Personal communication from Richard Rood, University of Michigan.
The change of sign in the Coriolis parameter at the Equator leads to a specific class of planetary-scale waves. Kelvin waves propagate eastward in the equatorial zone, which acts as a waveguide, and are trapped in the vicinity of the Equator. Their vertical wavelength is typically 10 km and their wavenumber 1 to 3. Mixed Rossby–gravity waves are also trapped waves, but with vertical wavelengths of 4–8 km in the vertical and wavenumber 3–5. They propagate westward.
Gravity waves are generated by local disturbances in the flow over mountain ranges or in relation to weather (frontal systems, convective storms). As these waves propagate upwards into progressively more rarified air, their amplitude increases until nonlinear effects cause the waves to break and to transfer momentum to the mean flow, predominantly in the mesosphere. This drives the mesosphere away from radiative equilibrium and generates a meridional flow directed from the summer to the winter hemisphere. The associated upwelling in the summer high latitudes leads to strong adiabatic cooling and explains the presence of a very cold summer mesopause (Figure 2.11).
Atmospheric tides are global-scale waves produced by the release of latent heat in the troposphere and the absorption of solar radiation by ozone and water vapor. They propagate upwards and break primarily in the lower thermosphere. Migrating tides are Sun-synchronous and so propagate westward with the apparent motion of the Sun. Since solar forcing is nearly a square wave that is rich in harmonics, waves with periods shorter than 24 hours (e.g., semi-diurnal wave) also are observed. Non-migrating tides produced by the release of latent heat in the troposphere do not follow the motion of the Sun. They may be stationary, or propagate westward or eastward.
Oscillations in the tropical zonal winds, including the quasi-biennial oscillation (QBO with a period of 22 to 34 months) in the lower stratosphere and the semi-annual oscillation (SAO) near the stratopause, are the result of interactions between dissipating waves and the mean flow. The QBO is the major cause of interannual variance of the zonal wind in the equatorial stratosphere. The amplitude of the easterly phase is about twice as strong as that of the westerly phase. The momentum source that produces the oscillation in the zonal wind is provided by the dissipation of Kelvin and mixed Rossby–gravity waves. The Arctic Oscillation (AO), an oscillation in temperature and pressure between the Arctic and mid-latitudes discussed in Section 2.9, extends to the stratosphere, where it affects the strength of the polar vortex with associated effects on stratospheric ozone.
Observations of long-lived tracers in the stratosphere have highlighted the existence of dynamical barriers (Figure 2.23) hindering the exchange of air between different atmospheric regions. The mid-latitude stratosphere is fairly isolated from tropical influences through a barrier against meridional transport situated at 20°–30° latitude. The resulting upward motion confined to the tropical stratosphere is called the tropical pipe. The polar vortex is another dynamical barrier that separates the polar from the mid-latitude stratosphere. The photochemically produced Antarctic ozone hole is sustained as long as the polar vortex remains present, but disappears when the vortex breaks down and strong mixing of air masses takes place.
Figure 2.23 Schematic representation of the most important dynamical barriers in the stratosphere. Thin contour lines represent uniform potential temperature (isentropes). The thick solid line is the mean altitude of the tropopause.
Reproduced from the World Meteorological Organization (WMO, 1999).
References
Aguado E. and Burt J. E. (2013) Understanding Weather and Climate, 6th edition, Pearson Education, Harlow.
Ahrens C. (2000) Essentials of Meteorology, Cengage Learning EMEA, Stamford, CT.
Andrews D. G., Holton J. R., and Leovy C. B. (1987) Middle Atmosphere Dynamics, Academic Press, New York.
Banks P. M. and Kockarts G. (1973) Aeronomy, Academic Press, New York.
Cunningham W. P. and Cunningham M. A. (2010) Principles of Environmental Sciences, Inquiry and Applications, McGraw Hill, New York.
Gill A. E. (1982) Atmosphere–Ocean Dynamics, Academic Press, New York.
Green J. (2004) Atmospheric Dynamics, Cambridge University Press, Cambridge.
Holton J. R. and Hakim G. J. (2013) An Introduction to Dynamic Meteorology, 5th edition, Elsevier Academic Press, Amsterdam.
Jacob D. J. (1999) Introduction to Atmospheric Chemistry, Princeton University Press, Princeton, NJ.
Lutgens F. K. and Tarbuck E. J. (2000), The Atmosphere, 8th edition, Prentice Hall, Englewood Cliffs, NJ.
Lutgens F. K., Tarbuck E. J., and Tasa D. (2013) The Atmosphere: An Introduction to Meteorology, 12th edition, Pearson Education, Harlow.
Mak M. (2011) Atmospheric Dynamics, Cambridge University Press, Cambridge.
Martin J. E. (2006) Mid-Latitude Atmospheric Dynamics: A First Course, Wiley, Chichester.
Pedlosky J. (1987) Geophysical Fluid Dynamics, Springer Verlag, Berlin.
Pidwirny M. (2006) Fundamentals of Physical Geography, 2nd edition, Physicalgeography.net.
Prölss G. W. (2004) Physics of the Earth’s Space Environment: An Introduction, Springer Verlag, Berlin.
Rogers R. R. and Yau M. K. (1989) A Short Course in Cloud Physics, Pergamon Press, Oxford.
Shepherd T. G. (2003) Large-scale atmospheric dynamics for atmospheric chemists, Chem Rev., 103 (12), 4509–4532, doi: 10.1021/cr020511z.
Trenberth K. E., Fasullo J. T., and Kiehl J. (2009) Earth’s global energy budget, Bull. Amer. Meteor. Soc., 90, 311–323, doi: http://dx.doi.org/10.1175/2008BAMS2634.1.
Vallis G. K. (2006) Atmospheric and Oceanic Fluid Dynamics: Fundamental and Large Scale Circulation, Cambridge University Press, Cambridge.
World Meteorological Organization (WMO) (1999) Scientific Assessment of Ozone Depletion: 1998, Global Ozone Research and Monitoring Project – Report No. 44, World Meteorological Organization, Geneva.
Xie P. and Arkin P. A. (1997) Global precipitation: A 17-year monthly analysis based on gauge observations, satellite estimates and numerical model outputs, Bull. Amer. Meteor. Soc., 78, 2539–2558.
Zdunkowski W. and Bott A. (2003) Dynamics of the Atmosphere: A Course in Theoretical Meteorology, Cambridge University Press, Cambridge.
3
Chemical Processes in the Atmosphere
3.1 Introduction
Atmospheric chemistry models simulate the concentrations of chemical species as determined by emissions, transport, chemical production and loss, and deposition. Chemical production and loss are computed for an ensemble of reactions described by kinetic equations and often involve coupling between species. The ensemble of reactions is the chemical mechanism of the model. The purpose of this chapter is to give a primer of important atmospheric chemical processes as a basis for understanding the construction of atmospheric chemistry models. Model equations will be introduced in Chapter 4 and a more in-depth formulation of the kinetic equations will be presented in Chapter 5. Numerical methods for solving complex chemical mechanisms (chemical solvers) are presented in Chapter 6. Emission processes, including global emission budgets for major species, are described in Chapter 9. A sample chemical mechanism is given in Appendix D.
The major components of the atmosphere are molecular nitrogen (N2), molecular oxygen (O2), argon (Ar), and water vapor (H2O). Argon has no chemical reactivity, but N2, O2, and H2O react in the atmosphere to drive chemical processes. Many other species present in trace amounts also contribute to drive chemical processes as described in this chapter. A species directly emitted to the atmosphere is called primary, while a species chemically produced within the atmosphere is called secondary.
Fast chemistry generally involves radical-assisted reaction chains. Radicals are chemical species with unfilled electron orbitals. An orbital can contain two electrons, and having filled orbitals lowers the internal energy of an atom or molecule. An atom or molecule with an odd number of electrons has high reactivity due to its unfilled or
bital. Reaction of a radical (odd number of electrons) with a non-radical (even number of electrons) necessarily produces a radical (since the sum of electrons in the product species is odd), thus propagating a chain reaction. Radicals originate in the atmosphere from cleavage of non-radicals, usually by solar radiation (photolysis). Energetic input from solar radiation is thus critical to driving the chemistry of the atmosphere. The ensemble of chemical reactions enabled by solar radiation is called photochemistry, and chemical mechanisms in models are often called photochemical mechanisms.
Our intent here is to give a compact summary of major processes relevant to atmospheric chemistry. Detailed presentations of chemical and aerosol processes in the atmosphere can be found, for example, in the books by Warneck (1999), Brasseur et al. (1999, 2003), Finlayson-Pitts and Pitts (2000), and Seinfeld and Pandis (2006).
3.2 Oxygen Species and Stratospheric Ozone
Molecular oxygen is photolyzed in the atmosphere by solar UV radiation:
O2 + hv (λ < 242 nm) → O + O
(3.1)
where the reaction threshold of 242 nm corresponds to the minimum energy required to dissociate the molecule. The oxygen atoms combine with O2 by a three-body reaction to produce ozone (O3):
O + O2 + M → O3 + M
(3.2)
The third body M is an inert molecule such as N2 or O2 that collides with the excited O3* product of the collision of O and O2 and takes up its internal energy, thus allowing stabilization of O3* to ground-state O3. The internal energy of the excited M* eventually dissipates as heat. A third body is needed for any reaction where two reactants combine to form a single product. It is standard practice to include M in the expression of a three-body reaction as it may play a limiting role in the kinetics. See Chapter 5 for discussion of the kinetics of three-body reactions.
Reaction (3.1) is the main source of ozone in the stratosphere. Solar radiation of wavelength shorter than 242 nm is efficiently absorbed by both O2 and O3 as it propagates down through the atmosphere, so the rate of (3.1) becomes negligible below 20 km altitude. Production of ozone in the troposphere takes place by a different mechanism, described in Section 3.6.
Ozone is loosely bound and is photolyzed rapidly in daytime:
O3 + hv (λ < 1180 nm) → O + O2
(3.3)
The principal bands for absorption of solar radiation by ozone are the Hartley (200–290 nm), Huggins (310–400 nm) and Chappuis (400–850 nm) bands. Radiation of wavelengths longer than 320 nm produces the O atom in its electronically ground state (3P). Radiation of shorter wavelengths produces the O atom in its electronically excited 1D state (Figure 3.1):
O3 + hv (λ < 320 nm) → O (1D) + O2
(3.4)
Radiation at wavelengths shorter than 234 nm can produce an even more excited state of the O atom (1S). Standard chemical notation omits mention of the spectroscopic state when the species is in its ground state (here O(3P) ≡ O) and retains it when the species is in an excited state (O(1D), O(1S)). Deactivation of excited species occurs by collision with other molecules, a very rapid process in the lower atmosphere. The O(1D) atom is thus deactivated to the ground state:
O (1D) + M → O + M
(3.5)
A small fraction of O(1D) can also react with non-radical species to produce radicals and initiate radical-assisted reaction chains. This will be discussed in Sections 3.3 and 3.4.
Figure 3.1 Enthalpy of formation [kcal mol–1] of gas-phase oxygen species.
The ground-state oxygen atom produced by (3.3) recombines with O2 by (3.2) to regenerate ozone. It can also react with ozone to form two O2 molecules:
O + O3 → O2 + O2
(3.6)
Reactions (3.1), (3.2), (3.3), and (3.6) comprise the Chapman mechanism for stratospheric ozone, originally proposed by Sydney Chapman in 1930. Reactions (3.2) and (3.3) interconvert O and ozone. The lifetime of O against loss by (3.2) is less than a second in the stratosphere and troposphere, so that O and ozone are in photochemical equilibrium during daytime with [O]/[O3] ≪ 1. It follows that reaction (3.3) is not a true sink for ozone because O atoms will immediately return ozone by (3.2). Nor is (3.2) a true source of ozone if the reactant O atoms originated from ozone photolysis by (3.3). Ozone concentration is thus actually controlled by production in (3.1) and loss through (3.6). Accounting is aided by defining an “odd oxygen” family (Ox ≡ O + O3) produced by (3.1), lost by (3.6), and unaffected by (3.2) and (3.3). Since[O]/[O3] ≪ 1, the budget of ozone is actually that of Ox. The general concept of chemical families is important for atmospheric chemistry modeling and is described further in Box 3.1.
Box 3.1 Chemical Families
The concept of “chemical family” is central to atmospheric chemistry. It enables convenient accounting of the budgets of species cycling rapidly with each other. It is nothing more than an accounting device; it does not imply any similarity in the chemical properties of different members of the family. Consider an ensemble of species {A1, …An} cycling with each other by chemical reactions. If this cycling is sufficiently fast, then a chemical equilibrium is established defining concentration ratios [Ai]/[Aj]. Consider now a chemical family Ax representing the ensemble of these species: Ax ≡ A1 + … + An such that . Writing[Ai] = [Ax]([Ai]/[Ax]), we see that the budget of Ai can be defined from the budget of the family Ax and the chemical partitioning [Ai]/[Ax] within the family. In the case where Ai is the dominant member of the family such that[Ai]/[Ax] ≈ 1, the budget of Ai is solely defined by that of Ax. The chemical family is a useful accounting tool if the lifetime of Ax is longer than that of any individual family member, so that Ax is a more conserved quantity in the atmosphere. It is most useful when the family members are in equilibrium so that chemical partitioning within the family can be easily derived.
In the case of the Chapman mechanism described in Section 3.2, there is rapid cycling between O and ozone so that it is useful to group them into a chemical family. That chemical family is commonly called odd oxygen: Ox ≡ O3 + O. The terminology “odd” simply refers to ozone and O having an odd number of O atoms. Since [O]/[O3] < < 1, the ozone budget is well approximated by the Ox budget. The budget of the O atom is defined by that of Ox together with the [O]/[Ox] ≈ [O]/[O3] ratio from chemical equilibrium.
Following on the above, we can qualitatively explain the distribution of ozone in the stratosphere in terms of the odd oxygen budget. The source of odd oxygen from O2 photolysis (3.1) peaks at about 40 km altitude, reflecting opposite trends in O2 number density, which decreases with altitude, and the UV photon flux, which increases with altitude. The maximum ozone number density occurs at a somewhat lower altitude (Figure 3.2) because the sink of odd oxygen from reaction (3.6) increases with altitude as the O atom concentration increases (the O loss rate from (3.2) has a quadratic pressure dependence). The lifetime of odd oxygen in the upper stratosphere is less than a day, sufficiently short that the ozone concentration is determined by the local chemical steady state between production and loss of odd oxygen. Below 30 km, the lifetime of odd oxygen is sufficiently long that the distribution of ozone is affected by transport on a global scale. Coupling of chemistry and transport results in a minimum ozone column in the tropical stratosphere (Figure 3.3), as the Brewer–Dobson circulation carries low-ozone air from the troposphere upward (see Figure 2.23). Box 3.2 gives historical milestones in the development of our knowledge of stratospheric ozone.
Figure 3.2 Typical vertical profile of atmospheric ozone measured by ozonesonde. About 90% of total atmospheric ozone is located in the stratosphere. The origin of the “smog” ozone near the surface is discussed in Section 3.6.
Source: www.esrl.noaa.gov.
Figure 3.3 Total ozone columns measured by the Global Ozone Monitoring Experiment (GOME) satellite instrument in June 2002. The column concentration is expressed in Dobson units (DU), where 1 DU is defined as 0.01 mm of pure ozone at standard conditions of temperature and pressure: 1 DU = 2.69 × 1016 molecules cm–2.
Bo
x 3.2 Historical Milestones in our Understanding of Stratospheric Ozone
The First Steps
1839: Christian Friedrich Schönbein (Basel, Switzerland) identifies a particular odor following electric discharges in air. He calls this “property” ozone from the Greek word οζειν (to smell), and recognizes that it represents a gas.
1845: Jean-Charles de Marignac and Auguste de la Rive (Geneva, Switzerland) suggest that this gas is produced by a transformation of oxygen.
1863: Jean-Louis Soret (Geneva, Switzerland) determines experimentally that ozone is made of three oxygen atoms.
The First Observations
1853: Schönbein detects ozone in the atmosphere.
1858: André Houzeau measures atmospheric ozone in Rouen, France.
1877–1907: Albert Levy conducts systematic observations of ozone at Parc Montsouris in the outskirts of Paris.
Laboratory Investigations and Spectroscopic Observations
Starting in the 1870s, the spectroscopic properties of ozone are investigated and are used to observe ozone in the atmosphere. Important contributions are due to Alfred Cornu (1878), J. Chappuis (1880), Walter N. Hartley (1880/1881), William Huggins (1890), A. Fowler and R. J. Strutt (1917), Charles Fabry and Henri Buisson (1913–1920), Gordon M. B. Dobson (1920s), and F. W. P. Götz (1924).
1920: Fabry and Buisson make the first quantitative observation of the ozone column abundance.
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