Atmospheric Ozone

Ozone (03) is a reactive oxidant gas produced naturally in trace amounts in the Earth's atmosphere. Ozone was discovered by C. F. Schonbein in the middle of the nineteenth century; he also was first to detect ozone in air (Schonbein 1840, 1854). Schonbein (1840) suggested the presence of an atmospheric gas having a peculiar odor (the Greek word for "to smell" is ozeiri). Spectroscopic studies in the late nineteenth century showed that ozone is present at a higher mixing ratio in the upper atmospheric layers than close to the ground. Attempts to explain the chemical basis of existence of ozone in the upper atmosphere began nearly 80 years ago. Within the last 40 years, however, while increased understanding of the role of other trace atmospheric species in stratospheric ozone was unfolding, it became apparent that anthropogenically emitted substances have the potential to seriously deplete the natural levels of ozone in the stratosphere. At about the same period, ironically, it was realized that anthropogenic emissions could lead to ozone increases in the troposphere. Whereas stratospheric ozone is essential for screening of solar ultraviolet radiation, ozone at ground level can, at elevated concentrations, lead to respiratory effects in humans. This paradoxical dual role of ozone in the atmosphere has, on occasion, led to the dubbing of stratospheric ozone as "good" ozone and tropospheric ozone as "bad" ozone.

Most of the Earth's atmospheric ozone (about 90%) is found in the stratosphere where it plays a critical role in absorbing ultraviolet radiation emitted by the Sun. Figure 2.6 shows the stratospheric ozone at 35°N in September 1996. The peak in ozone molecular number density (concentration) occurs in the region of 20-30 km. The so-called stratospheric ozone layer absorbs virtually all of the solar ultraviolet radiation of wavelengths between 240 and 290 nm. Such radiation is harmful to unicellular organisms and to surface cells of higher plants and animals. In addition, ultraviolet radiation in the wavelength range 290-320 nm, so-called UV-B, is biologically active. A reduction in stratospheric ozone leads to increased levels of UV-B at the ground, which can lead to increased incidence of skin cancer in susceptible individuals. (An approximate rule of thumb is that a 1% decrease in stratospheric ozone leads to a 2% increase in UV-B.) The stratospheric temperature profile

Ozone Concentration Ozone Mixing Ratio

Ozone Concentration Ozone Mixing Ratio

Ozone Mixing Ratio
FIGURE 2.6 Stratospheric ozone profile over Northern Hemisphere midlatitude (35°N) in September 1996 as measured by satellite with the Jet Propulsion Laboratory FTIR (Fourier transform infrared) spectrometer. Note that molecular concentration and mixing ratio peak at different altitudes.

(see Figure 1.1) is the result of ozone absorption of radiation. Stratospheric chemistry centers around the chemical processes influencing the abundance of ozone.

As compared with the stratosphere, natural concentrations of tropospheric ozone are small—usually a few tens of parts per billion (ppb) in mixing ratio (molecules of 03/molecules of air; 10 ppb — 2.5 x 10'1 molecules cm"" at sea level and 298 K) versus peak stratospheric mixing ratios of more than 10,000 ppb (10 ppm). Since the atmospheric molecular number density thins out exponentially with altitude, Lhe peak in ozone mixing ratio occurs at a higher altitude than does its peak in concentration. Still, a significant amount of naturally occurring ozone, about 10-15% of the atmospheric total, is found in the troposphere (Fishman et al. 1990). The total amount of O3 in the atmosphere, stratosphere and troposphere combined, is extremely small. In the pristine, unpolluted troposphere ozone mixing ratios are in the range of 10-40 ppb with somewhat higher mixing ratios in the upper troposphere. Ozone reaches a maximum mixing ratio of about 10 ppm at an altitude of 25-30 km in the stratosphere.

Over the past three decades, the improved ability to monitor atmospheric ozone with ground-siled. aircraft-mounted, and satelliteborne instruments has led to definitive data about changes in the amounts of ozone found in the atmosphere. One of the most significant environmental issues facing the planet is the catalytic destruction of significant portions of the stratospheric ozone layer, most dramatically seen each Antarctic spring, which is the result of stratospheric halogen-induced photochemistry (see Chapter 5). The current and projected loss of stratospheric ozone has stimulated a major, continuing global research program in stratospheric chemistry and an international treaty that restricts the release of chlorofluorocarbons into the atmosphere.

The Dobson Unit Imagine that all the ozone in a vertical column of air reaching from the Earth's surface to the top of the atmosphere were concentrated in a single layer of pure 03 at the surface of the Earth, at 273 K and 1.013 x 10! Pa. The thickness of that layer, measured in hundredths of a millimeter, is the column abundance of Oi expressed in Dobson units (DU), The unit is named after G. M. B. Dobson, who, in 1923, produced the first ozone spectrometer, the standard instrument used to measure ozone from the ground. The Dobson spectrometer measures the intensity of solar UV radiation at four wavelengths, two of which are absorbed by ozone and two of which are not.

One DU is equivalent to 2.69 x 1016 molecules 03 cm 2. A normal value of 03 column abundance over the globe is about 300 DU, corresponding to a layer of pure 03 at the surface only 3 mm thick.

If the ozone concentration as a function of altitude is «q3 (z) (molecules cm-3), the 03 column burden is roo

To find what this «q, translates to in terms of DU, we need to determine the thickness of the layer of pure 03 at 273 K and 1 atm corresponding to this burden. Over 1 cm2 of area, this corresponds to molecules of 03. The volume (cm3) occupied by this number of molecules of 03 can be written as V = 1 (cm2) x /¡(cm). One DU corresponds to a thickness, h, of 0.01 mm (0.001 cm).

From the ideal-gas law

where «Oj/^Va is the number of moles of Ch, and the factor of 106 is needed to convert m to cm3. At 273 K and ! atmosphere, the thickness of the layer of pure 0-, is

For h — 0.01 mm = 0.00! cm (! DU), the column burden of is

Whereas stratospheric ozone is thinning, tropospheric ozone is increasing (WMO 1986, 1990). Much of the evidence for increased baseline levels of tropospheric ozone comes from Europe, where during the late 1800s there was much interest in atmospheric ozone. Because ozone was known to be a disinfectant, it was believed to promote health (Warneck 1988). Measurements of atmospheric ozone made at Montsouris, near Paris, from 1876 to 1910 have been reanalyzed by Vo!z and Kley (1988), who recalibrated the original measurement technique. Their analysis showed that surface ozone mixing ratios near Paris over 100 years ago averaged about lOppb; current mixing ratios in the most unpolluted parts of Europe average between 20 and 45 ppb (Volz and Kley 1988; Bojkov 5988;Crutzen 1988;Staehelin and Schmid 199i;Oltmans and Levy 1994;Janach 1989). An analysis of ozone measurements made in relatively remote European sites indicates a 1 to 2% annual increase in average concentrations over the past 40 years (Janach 1989).' Based on total ozone mapping spectrometer (TOMS) data between the high Andes and the Pacific Ocean, from 1979 to 1992 tropospheric ozone concentrations in the tropical Pacific South America apparently increased by i,48% per year (Jiang and Yung 1996). The integrated ozone column in the troposphere can be determined from the difference of the measurements of two satellite instruments, TOMS and SAGE, which detect total column ozone and stratospheric ozone, respectively (Fishman et al. 1990, 1992). Tropospheric ozone column densities average about 30 DU, but there is significant variation with season and hemisphere.


Particles in the atmosphere arise from natural sources, such as windborne dust, seas pray, and volcanoes, and from anthropogenic activities, such as combustion of fuels. Whereas an aerosol is technically defined as a suspension of fine solid or liquid panicics in a gas,

1 If stratospheric ozone concern rations remained constant, the 10% increase in tropospheric ozone would increase the total column abundance of ozone by aboui 1%. Thus the additional tropospheric ozone is believed to have counteracted only a small fraction of the stratospheric loss, even if the trends observed over Europe are representative of the entire northern midlatitude region.

TABLE 2.18 Terminology Relating to Atmospheric Particles

Aerosols, aerocolloids, aerodisperse systems Dusts


Hazes Mists





Tiny particles dispersed in gases

Suspensions of solid particles produced by mechanical disintegration of material such as crushing, grinding, and blasting; Dp > 1 |_im A term loosely applied to visible aerosols in which the dispersed phase is liquid; usually, a dispersion of water or ice, close to the ground The solid particles generated by condensation from the vapor state, generally after volatilization from melted substances, and often accompanied by a chemical reaction such as oxidation; often the material involved is noxious; Dp < 1 |im An aerosol that impedes vision and may consist of a combination of water droplets, pollutants, and dust; Dp < 1 |xm Liquid, usually water in the form of particles suspended in the atmosphere at or near the surface of the Earth; small water droplets floating or falling, approaching the form of rain, and sometimes distinguished from fog as being more transparent or as having particles perceptibly moving downward; Dp > 1 nm An aerosol particle may consist of a single continuous unit of solid or liquid containing many molecules held together by intermolecular forces and primarily larger than molecular dimensions (>0.001 fim); a particle may also consist of two or more such unit structures held together by interparticle adhesive forces such that it behaves as a single unit in suspension or on deposit A term derived from smoke and fog, applied to extensive contamination by aerosols; now sometimes used loosely for any contamination of the air

Small gasbome particles resulting from incomplete combustion, consisting predominantly of carbon and other combustible materials, and present in sufficient quantity to be observable independently of the presence of other solids. Dp > 0.01 |im Agglomerations of particles of carbon impregnated with "tar," formed in the incomplete combustion of carbonaceous material common usage refers to the aerosol as the particulate component only (Table 2.18). Emitted directly as particles (primary aerosol) or formed in the atmosphere by gas-to-particle conversion processes (secondary aerosol), atmospheric aerosols are generally considered to be the particles that range in size from a few nanometers (nm) to tens of micrometers (|im) in diameter. Once airborne, particles can change their size and composition by condensation of vapor species or by evaporation, by coagulating with other particles, by chemical reaction, or by activation in the presence of water supersaturation to become fog and cloud droplets. Particles smaller than 1 |im diameter generally have atmospheric concentrations in the range from around ten to several thousand per cm3; those exceeding 1 |im diameter are usually found at concentrations less than 1 cm-3.

Particles are eventually removed from the atmosphere by two mechanisms: deposition at the Earth's surface (dry deposition) and incorporation into cloud droplets during the formation of precipitation (wet deposition). Because wet and dry deposition lead to relatively short residence times in the troposphere, and because the geographic distribution of particle sources is highly nonuniform, tropospheric aerosols vary widely in concentration and composition over the Earth. Whereas atmospheric trace gases have lifetimes ranging from less than a second to a century or more, residence times of particles in the troposphere vary only from a few days to a few weeks.

2.7.1 Stratospheric Aerosol

The stratospheric aerosol is composed of an aqueous sulfuric acid solution of 60-80% sulfuric acid for temperatures from - 80 to - 45°C, respectively (Shen et al. 1995). The source of the globally distributed, unperturbed background stratospheric aerosol is oxidation of carbonyl sulfide (OCS), which has its sources at the Earth's surface. OCS is chemically inert and water insoluble and has a long tropospheric lifetime. It diffuses into the stratosphere where it dissociates by solar ultraviolet radiation to eventually form sulfuric acid, the primary component of the natural stratospheric aerosol. Other surface-emitted sulfur-containing species, for example, S02, DMS, and CS2, do not persist long enough in the troposphere to be transported to the stratosphere.

A state of unperturbed background stratospheric aerosol may be relatively rare, however, as frequent volcanic eruptions inject significant quantities of S02 directly into the lower and midstratosphere. Major eruptions include Agung in 1963, EI Chichon in 1982, and Pinatubo in 1991. The subsequent sulfuric acid aerosol clouds can, over a period of months, be distributed globally at optical densities that overwhelm the natural background aerosol. The stratosphere's relaxation to background conditions has a characteristic time on the order of years, so that, given the frequency of volcanic eruptions, the stratospheric aerosol is seldom in a state that is totally unperturbed by volcanic emissions. With an estimated aerosol mass addition of 30 Tg to the stratosphere, the June 1991 eruption of Mt. Pinatubo was the largest in the 20th century and led to enhanced stratospheric aerosol levels for over 2 years.

2.7.2 Chemical Components of Tropospheric Aerosol

A significant fraction of the tropospheric aerosol is anthropogenic in origin. Tropospheric aerosols contain sulfate, ammonium, nitrate, sodium, chloride, trace metals, carbonaceous material, crustal elements, and water. The carbonaceous fraction of the aerosols consists of both elemental and organic carbon. Elemental carbon, also called black carbon, graphitic carbon, or soot, is emitted directly into the atmosphere, predominantly from combustion processes. Particulate organic carbon is emitted directly by sources or can result from atmospheric condensation of low-volatility organic gases. Anthropogenic emissions leading to atmospheric aerosol have increased dramatically over the past century and have been implicated in human health effects (Dockery et al. 1993), in visibility reduction in urban and regional areas (see Chapter 15), in acid deposition (see Chapter 20), and in perturbing the Earth's radiation balance (see Chapter 24).

Table 2.19 presents data summarized by Heintzenberg (1989) and Solomon et al. (1989) on aerosol mass concentrations and composition in different regions of the troposphere. It is interesting to note that average total fine particle mass (that associated with particles of diameter less than about 2 pm) in nonurban continental, (i.e., regional) aerosols is only a factor of 2 lower than urban values. This reflects the relatively long residence time of particles. Correspondingly, the average compositions of nonurban

TABLE 2.19 Mass Concentrations and Composition of Tropospheric Aerosols

Percentage Composition

Percentage Composition

TABLE 2.19 Mass Concentrations and Composition of Tropospheric Aerosols


Mass (ng m 3)

C (elem)

C (org)




Remote (11 areas)"







Nonurban continental







(14 areas)*

Urban (19 areas)"







Rubidoux, California*







(1986 annual average)

"Heintzenberg( 1989). 6Solomon et al. (1989).

(1986 annual average)

"Heintzenberg( 1989). 6Solomon et al. (1989).

continental and urban aerosols are roughly the same. The average mass concentration of remote aerosols is a factor of 3 lower than that of nonurban continental aerosols. The elemental carbon component, a direct indicator of anthropogenic combustion sources, drops to 0.3% in the remote aerosols, but sulfate is still a major component. This is attributable to a global average concentration of nonseasalt sulfate of about 0.5 pg m 3. Rubidoux, California, located about 100 km east of downtown Los Angeles, routinely experiences some of the highest particulate matter concentrations in the United States.

2.7.3 Cloud Condensation Nuclei (CCN)

Aerosols are essential to the atmosphere as we know it; if the Earth's atmosphere were totally devoid of particles, clouds could not form. Particles that can become activated to grow to fog or cloud droplets in the presence of a supersaturation of water vapor are termed cloud condensation nuclei (CCN). At a given mass of soluble material in the particle there is a critical value of the ambient water vapor supersaturation below which the particle exists in a stable state and above which it spontaneously grows to become a cloud droplet of 10 pm or more diameter. The number of particles from a given aerosol population that can act as CCN is thus a function of the water supersaturation. For marine stratiform clouds, for which supersaturations are in the range of 0.1-0.5%, the minimum CCN particle diameter is 0.05-0.14 pm. CCN number concentrations vary from fewer than 100 cm-3 in remote marine regions to many thousand cm-3 in polluted urban areas. An air parcel will spend, on average, a few hours in a cloud followed by a few days outside clouds. The average lifetime of a CCN is about 1 week, so that an average CCN will experience 5-10 cloud activation/cloud evaporation cycles before actually being removed from the atmosphere in precipitation.

2.7.4 Sizes of Atmospheric Particles

Atmospheric aerosols consist of particles ranging in size from a few tens of angstroms (A) to several hundred micrometers. Particles less than 2.5 pm in diameter are generally referred to as "fine" and those greater than 2.5 um diameter as "coarse." The fine and coarse particle modes, in general, originate separately, are transformed separately, are removed from the atmosphere by different mechanisms, require different techniques for their removal from sources, have different chemical composition, have different optical properties, and differ significantly in their deposition patterns in the respiratory tract. Therefore the distinction between fine and coarse particles is a fundamental one in any discussion of the physics, chemistry, measurement, or health effects of aerosols.

The phenomena that influence particle sizes are shown in an idealized schematic in Figure 2.7, which depicts the typical distribution of surface area of an atmospheric aerosol. Particles can often be divided roughly into modes. The nucleation (or nuclei) mode comprises particles with diameters up to about 10 nm. The Aitken mode spans the size range from about lOnm to lOOnm (0.1 pm) diameter. These two modes account for the

Atmospheric Aerosol Processes

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  • Satu
    Why do molecular number concentration and mixing ratio peak at different altitudes ozone?
    9 years ago
  • Bradley
    How are mixing ratios different at different altitudes?
    8 years ago

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