Acid Deposition

An "acid" is essentially any subslance that releases hydrogen ions (H+) when dissolved in water. Several atmospheric trace constituents dissociate into positive and negative ions when dissolved in water, and some are acidic to varying degrees. The strongest acids in atmospheric waters are dissolved sulfuric acid (H2S04) and nitric acid (HN03). Numerous other acidic substances have been identified in the atmosphere, such as sulfur dioxide (S02), organic acids, hydrochloric acid (HCl), carbon dioxide (C02), and even water itself, but these substances are either relatively weak acids or are present at relatively small concentrations and thus do not usually contribute appreciably to measured acidity. For any solution, there is an equal concentration of dissolved positive and negative ion charge, and a typical ion balance in cloudwater and precipitation is

[H+] + [Na+] + [NH4+] + [soil ions] = (positive ions)

= 2[S04=] + [NO,-] + [Cr] + [HCOj-] (negative ions)

Na~" and Cl~ ions arise from dissolved sea salt aerosols and are usually present in approximately equal concentrations. NH4" is dissolved ammonia, "soil ions" refers to calcium and magnesium cations that are typically associated with carbonates

(HC03~) in soil dust, and S04= and HN03~ are sulfuric and nitric acid. Since an ion balance is always maintained in atmospheric waters, the concentration of H+ is

[H+] = 2[S04=] + [N<V] + [HCO3-] - [NH4+] - [soil ions]

Thus the concentration of the hydrogen ion is proportional to the concentrations of sulfates, nitrates, bicarbonate (dissolved C02), ammonia, and carbonate-laden soil dust dissolved in cloudwater and precipitation. Measured concentrations of H+ vary over several orders of magnitude, and therefore a logarithmic pH scale is used to quantify acidity levels in water pH = -log10[H+]

Using the pH scale, a decrease in one pH unit corresponds to a 10-fold increase in acidity or H+ concentration. Also, as the pH decreases, the concentration of H+ and acidity increases. Pure water in equilibrium with atmospheric C02 has a pH near 5.6, but the concentrations of sulfate, nitrates, ammonia, or soil cations in cloud and rainwater usually greatly exceed the concentrations of dissolved C02, even in remote areas. Typical "clean" atmospheric waters have a pH of 4.5 to 5.5. In more polluted areas, pHs in precipitation range from 3 to 4, and in some low liquid water content clouds, pHs as low as 2 to 3 have been measured.

Formation Of Acids. Sulfuric and nitric acids are produced by reactions between atmospheric oxidants and emitted sulfur and nitrogen oxides (S02 and NOt ), which are by-products of fossil fuel combustion and other industrial activities. The dominant reactions converting S02 to sulfuric acid include reactions with hydrogen peroxide (H202) in clouds and hydroxyl radical (HO) in air. Nitric acid is produced by the oxidation of N02 by the HO radical, and also at night by a heterogeneous reaction involving ozone, N02, and N03 radicals. Atmospheric oxidants responsible for acid formation are produced via a complex sequence of photochemical reactions, and some acid-generating chemical reactions occur among dissolved gaseous constituents in atmospheric clouds or aerosols. Typically, emitted NO-,, is converted to nitric acid within a day or less, and S02 is converted to sulfuric acid within several days following emission. The concentrations of the oxidants, and the time scale for chemical reactions vary strongly with season, latitude, time of day, sunlight intensity, background concentrations of NOj and organic compounds, and many other chemical and meteorological factors.

Strong acids have an affinity for water, and therefore hygroscopically grow or combine with water vapor to form "haze" aerosols containing sulfuric acid, nitric acid, and varying degrees of neutralizing ammonia (NH3), especially when atmospheric relative humidities are above 60 to 70%. Typically, ammonia and nitric acid are present as both gases and aerosols in the atmosphere, while sulfate partitions predominantly into condensed aerosols. These sulfate, nitrate, and ammonium-containing aerosol particles constitute a significant fraction of cloud condensation nuclei (CCN), and thus acid-containing aerosols are readily incorporated into clouds.

Precipitation forming within clouds therefore contains dissolved CCN together with other soluble gases such as HN03 and NH3.

Undesirable Effects. At high concentrations or exposures, acidic solutions induce numerous undesirable reactions with surfaces. In conjunction with other pollutants, acid deposition contributes to potentially deleterious effects on aquatic, agricultural, and forest ecosystems. Chemical changes attributed to the deposition of acidity from the atmosphere have been measured in forest ecosystems and surface waters.

Concentrations of acids in lakes have been correlated with the concentrations and deposition rates of atmospheric acids, and high concentrations of acids in lakes and streams can adversely affect fish populations. Health effects associated with exposure to acid-containing particulates in humans remains an area of uncertainty since current studies of these effects are too limited to unambiguously discern dose-response relationships in humans. Acid deposition from the atmosphere has been shown to accelerate the deterioration rate of exposed metals, painted finishes, and concrete or stone surfaces.

In industrialized areas, concentrations of sulfuric and nitric acids in cloud water and precipitation are up to 50 to 100 times greater than values measured in areas that are not influenced by upwind emissions of anthropogenic pollutants. The relative concentrations of deposited sulfur and nitrogen acids are correlated with the relative emission rates of sulfur and nitrogen pollutants over larger areas.

The amount of acidity within precipitation is strongly influenced by numerous meteorological and chemical factors, as well as the emission rates of precursor sulfur and nitrogen pollutants. Therefore a thorough understanding of larger-scale meteorology, cloud dynamics and microphysics, and atmospheric chemistry is required to fully quantify and study atmospheric acidity

REFERENCES

Köhler H., Zur thermodynamic der kondensation an hygroskopichen kernen und bemerkungen

über das zusammenfliessen der tropfen, Medd. Met. Hydr. Anst. Stockholm, 3(8), 1926. Lipps, F. B., and R. S. Hemler, Numerical simulation of deep tropical convection associated with large-scale convergence, J. Atmos. Sei., 43, 1796-1816, 1986. Pruppacher, H. R., and J. D. Klett, Microphysics of Clouds and Precipitation, D. Reidel

Publishing, 714 pp., 1978. Twomey, S., The nuclei of natural cloud formation: The supersaturation in natural clouds and the variation of cloud droplet concentration, Geophys. Pura Appi, 43, 243-249, 1959. Warner, J., J. Atmos. Sei., 27, 1035, 1970.

Weisman, M. L., and J. B. Klemp, The dependence of numerically simulated convective storms on vertical wind shear and buoyancy, Monthly Weather Rev., 110, 504-520, 1982.

dional components, respectively, of the horizontal velocity, Vh. The variable, oj is defined as

Dt and becomes the surrogate for the vertical velocity, w, in the (x,y,p) coordinate system.

For large-scale motions, the vertical component of velocity is typically several orders of magnitude smaller than for the horizontal velocity, and vertical accelerations can be neglected to a good approximation. The vertical component of the momentum equation then reduces to a diagnostic equation and can be expressed as

dp p

with R the gas constant for air and T the temperature.

The mass continuity equation in this coordinate system becomes

Finally, the thermodynamic energy equation becomes

with Cp the specific heat at constant pressure for air. The variable Q is the diabatic heating rate. For the stratosphere, Q is typically the radiative heating rate and can be calculated as a function of the other dependent variables, knowing the distribution of radiatvely active species such as ozone, water vapor, and carbon dioxide (Goody, 1995).

Equations (1) to (4) represent the primitive equations in the (x,y,p) coordinate frame and form a determinate system of equations for the variables u, v, to, <D and T. These equations are inherently nonlinear and must be integrated in time with suitable initial and boundary conditions. Direct solution of the primitive equations requires the use of sophisticated numerical techniques implemented on digital computers.

A relationship of fundamental importance can be derived from the primitive equations [see Pedlosky (1979) for the derivation] in terms of Ertel's potential vorticity, n, namely for frictionless, adiabatic conditions. Here, II, is given by n = -(£+/) V©

Note that in Eq. (6) the total derivative is now expressed in the Cartesian coordinate system (x, y, z) as

Also, V in Eq. (6) is now the three-dimensional gradient operator in the (x, y, z) coordinate system with p the density, t the relative vorticity (curl of the velocity V), and © the potential temperature. For frictionless, adiabatic flow, this relationship requires that II be conserved following the motion. Physically, potential vorticity is representative of the ratio of the rotation of a fluid vortex column to the depth of the column. Conservation of potential vorticity for a compressible fluid can be thought of as analogous to conservation of angular momentum for a solid body. This conservation property for II represents a very powerful constraint on the motions, particularly in the lower stratosphere where Eq. (5) is a reasonable approximation for time scales up to about 10 days. Distributions of n on an isentropic surface (i.e., a constant potential temperature surface) thus represent a conserved dynamical tracer and, as we shall later see, provide much useful insight into the nature of the transport.

2 VERTICAL TEMPERATURE STRUCTURE

A traditional means of separating Earth's atmosphere into regions with quite distinct characteristics originated from considerations of the vertical temperature structure. A reference temperature profile for middle latitudes is shown in Figure 1 as a function of the geometric height using data from the compilation, U. S. Standard Atmosphere (1976). The difference between the stratosphere and troposphere in terms of the vertical temperature profile is readily apparent. Note that in the troposphere, the temperature T, decreases with increasing height z, up to the level of the tropopause. In contrast, the temperature in the lower stratosphere is nearly constant and then increases with increasing height through the middle and upper stratosphere until the level of the stratopause is reached. It is this difference in the temperature lapse rate, r, where that results in a difference in the stability characteristics between these two regions of the atmosphere.

The atmosphere is said to be statically stable, if after an air parcel is adiabatically displaced in the vertical dimension, the net forces acting on the parcel tend to restore

than is typical for the troposphere, where air parcels can be transported vertically between the surface and the tropopause on time scales of several days or even as little as a few minutes in the case of very strong convective activity associated with the largest thunderstorms (Wallace and Hobbs, 1977).

This stable stratification with relatively slow vertical mixing is an important influence on the stratospheric circulation and gives rise to the long residence times in the stratosphere that have been deduced for long-lived constituents such as the chlorofluorocarbons (CFC) [appproximately 50 years for CFC-11 and 100 years for CFC-12, the two most abundant CFC compounds (WMO, 1994)]. Here, long-lived is used in the sense that the characteristic time scale for chemical change is very much greater than that associated with the transport of a constituent.

3 ZONAL-MEAN CLIMATOLOGY OF TEMPERATURES AND ZONAL WINDS

Climatologies of zonal-mean (spatial averages at constant latitude) values of temperatures and zonal (east-west component) winds are displayed as a function of height and latitude in Figure 2a and 2b, respectively. An inspection of the temperature cross section shown in Figure 2a reveals marked variations in temperatures between the winter and summer hemispheres in the stratosphere. Note that the tropopause above the equator occurs at a much higher altitude than above the polar regions and that distinct discontinuities or "breaks" occur in the tropopause at middle latitudes. In the lower, summer stratosphere the temperature increases from the equator toward the pole. In contrast, there is a midlatitude warm belt (Rama-nathan and Grose, 1978) in the lower, winter stratosphere. Coldest temperatures occur in the lower stratosphere over the south polar region during winter. The northern polar regions exhibit more dynamical variability during winter and are typically not quite as cold. In the upper stratosphere the temperature increases monotonically from winter to summer pole.

The corresponding cross section of zonal winds is shown in Figure 2b. Note the presence of easterly winds (from east toward the west) in the summer hemisphere and westerly winds (from west toward the east) in the stratosphere. The seasonal reversal of summer easterlies to winter westerlies in each hemisphere is observed to be a prominent feature of the stratospheric circulation. The axis of the wintertime, westerly jetstream tilts poleward with decreasing height in the stratosphere resulting in a high latitude jetstream in the lower stratosphere, the so-called polar night jet.

4 ZONAL-MEAN MERIDIONAL CIRCULATION

The zonally averaged, meridional (north-south component) and vertical winds are, on average, at least an order of magnitude smaller than the zonal winds described in the previous section. A useful depiction of the circulation in the meridional plane (latitude vs. height), was originally derived by Murgatroyd and Singleton (1961).

Polar Hydrology

Figure 3 Schematic streamlines of the diabatic circulation for solstice conditions. S, the summer pole and W the winter pole (from Dunkerton, 1978).

km 70

60 30 EO

30 60

Figure 3 Schematic streamlines of the diabatic circulation for solstice conditions. S, the summer pole and W the winter pole (from Dunkerton, 1978).

This circulation is now generally referred to as the diabatic circulation. It is conceptually useful in that it illustrates the sense of the actual mass motions in the meridional plane. Using net radiative heating rates originally compiled by Murgatroyd and Singleton for solstice conditions, Dunkerton (1978) derived the vertical velocities necessary to balance these heating rates. From considerations of mass continuity, the corresponding meridional velocities were then calculated. Figure 3 shows the resultant streamlines inferred by Dunkerton (1978) from the calculated velocity fields. The large-scale stratospheric circulation in the meridional plane exhibits rising motion in the summer hemisphere with a slow (seasonal or longer time scale) meridional drift and subsidence over the winter pole (Andrews et al. 1987). A Coriolis torque associated with the meridional drift influences the production of the summertime (wintertime) zonal-mean easterlies (westerlies) in the stratosphere seen in Figure 2b.

5 WAVE MOTIONS

The depiction of the large-scale stratospheric circulation discussed in the preceding two sections is a zonally averaged perspective as noted. However, wavelike disturbances (commonly referred to as "waves"), which propagate vertically from the troposphere producing departures from zonal symmetry, are known to be important in determining the circulation and the transport of constituents in the stratosphere.

Important departures from the climatological zonal mean state shown in Figure 2a and 2b occur as a result of sudden stratospheric warmings, the quasibiennial oscillation (QBO) and the semiannual oscillation (SAO). Occurrence of these latter two phenomena is presently believed to result at least in part from vertically propagating Kelvin, Rossby-gravity, and/or gravity waves. The QBO manifests itself as an oscillation of the zonal winds with alternating westerlies and easterlies in the equatorial lower stratosphere. The period of the oscillation varies, but averages about 27 months. The SAO is also an oscillation of equatorial zonal winds with alternating westerlies and easterlies, but this phenomenon occurs in the upper stratosphere (and lower mesosphere) on a semiannual basis as the name implies. The mechanisms responsible for the QBO and SAO are quite different. The interested reader is referred to Andrews et al. (1987) for further details.

The sudden stratospheric warming phenomenon occurs during some years in association with anomalous enhancement of the amplitude of vertically propagating, planetary-scale disturbances into the wintertime stratosphere. Major warming events are characterized by very rapid increases in polar temperatures (50 to 70 K in a week or less) and severe disruptions of the wintertime westerly polar vortex with zonal easterlies replacing zonal westerlies at high latitudes. These warming events can occur in either hemisphere, although generally Southern Hemisphere warming events tend to be somewhat less spectacular than their counterpart in the Northern Hemisphere.

The various different types of waves [e. g., gravity waves, Rossby waves, Kelvin waves, and Rossby-gravity waves; see Andrews et al. (1987), for further description of these and other types of waves] are often distinguished by the restoring mechanism that is responsible for producing the wavelike motions. In particular, both gravity waves and Rossby waves play a most significant role with respect to the large-scale motions and the concomitant transport of constituents observed in the stratosphere.

The restoring force for the gravity wave is the buoyancy force, which is proportional to N2 and results from stable density stratification in the atmosphere. One of the most important consequences of gravity waves propagating upward into the stratosphere is their role in determining the structure of the stratospheric jets. Gravity waves propagate upward and "break" (Lindzen, 1981) at some level in the stratosphere or in the mesosphere. Breaking occurs as the gravity wave grows in amplitude, eventually producing an unstable lapse rate with resultant turbulence and mixing. Momentum deposition occurs as a result of the breaking process and effectively creates a net drag on the zonal momentum budget and decelerates the zonal-mean flow. Gravity wave breaking in the lower stratosphere is believed to contribute to the separation of the tropospheric and stratospheric jets as seen in Figure 2b. In a similar fashion, gravity wave breaking in the mesosphere contributes to deceleration and closing off above the stratospheric jets in the mesosphere.

The Rossby wave restoring force ultimately results from the latitudinal gradient in the Coriolis parameter, /,

where Q is the angular velocity of rotation of Earth and <fi is the latitude. The Rossby wave is responsible for much of the irreversible quasi-horizontal transport that occurs in the extratropical wintertime stratosphere by a process termed Rossby wave breaking (Mclntyre and Palmer, 1984), which is fundamentally different from the gravity wave breaking process just discussed. Contours of constant values of Ertel's potential vorticity, II, on an isentropic surface (a constant 0 surface) represent material lines when Eq. (5) is valid. Rossby wave breaking occurs as the wave amplifies, and material lines buckle and deform irreversibly. A graphic illustration of the process can be seen in Figures 4a and 4b. This sequence of figures depicts conditions in the midstratosphere (about 30 km) during a stratospheric warming in January 1979. The shaded area in Figure 4a correlates very well with the polar vortex, which was nearly centered on the pole on January 17, 1979. Ten days later (Fig. 4b), a large-amplitude wave resulted in elongation of the polar vortex with a tongue of low potential vorticity air from the vortex being drawn around the Aleutian anticyclone (which is centered near 65°N, 150°W at this time), and discrete parts of this tongue are seen to be scattered around the anticyclone. Mclntyre and Palmer (1983) constructed these isentropic distributions of II from meteorological analysis data and refer to them as a "coarse-grained" view of the potential vorticity distribution in the stratosphere.

Concurrent observations of ozone (Leovy et al., 1985) show a very high degree of correlation with the potential vorticity distribution and testify to the essential correctness of the Rossby wave-breaking paradigm of Mclntyre and Palmer (1983) for transport and mixing in the extra-tropical stratosphere.

6 SUMMARY

Earth's atmosphere is a thin layer of gas rotating with the planet and externally forced by differential radiative heating by the Sun. The somewhat simplified picture of the large-scale circulation in the stratosphere that has been presented here has been separated into two component parts for convenience, a zonally averaged representation and a contribution from wavelike disturbances. The zonally averaged circulation described here is one that is thermally driven with air parcels heated and ascending at low latitudes, drifting slowly poleward, and finally cooling and descending at higher latitudes. Concurrently, wavelike disturbances propagating upward produce departures from the zonally averaged state. These disturbances transport and mix heat, momentum, and constituents. The reader should be reminded that chemical transformations are also taking place that not only alter the composition of the stratosphere, but also affect the radiative heating as the composition changes (principally changes in ozone, water vapor, and carbon dioxide). It should also be noted that the process by which air enters and leaves the stratosphere (the stratospheric-tropospheric exchange process) is considerably more complicated than described here. A recent comprehensive review of stratospheric-tropospheric exchange can be found in Hoi ton et al. (1995). The stratosphere should be viewed as a very complex system in which radiative, chemical, and dynamical processes mutually interact to determine the structure and composition.

Figure 4 Potential vorticity distribution on the 850K isentropic surface for (a) January 17, 1979 and (b) January 27, 1979. Polar stereographic projection with outermost latitude circle 20°N (from Mclntyre and Palmer, 1983).

REFERENCES

Andrews, D. G., J. R. Holton, and C. B. Leovy, Middle Atmosphere Dynamics, Academic, London, 1987.

Dunkerton, T. J., On the mean meridional mass motions of the stratosphere and mesosphere, J. Atmos. Sei., 35, 600-614, 1978.

Goody, R., Principles of Atmospheric Physics and Chemistry. Oxford University Press, New York, 1995.

Holton, J. R., An Introduction to Dynamic Meteorology, Academic, San Diego, CA, 1992.

Holton, J. R., P. H. Haynes, M. E. Mclntyre, A. R. Douglass, R. B. Rood, and L. Pfister, Stratosphere-troposphere exchange, Rev. Geophys., 33, 403 439, 1995.

Lindzen, R. S. Turbulence and stress owing to gravity wave and tidal breakdown, J. Geophys. Res., Vol. 86, 9707-9714, 1981.

Mclntyre, M. E., and T. N. Palmer, Breaking planetary waves in the stratosphere, Nature, 305, (5935), 593-600, 1983.

Murgatroyd, R., The structure and dynamics of the stratosphere, in G. A. Corby (Ed.), The Global Circulation of the Atmosphere, Royal Meteorological Society, London, 1969.

Murgatroyd, R., and F. Singleton, Possible meridional circulations in the stratosphere and mesosphere, Q. J. R. Meteorol. Soc., 87, 125-135, 1961.

Pedlosky, J., Geophysical Fluid Dynamics, Springer-Verlag, New York, 1979.

Ramanathan, V, and W. L. Grose, A numerical simulation of seasonal stratospheric climate: Part I. Zonal temperatures and winds, J. Atmos. Sei., 35, 600-614, 1978.

U. S. Standard Atmosphere. U.S. Government Printing Office, Washington, DC, 1976.

Wallace, J. M., and P. V Hobbs, Atmospheric Science: An Introductory Survey, Academic, San Diego, 1977.

WMO Scientific Assessment of Ozone Depletion, WMO Global Ozone Research and Monitoring Project, Report No. 37, World Meteorological Organization, Geneva, Switzerland, 1994.

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