hydroxyl radical (and subsequent reaction with H2O) and the resulting sulfuric acid accumulates on the aerosol. These aerosols have a residence time of 1 to 2 years due to their sedimentation and mixing back into the troposphere (where they are scavenged by surfaces and rain). Sulfate has a number of strong absorption features between 3 and 20 µm and in addition to reflecting sunlight these aerosols absorb terrestrial radiation efficiently, heating the stratosphere and producing a small greenhouse effect (Figure 11.6). Figure 11.7 shows that for a visible optical depth of 0.1, provided that the particle size is less than 2 µm, the net influence of stratospheric aerosol is to cool the surface. Mt. Pinatubo is estimated to have produced a net radiative forcing of –4 Wm-2. Note that this implies that the asymmetry factor is nearly 1.0 (why?). The stratospheric aerosol is lost by advection into the troposphere (important for nearly all sizes), and by sedimentation (only important for particles > 1 µm).

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Hartmann

Hartmann

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Explosive eruptions have produced significant short-term cooling of the surface. The cooling is moderated by the thermal inertia of the climate system (particularly the oceans) over the few year residence time of stratospheric aerosol. In 1815 Mt. Tambora in Indonesia exploded and within a few months the optical effects of the stratospheric aerosol was observed in Europe. It is estimated that 100 Tg Sulfur was added to the stratosphere in this single eruption. The sun was dimmed noticeably for nearly 2 years with stratospheric optical depths greater than 1 at its peak. The year 1816 was anomalously cold with crop failures widespread. The influence of the volcano was embedded within what was already an anomalously cold decade (1810-1820).

Robock

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Mt. Pinatubo

SODEN ET AL., SCIENCE 296, 727(2002)

Tropospheric Sulfate Aerosol Just as with the stratosphere, sulfuric acid is important for

formation of tropospheric aerosol. Unlike the stratosphere, however, the residence time of tropospheric aerosol is much shorter (one-two weeks). As a result, a much larger source of sulfur is required to produce high optical depths. Such high optical depths do occur, but the aerosol distribution is quite variable (as a result of the short lifetime). Emission of SO2 by industrial activity (from sulfur containing fossil fuels) accounts for the majority of all sulfur emissions to the atmosphere (80 Tg S / yr). Volcanic emissions and reduced sulfur gas production in the oceans (H2S and DMS which are relatively quickly oxidized to H2SO4) account for the most of the rest (35 Tg S/yr). These sources, however, produce sulfate with a longer residence time because they form at higher altitude. Thus the sulfate burden is more evenly distributed than the source strength's would imply. Volcanic emissions to the troposphere are also quite important for the same reason.

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From Ruddiman.

U.S. SO2 Emissions for

Phase I and Phase II Units

http://www.epa.gov/airmarkets/arp/overview.html

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Average SO2 burdens over USA, Europe and China: 2005-2006

25.5 million tons of SO2 was emitted by Chinese factories in 2005

up 27% from 2000

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Dust Soil dust is a major contributor to aerosol loading and optical thickness,

especially in sub-tropical and tropical regions. Dust sources are mostly from the deserts, dry lake beds (go to the Owen's valley), and in agricultural areas during soil disturbance. It is estimated that ~1/2 of the dust results from soil disturbance. The residence time depends (obviously) on the size of the particles. The largest particles fall quickly, while the submicron sizes can be transported over long distances. During the ACE-Asia campaign (which the Seinfeld/Flagan groups worked on), a huge dust storm from China was observed to spread a pall over the entire Pacific. In addition to radiative effects, these dust particles carry iron to the oceans where iron can be the limiting nutrient for biological production. Dust has a single scattering albedo significantly less than 1, the resulting direct forcing is small due to the partial cancellation of solar and thermal forcing.

Sea Salt: The action of waves on the ocean produces sea spray and with the bursting

of entrained air bubbles sea salt aerosol is formed. Where winds are strong, sea salt aerosol is often the most important contributor to both light scattering and cloud nuclei. Sea salt particles cover a wide size range (0.05 - 10 um diameter), and thus have a correspondingly diverse atmospheric lifetime. For the present climate, it is estimated that more than 3000 Tg/yr of sea salt aerosol is formed (IPCC 2000).

Carbonaceous Aerosol: Carbonaceous aerosol make up a large (but highly variable) fraction of

atmospheric aerosol. Organics are the largest single component of biomass burning aerosol. Measurements over the Atlantic suggest organics are as important as sulfur to the aerosol mass (IPCC 2000). In the upper troposphere, organics can comprise the majority of aerosol mass. Much of the organic appears to be oxygenated and polar (low vapor pressure), particularly carboxylic and dicarboxylic acids. As a result, these particles are quite hydroscopic and participate as CCN. Carbonaceous aerosols form both directly (as in fires) or by secondary accumulation of oxidation products formed in the gas phase. The formation of so-called secondary aerosol is very important in the LA basin, for example. Primary biogenic aerosols are also produced by ablation of organic material at the surface and lofting of small particles (bacteria, fungi, viruses, algae, pollen, etc.). Humic-like substances are formed by ablation of leaf waxes. These aerosols are often quite efficient for absorption of light shortward of 400 nm.

A second class of carbonaceous aerosol is called "black carbon". This is largely elemental carbon such as soot. Small amounts of black carbon can greatly influence the radiative impact of aerosols particularly in the presence of high optical depths of non-absorbing aerosol. This is an area of very active research.

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Nitrates: Aerosol nitrite is also quite important. It is observed that many

aerosols are near neutral pH. Nitrate and ammonia are quite efficient for forming aerosol. Ammonium nitrate aerosol is ubiquitous and very efficient absorber. Although nitrate aerosol is thought to be only perhaps 1/10 as important radiatively as sulfur at present, it is expected to become grow in importance with the further industrialization in Asia. Locally, nitrate is already quite important. Observations suggest for example, that nitrate aerosol is more abundant over India than sulfate.

If you travel toward Riverside, past the agricultural areas of the basin, you can directly observe the very high aerosol optical depths produced when acid aerosol grow rapidly when neutralized with addition of ammonia (to produce ammonium nitrate and sulfate).

Aerosols – The indirect radiative effect From - K. Noone, Department of Meteorology, Stockholm University,

Sweden http://web.mit.edu/igac/www/newsletter/igac23/noone.html

The indirect effect of aerosols (how changes in aerosol influence clouds) is thought to be one of the largest uncertainties in current global climate models. IPCC 2000. There are several indirect ways that aerosols influence the radiative balance of the Earth. Aerosols contribute to determining the cloud albedo ("Twomey effect"). They also affect precipitation processes. Aerosols influence the initial droplet size distributions produced close to cloud base, and can subsequently influence the effectiveness of coalescence at a later stage of cloud development through changing the spread in drop sizes. Aerosols also influence freezing processes in mixed-phase clouds. Both these mechanisms (coalescence and freezing) influence precipitation development. Precipitation is a key component in determining the lifetime and extent of clouds. It also is a key component in the atmospheric energy balance through the redistribution of latent heat. Finally, it is possible that aerosols can influence the dynamical processes in the atmosphere that drive cloud formation and development.

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An early observation that anthropogenically-produced aerosols could change cloud properties came when “anomalous cloud lines”—bright curvilinear structures in low clouds—were seen in early TIROS satellite imagery [Conover JH, J Atmos. Sci, 23 ,778, 1966]. The suspicion was that these areas of increased cloud albedo were caused by aerosol emissions from ships. These features have subsequently become known as ship tracks and serve as a textbook example of the effect pollution aerosols can have on the albedo of warm clouds.

The Albedo Effect Perhaps the first to be recognized (and receive the most attention) is the albedo or “Twomey” effect—the increase in cloud albedo due to an increase in aerosol concentration. For a dynamic forcing that creates a cloud with a given vertical extent and liquid water content, an increase in aerosol concentration going into the cloud can result in the formation of a larger number of smaller droplets as compared to an unperturbed cloud. The end result is in an increase in cloud albedo [Twomey, Atmos. Environ., 8, 1251, 1974].

The albedo effect does not necessarily imply an increase in cloud reflectivity. The albedo (α) of a cloud is a function of several variables: α = α(τ, ω0, cosθ) where τ is the optical depth of the cloud, ωo is the single scattering albedo of the droplets, and cosθ is the asymmetry parameter. An aerosol/cloud interaction that can change any of these parameters can also affect cloud reflectivity. For instance, ω0 is approximately unity for water droplets at visible wavelengths. If a significant amount of absorbing aerosol were to be incorporated into cloud droplets, it could reduce ϖ0 and decrease α. On the whole, however, increasing aerosol concentrations are expected to increase cloud albedo for all but the thickest clouds.

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Precipitation: Cloud lifetime, extent and energy redistribution

The same processes that increase cloud albedo in low-level clouds (a production of more and smaller droplets) tend to decrease the efficiency with which precipitation is formed. Albrecht et al., ["observations of marine stratocumulus clouds during fire", Bull. Am. Met. Soc., 69, 618, 1988], proposed that a decrease in drizzle production in these kinds of clouds could increase both the cloud liquid water content (and thus liquid water path) and the fractional cloudiness. Aerosol-induced precipitation suppression has been observed both with in situ measurements [Ferek RJ et al., "Drizzle suppression in ship tracks", J. Atmos. Sci., 57, 2707, 2000], and in satellite observations [Rosenfeld D., "Suppression of rain and snow by urban and industrial air pollution", Science, 287: 1793, 2000] showing that the effect does in fact occur in the atmosphere.

If precipitation is suppressed, water that would have been removed from the atmosphere remains aloft and can be transported to other locations before it is deposited to the surface. The same is true for the energy associated with this water—the latent heat released on condensation in clouds and the energy required for evaporation of water from the surface. This redistribution of water and latent heat due to precipitation suppression may have the potential to influence circulation patterns.

Another kind of indirect effect may arise due to the presence of absorbing (black carbon) aerosols. It has been hypothesized that heating of the boundary layer by absorbing aerosols may influence cumulus cloud formation by stabilizing the layer and reducing relative humidity [Ackerman AS et al., Reduction of tropical cloudiness by soot, Science, 288, 1042, 2000]. Field experiments in polluted winter fogs showed that despite the presence of considerable amounts of absorbing aerosol, fog still persisted for several days [Noone, K.J., Tellus B, 44, 489, 1992].

These results are not necessarily contradictory—what they do show is that this subject is not yet very well understood. The number and complexity of the steps in the causal relationship between aerosols and precipitation makes quantifying this indirect radiative effect very challenging. Even more challenging is understanding and quantifying the consequences of the effect in terms of circulation patterns and energy redistribution. Clearly, there is still plenty of work to be done in this area.