21 Properties of Aerosols

Dr. Vijay Shridhar

Contents

1. Introduction
2. Concentrations
3. Morphological Properties
4. Condensation of aerosol particles Molality, mol/kg
5. Coagulation of aerosol particles
6. Chemical speciation of Atmospheric Particulate Matter
7. Size differentiated particulate matter
8. Aerosol Optical Depth (AOD) or Aerosol Optical Thickness (AOT)
9. Black carbon (BC) and its radiative forcing
10. References

Introduction

Aerosol is a suspension of fine particles in a gas, usually air, and is generally taken to include both solid and liquid particles with dimensions ranging from a few nanometers up to around 100 micrometers in diameter. Aerosol science is the study of the physics and chemistry of aerosol properties and behavior in the atmosphere and in technical processes involving particulate matter such as combustion. This includes techniques of generating particles of nanometer and micrometer dimensions, size classification and measurement of transport and deposition properties, chemical properties of aerosol particles in the atmosphere and in industry, as well as health effects.

The environmental impacts of atmospheric particles depend on their physical, chemical properties, lifetimes and abundances. The concentration, size distribution and composition of atmospheric aerosol particles are highly variable in both time and space.

Airborne behavior, such as settling velocity, is a function of: Size, Specific gravity, Shape, Surface properties. The Aerodynamic Equivalent Diameter (AED) of a particle is the diameter of a unit density sphere that would have the identical settling velocity as the particle.

Concentrations:

Different measures are commonly used to describe the aerosol concentration: these are number, area volume and mass concentrations

Name	Unit	Description
Number concentration	cm-3	N
area concentration	µm2cm-3	4πr2N
Volume concentration	µm3cm-3	4/3πr3N
mass concentrations	µg m-3	4/3πr3Nρ

In the lower troposphere, the total particle number concentration typically varies in the range of about 102–105 cm–3, and typical mass concentration varies between 1 and 100 µg m–3. Aerosol concentrations in the free troposphere are typically 1–2 orders of magnitude lower than in the atmospheric boundary layer.

Morphological Properties

Shape

Aerosols may have range of shape depending on its formation and chemical composition. These shapes can be divided into three class.

1. Isomeric particles are those for which all three dimensions are roughly the same. Spherical, regular polyhedral or particles approximating these shapes belong in this class.
2. Platelets are particles that have two long dimensions and a small third dimension. i.e leaf fragments, scales, disc etc
3. Fibers are particles with high length in one dimension compared to much smaller lengths in other two dimension.

1. Size: All particles having similar settling velocities are considered to be of the same size, regardless of their actual size, composition or shape. Aerosols are mono disperse if all Particles are of only a single size. Polydisperse contain particles of more than one size. Similarly Homogenous aerosols are the particles having one chemical composition while Heterogenous aerosols constitute a combination of chemical composition .To study the polydisperse aerosol, measurement of the diameter of a single particle is not sufficient to describe all particles and need to make certain assumption in place of measuring particle diameter on the basis of Feret’s diameter (maximum distance from edge to edge) or Martin diameter(length of line that divide each particle in to two equal portion). Hence for simplification of measurement problem, projected area diameter has been widely used. PMx (particulate matter with diameter smaller than x µm) is another often-used measure to describe the aerosol mass concentration. PM (2.5) and PMSize distribution

   Aerosol particles in the atmosphere have extensively variable shapes. Their dimensions are usually characterized by a particle diameter, which span over four orders of magnitude, from a few nanometers to around 100 μm. Particle size is one of the most important parameters to describe the behavior of aerosols, affecting both their lifetime, physical and chemical properties. Distribution of aerosol particles is generally defined by their number, surface or volume.

   The smallest range of these particles (with particle diameter is lower than 0.01 µm) are produced by homogeneous and heterogeneous nucleation processes. They can form during natural gas-to particle condensation or during condensation of hot vapour in combustion processes. Due to their rapid coagulation or random impaction onto surfaces, the lifetime of these small particles is very short (order of minutes to hours).

   Larger aerosol particles in the size range 0.1 to 1 µm in diameter can accumulate in the atmosphere because their removal mechanisms are least efficient. Their lifetime in the atmosphere is 7–10 days (10) are routinely monitored values and during this period they can transported to a long distance from their sources. Particles belonging to this accumulation mode are formed mainly by coagulation of smaller particles or condensation of vapours onto existing particles, and during the process they grow into this size range. At the same time, these are also being emitted to the atmosphere from different sources, mainly from incomplete combustion. Accumulated particles get removed from the atmosphere mainly by wet deposition. The Coarse mode contains particles with diameter larger than 2.5 μm. These particles mostly get emitted into the atmosphere during mechanical processes from both natural and anthropogenic sources (e.g. sea salt particles from ocean surface, soil and mineral dust, biological materials). Due to their relatively large mass, they have short atmospheric lifetimes because of their rapid sedimentation. Gravitational sedimentation is a state when a particle is released in air, the particle undergoing gravitational settling and it quickly reaches its terminal settling velocity, a condition of constant velocity wherein the drag force of the air on the particle, FD, is exactly equal and opposite to the force of gravity, FG. Under this condition, FD = FG = mg. The distribution of atmospheric aerosol particles can be seen in fig1. Small particles in nucleation mode constitute the majority of atmospheric particles by number. However due to their small sizes, their contribution to the total mass of aerosols are very small (around a few percent). The Accumulation mode particles have the greatest surface area. The mass or volume concentration is dominated by the aerosols in coarse and accumulation modes. Size, area and volume distributions of aerosol particles show characteristic pattern at different locations (e.g. urban, rural, remote continental or marine regions). On the basis of frequency distribution of mass or volume, the distribution may be characterized as two or three modal distribution.

   Study has been conducted to investigate the role of aerosol sedimentation, i.e. the downward motion and settling on the Earth’s surface of aerosol particles due to gravity. The impact of aerosol sedimentation and of its numerical implementation has been analyzed and incorporated in the global climate model.

   

   Condensation of aerosol particles

   Gas to particle conversion may result from homogenous gas phase processes or it may be controlled by processes in the particulate phase. Gas phase processes, either physical (include adiabatic expansion or mixing with cool air, radiative or conductive cooling) or chemical(i.e oxidation of SO2 to sulphuric acid ) can produce a super saturated state which then collapses by aerosol formation. Once a condensable species has been formed in the gas phase, the system is in a non equilibrium state. It may pass toward equilibrium by the generation of new particles (homogenous nucleation) or by condensation on existing particles (heterogenous condensation).

Coagulation of aerosol particles

Coagulation of aerosols, which is used to describe the growing process of aerosol particles in contact with each other, causes a continuous change in concentration and size distribution of agglomerates, keeping the total particle volume constant.

Structure:

Aerosol particles may occur by themselves or may be formed into chains of sphere. These are called agglomerates or flocs. Agglomerates are usually formed from highly charged small particles such as are found in metal fumes or dense smokes.

Surface properties

As we know aerosol particles have very small size and due to their small size, they provide large surface for chemical reactions such as adsorption , absorption , burning or other chemical reaction. The amount of area per gram of material increases as the particle size decreases, and for a given average size, increasing polydispersity decreases the surface area per gram.

Chemical speciation of Atmospheric Particulate Matter

Chemical speciation is essential for establishing more specific relationships between particle concentrations and measures of public health. Aerosol contaminants in the city air are extremely complex in chemical composition. More than 35 metallic elements have been found by chemical and spectroscopic analysis of the inorganic fractions. Some of the most abundant metallic elements are Si, Ca, Na, Al and Fe. Relatively high quantities of Mg, Pb, Cu, Zn and Mnare also found in the atmosphere. The concentration of these elements will depend on the nature of the area and the type of industries present. The largest contribution to coarse mass is typically from soil dust, which is dominated by salts of silicon, aluminum, iron and calcium. Other sources of coarse aerosol are sea spray. The crustal element Al, Si and Ca showed a close association. The distribution pattern for these crustal related elements suggests that crustal dust is contributing significantly not only to the coarse fraction but also to the fine fraction. Fe, Mn, Al, Mg and Ca (Calcium total) had high correlation with SPM because all of them originate from terrestrial sources. The elements like Zn, Cu, Cd, Pb, Ni and Cr had low correlations with SPM and may be said to have originated from anthropogenic source. Trace metals are found in almost all atmospheric aerosol size fractions. The presence of trace metals in concentrations exceeding critical thresholds may cause toxic effects within terrestrial ecosystems. In the absence of local industrial or mining activities, a major source of metals is usually deposition from the atmosphere. Accumulation mode particles (those possessing diameter around 0.1-1.0μm) deposit slowly and can therefore be transported over long distances having consequent effects in regions remote from the source.

Size differentiated particulate matter

Chemical composition of particulate matter along with size fraction plays important role in efficacy of monitoring work to know the extent PM will affect exposed public life. The particle size and chemical composition of inhalable particles are important in determining potential impacts on human health, since the efficiencies of both inhalation and respiratory deposition are dependent upon particle size while chemical composition mediates toxicity due to the presence of specific toxic elements and influence the non specific toxicity of particles. Various studies showed that heavy metals emissions from various anthropogenic sources, showed a tendency towards increasing concentration as particle size decreases. Explanation lies in facts that these metals volatilized during combustion and preferentially adsorbed or condensed through thermal coagulation on to small particles, which could most readily pass through conventional control equipment. Hence, particles having size <5μm are extremely harmful to human health, but more understanding is required for conclusion.

Atmospheric lifetime

The lifetime of atmospheric aerosol particles depends on their properties (size, chemical composition, etc.) and on altitude range, too (Figure 9.13). In the atmospheric boundary layer (lower troposphere), the residence time of aerosol particles is usually less than a week, often on the order of a day, depending on aerosol properties and meteorological conditions. In the free troposphere, the typical particle lifetime is 3–10 days on average. During this time, particle can easily be transported to a long distance. Therefore, there is a large variability in particle concentration, reflecting the geographical distribution of sources and sinks. The stratosphere also contains aerosol particles, which have much longer lifetime (up to 1 year), than in the tropospheric particles, due to the lack of precipitation. Smaller particles are efficiently removed by coagulation with other particles. Therefore, their lifetime is very short (in a range of ten minutes to day). Similarly, the large particles spend only a short time in the atmosphere due to the sedimentation. Particles in the accumulation mode have the longest lifetime (7–10 days on average), as in this range, both the Brownian diffusion and sedimentation are less important. These particles removed from the atmosphere predominantly by wet deposition.

Optical properties of Aerosol

Optical effects of aerosols represent the most spectacular and most important of all aerosol characteristics. The interaction of aerosols with radiation is usually measured by aerosols’ optical properties, e.g., the scattering coefficient, absorption coefficient, aerosol optical depth (AOD), single scattering albedo (SSA), and Angstrom exponent (AE). Aerosol particles directly affect the earth– atmosphere radiative balance through absorbing or scattering radiation and can also serve as cloud condensation nuclei to change global and regional climate. Therefore, aerosol optical properties and size distribution are major factors influencing the radiative balance, and knowledge of them is required to predict global climate change. The effects of aerosols on the climate and environment have been researched and discussed worldwide and showed that aerosol optical depth (AOD) and the Ångström exponent (α) are two basic optical parameters of aerosol particles. Recently scientists demonstrated that the single scattering albedo (SSA) is one of the most important factors indicating aerosol radiative effects. The Ångström exponent can also be used to study the size and growth of particles through its spectral dependence. The AOD is a parameter used to measure the magnitude of aerosol extinction due to scattering and absorption, integrated in the vertical column. It represents the e-folding length of the decrease in a direct beam when traveling through the aerosol layer. The SSA is a ratio of the scattering coefficient to the extinction coefficient and measures the relative importance of scattering and absorption. The aerosol effects on the radiation budget at the top of the atmosphere (TOA) switch from net cooling to warming at a certain value of the SSA, depending on the local surface albedo . The AE represents the wavelength dependence of AOD, with high values of AE indicating small particles and low values representing large particles.

Long-term monitoring of aerosols from the ground is required to obtain the temporal and spatial variation data necessary to understand the effects of aerosols on climate, regional climate and air quality. Aerosol optical properties  such as  AOD, α, SSA,  volumes  of aerosol size distributions, absorption AOD (AAOD) and direct radiative forcing using sun-photometer data in metropolitan cities have   been   studied.  Various   countries   have   established   the   remote   sensing   and  ground-based measurements networks for the recording and obtaining the data and thereby derive aerosol optical properties.  Knowledge  of an  atmospheric  particle’s  chemical  composition  is  of  importance  as  it determines the optical properties of particles and affects atmospheric chemistry in the gas as well as in the particulate phase. Moreover, the aerosol chemical composition influences the ability of particles to act as cloud condensation nuclei CCN   or ice nuclei (IN). For instance, while some particles (such as minerals) do not make very good CCN, they act as very good IN in colder parts of the atmosphere. Due to high inhomogeneous horizontal and vertical distributions of aerosols and impractical direct assessment of column leads to create the large uncertainties in estimating aerosol radiative effects.

Satellite remote sensing is the best, and indeed, the only way to observe aerosols on the global scale due to the short lifetime of aerosols and their complex chemical composition. The effectively determination of aerosol’s properties, some sensors such as Advanced Very High Resolution Radiometer (AVHRR)and Total Ozone Meteorological Satellite (TOMS) have been designed to monitor aerosols from space. Satellite remote sensing now includes new and enhanced sensors such as Polarization and Directionality of the Earth’s Reflectance, Moderate Resolution Imaging Spectro radiometer (MODIS), and Multi-angle Imaging Spectro Radiometer (MISR). The launch of satellite-borne lidars such as Geoscience Laser Altimeter System (GLAS) and Cloud Aerosol Lidar and Infrared Pathfinder Satellite Observation (CALIPSO)has further enabled profiling of vertical aerosol distribution. These sensors allow quantitative analysis of aerosol optical properties, especially AOD, and provide additional information regarding aerosol size, SSA, and refractive index. Such advanced sensors will also provide aerosol global distribution information, seasonal and inter annual variations in the sources, optical properties, and the direct and indirect effects of aerosols. Combining ground-based measurements and satellite remote sensing makes it possible to obtain relatively reliable information regarding global aerosol distribution.

Aerosol Optical Depth (AOD) or Aerosol Optical Thickness (AOT)

AOD (τ) is a measure of the extinction of the solar beam by aerosols. In other words, aerosols in the atmosphere can block sunlight by absorbing or by scattering light beam. AOD is defined as the integrated extinction coefficient over a vertical column of unit cross section. AOD is the dimensionless number. AOD just tell us how much the signal is attenuated due to the transmission along the path. Simply AOD tells how much direct sunlight prevented from reaching ground by these aerosols. Therefore, AOD is directly proportional to the aerosol concentration. AOD measurement instruments measures the aerosol extinction as a whole column, means between the Earth’s surface and the Top of Atmosphere (TOA). TOA is the outer most layer of Earth’s atmosphere; it extends from 700km a.s.l. to 10,000km a.s.l.

The larger the AOT at a particular wavelength, the less light of that wavelength reaches Earth’s surface. This information is important for determining the concentration, size distribution, and variability of aerosols in the atmosphere. An AOD of less than 0.1 indicates a crystal clear sky with maximum visibility, 0.4 would correspond to a very hazy condition, whereas a value of 4 indicates the presence of aerosols so dense that people would have difficulty seeing the Sun, even at mid-day.

Along with aerosols, other atmospheric constituents can also scatter light. That must be considered when calculating the AOD. The optical depth due to water vapor, Rayleigh scattering, and other wavelength-dependent trace gases must be subtracted from the total optical depth to obtain the AOD:

τ(λ)Aerosol = τ(λ)TOT – τ(λ)water – τ(λ)Rayleigh – τ(λ)O3 – τ(λ)NO2 – τ(λ)CO2 – τ(λ)CH4

Where, τ(λ)Aerosol is aerosol optical depth, τ(λ)TOT is total optical depth, τ(λ)water is optical depth due to water vapour, τ(λ)Rayleigh optical depth due to Rayleigh scattering, τ(λ)O3 is optical depth due to ozone, τ(λ)NO2 is optical depth due to NO2, τ(λ)CO2 is optical depth due to CO2 and τ(λ)CH4 is optical depth due to CH4.

Applications for AOD data:

1) Air Quality

2) Health and Environment

3) Earth Radiation Budget

4) Radiative Transfer Model

5) Atmospheric correction of remotely sensed surface features

6) Monitoring of sources and sinks of aerosols

7) Monitoring of forest fire and volcanic eruptions

8) Climate Change

Many techniques have been developed in order to monitor AOD.

A well established method consists of observing direct solar radiation using ground based sun photometer networks such as AERONET (Aerosol Robotic NETwork). This method gives good temporal AOD measurements but sparse spatial information because only sun photometer placed one hundred sites in all over the worlds.

A second important technique is based on inversion algorithms, which use dark targets, present in remotely sensed images. This technique has been successfully applied over dense dark vegetation (DDV) and ocean pixels using satellite based sensors such as MODIS (Moderate Resolution Imaging Spectro-radiometer), AVHRR (Advanced Very High Resolution Radiometer) and POLDER (Polarization and Directionality of the Earth’s Reflectance). Satellite based inversion techniques give more inclusive spatial information but are limited to a few images per day or week and are typically poorer in accuracy.

Radiative properties of aerosols:

Aerosols may influence the radiation budget of the earth-climate system. Profoundly two types of radiative effects have been suggested: direct, indirect.

Direct radiative effect(DRE) : The findings leading to the potential climate impact of aerosols has resulted in a large research effort that has significantly improved our understanding of the role of aerosols in the earth’s radiation balance. With a full set of aerosol optical properties available, the direct radiative effect of aerosols can be calculated by using satellite-based, model-based , satellite-model integrated, or ground-model integrated approaches.

Indirect radiative effect

Aerosol cloud interaction:

Clouds themselves are important regulators of the earth’s radiation budget. Overall, clouds cool the earth-atmosphere system at the TOA. Losses of 48 W m−2 at Top Of the Atmosphere (TOA) in the solar spectrum by clouds are only partially compensated for 30 W m−2 by cloud trapped infrared radiation. Aerosol particles serve as condensation nuclei for the formation of cloud droplets and atmospheric ice particles. Aerosol-cloud interaction includes the indirect radiative effect (IRE) and the semi-direct radiative effect. The IRE consists of two components: the albedo effect and the cloud lifetime effect. Each affects the size distribution and chemical nature of atmospheric aerosols as well as the chemical composition of clouds and precipitation.

Albedo effect: Because solar radiation is mainly scattered but only minimally absorbed by cloud droplets, an increase in cloud condensation nuclei (CCN) at constant liquid water content leads to a large concentration of small radius cloud droplets. This enhances cloud reflectivity, rendering the radiative forcing negative. This aspect of the aerosol indirect effect is referred to as the albedo effect.

Climatic Effects:

Aerosols play important role in the balance of the Earth’s climate. Due to the increasing anthropogenic emission of aerosols since the industrial revolution, they can also effect the global climate change. However, the effects of aerosols on climate are not one-way, moreover excessively uncertain. The climate forcing by aerosols can be realized in two ways, basically: in direct and indirect radiative forcing.

Direct effects: direct radiative forcing due the scattering radiation

Aerosol particles reflected a part of shortwave solar radiation back into the space, cooling the Earth’s atmosphere. This cooling effect of aerosols, especially by sulphate components may be compensated by the absorption of long wave terrestrial radiation primarily by elemental (black) carbon aerosols and dust particles.

Indirect effects: indirect radiative forcing through cloud formation effects:

Aerosol particles can also affect the radiation balance by formation of cloud droplets. Cloud droplets are formed in the troposphere by condensation of water vapour onto aerosol particles (cloud condensation nuclei, or ice nuclei) when the relative humidity exceeds the saturation level. Without these particles, a very large super-saturation (about 400%) would be necessary for the homogeneous condensation of water vapour. The properties and the number of particles can affect the formation and the characteristic of clouds and precipitation in many ways. The increased number of aerosol particles, and therefore the increased cloud optical thickness decrease the net surface solar radiation. The more numerous smaller cloud particles reflect more solar radiation called albedo effect. Smaller particles decrease the precipitation efficiency, thereby prolonging cloud lifetime. The absorption of solar radiation by soot particles may cause evaporation of cloud particles (semi-direct effect). In mixed-phased clouds, smaller cloud droplets delay the beginning of freezing and decrease the riming efficiency. However, more ice nuclei increase the precipitation efficiency.

Light scattering is one of two attenuating effects of aerosols on solar radiation (the other being absorption) and can be quantified directly by means of the total scattering coefficient.

Scattering coefficient:

Integrating nephelometers measure the total amount of light scattered by an aerosol. The `integration covers scattering angles from near forward to near backward.

Albedo:

Albedo is a non-dimensional, unitless quantity that indicates how well a surface reflects solar energy. Albedo varies between 0 and 1. Albedo commonly refers to the “whiteness” of a surface, with 0 meaning black and 1 meaning white. A value of 0 means the surface is a “perfect absorber” that absorbs all incoming energy. Absorbed solar energy can be used to heat the surface. A value of 1 means the surface is a “perfect reflector” that reflects all incoming energy. Albedo generally applies to visible light, although it may involve some of the infrared region of the electromagnetic spectrum. Sea ice has a much higher albedo compared to other earth surfaces, such as the surrounding ocean. A typical ocean albedo is approximately 0.06, while bare sea ice varies from approximately 0.5 to 0.7.

This means that the ocean reflects only 6 percent of the incoming solar radiation and absorbs the rest, while sea ice reflects 50 to 70 percent of the incoming energy. The sea ice absorbs less solar energy and keeps the surface cooler. Snow on glaciers or sea has an even higher albedo than sea ice, and so thick sea ice covered with snow reflects as much as 90 percent of the incoming solar radiation. This serves to insulate the sea ice, maintaining cold temperatures and delaying ice melt in the summer. After the snow does begin to melt, and because shallow melt ponds have an albedo of approximately 0.2 to 0.4, the surface albedo drops to about 0.75. As melt ponds grow and deepen, the surface albedo can drop to 0.15. As a result, melt ponds are associated with higher energy absorption and a more rapid ice melt.

Black carbon (BC) and its radiative forcing

Black carbon (BC) is a distinct type of carbonaceous material, formed only in flames during combustion of carbon-based fuels. It is distinguishable from other forms of carbon and carbon compounds contained in atmospheric aerosol because it has a unique combination of the following physical properties : (1) It strongly absorbs visible light with amass absorption cross section of at least 5 m2g-1 at a wavelength of 550 nm. (2) It is refractory; that is, it retains its basic format very high temperatures, with a vaporization temperature near 4000K. (3) It is insoluble in water, in organic solvents including methanol and acetone, and in other components of atmospheric aerosol. (4) It exists as an aggregate of small carbon spherules.

The strong absorption of visible light at all visible wavelengths by black carbon is the distinguishing characteristic that has raised interest in studies of atmospheric radiative transfer. No other substance with such strong light absorption per unit mass is present in the atmosphere in significant quantities.

A variety of combustion sources, both natural and anthropogenic, emits BC directly to the atmosphere. The largest global sources are open burning of forests and savannas, solid fuels burned for cooking and heating, and on-road and off-road diesel engines. Industrial activities are also significant sources, while aviation and shipping emissions represent minor contributions to emitted mass at the global scale.

The best quantified climate impact of BC is its atmospheric direct radiative forcing—the consequent changes in the radiative balance of the Earth due to an increase in absorption of sunlight within the atmosphere. Thus, it causes warming like green house gases. BC also produces warming when it is deposited on ice or snow because BC decreases the reflectivity of these surfaces, causing more solar radiation to be absorbed.

Visibility reduction

Atmospheric visibility is defined by the ability of our eyes to distinguish an object from the surrounding background. Scattering of solar radiation by aerosols is the main process limiting visibility in the troposphere. In the absence of aerosols our visual range would be more than 30 km, limited by scattering by air molecules. Anthropogenic aerosols in urban environments typically reduce visibility by one order of magnitude relative to unpolluted conditions. Degradation of visibility by anthropogenic aerosols is also a serious problem in our national parks and sanctuaries. The visibility reduction is greatest at high relative humidity when the aerosols swell by uptake of water increasing the cross-sectional area for scattering; this is the phenomenon known as haze.

References:

• John H. Seinfeld, Spyros N. Pandis., Atmospheric Chemistry and Physics: From Air Pollution to Climate Change, 3rd Edition (2016)., John Wiley & Sons
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• Parker C. Reist,Introduction to aerosol science.,1984. Macmillan Publishing Company, New York.
• Barbara Finlayson-Pitts, James Pitts, Jr.Chemistry of the Upper and Lower Atmosphere: Theory, Experiments and Applications., Academic press

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