Composition of Tropospheric Air Sources, Transport, and Sinks of Trace Gases

1. PHYS-575/CSI-655Introduction to Atmospheric Physics and ChemistryLecture Notes #5: Atmospheric Chemistry • Composition of Tropospheric Air • Sources, Transport, and Sinks of Trace Gases • Some Important Tropospheric Trace Gases • Tropospheric Aerosols • Air Pollution • Tropospheric Chemical Cycles • Stratospheric Chemistry

2. Atmospheric Chemistry Atmospheric chemistry concerns the sources, properties, and effects of the many chemical species that exist in the atmosphere. Effects: • Acid Rain (deposition) • Ozone Depletion (Antarctic Ozone Hole) • Air Pollution (Photochemical Smog) • Trace gases and aerosol formation (Climate effects) • Oxidation State of the Atmosphere • Chemical Destruction of Greenhouse Gases

4. Tropospheric Gases - Categories • Non-Reactive Species: • Ar, Ne, He • (2) Long-Lifetime Molecules: • N2, O2, CO2 • (3) Disequilibrium Gases: • CH4, N2O, SO2, H2S, NH3, CS2, DMS • (4) Radicals: • OH, HO2, NO, CH3 • (5) Photochemical Species: • O3, H2, NOy • (6) Biogenic Gases: • CH4, NH3, H2S, DMS NB – A given gas may belong to more than one category.

5. Atmospheric Abundances Abundances by Volume of Dry Air: Nitrogen (N2) = 78.084% Oxygen (O2) = 30.946% Argon (Ar) = 0.934% Carbon Dioxide (CO2) = 0.03% Total: = 99.99% The volumes occupied by different gases at the same temperature and pressure are proportional to the numbers of molecules of the respective gases, i.e., each atom or molecule occupies the same average volume. At standard temperature (273oK) and pressure (1bar = 1.013 x 105 Pa) the number of molecules in 1 m3 of air is given by the ideal gas law P = nkT, and is known as Loschmidt’s Number.

6. Concentration and Residence Time Loschmidt’s Number No = 2.687 x 1019 molecules cm-3 Example: N2O occupies 310 ppbv (parts per billion by volume) of air. Number Density of N2O = 310 x 10-9 x No = 8.33 x 1012 cm-3 Residence Time: M = mass/volume of constituent in the atmosphere (kg m-3 or gm cm-3) F = rate/volume of its removal from the atmosphere (kg m-3 s-1 or gm cm-3 s-1) Note carefully that the residence time a more general concept than the chemical lifetime. The chemical lifetime is the loss timescale due to only chemical processes. The residence time includes all loss/or removal processes that affect a gas.

7. Spatial and Temporal Variability - Scales Stable Molecules Radicals

8. 2. Sources and Sinks of Trace Gases • Major Categories of Sources and Sinks • Biogenic • Solid Earth • Oceanic • In situ formation/loss Biogenic examples Photosynthesis: CO2(g) + H2O(l) + hν CH2O(s) + O2(g) Methanogenic Bacteria (Cow stomachs & termites): CO, CO2 CH4 Nitrogen Fixing Bacteria: N2  NH3, N2O, NO Phytoplankton: CH4, SO2, COS  DMS, dimethyl disulfide (DMDS) Seaweed: NaCl, Br, I, CH4  CH3Cl, CH3Br, CH3I VOCs: several thousand Volatile Organic Compounds (plants & human activity) Biomass burning: CO2

9. Solid Earth Sources & Sinks Volcanoes: H2O, CO2, SO2, H2S, COS, HCl, etc. Radiogenic elements: He, Ar, Rn Weathering of Rocks: O2, CO2 loss Oceanic Sources & Sinks Soluble gases: CO2, N2, O2 In situ Formation/loss (formation of chemicals by reactions driven by sunlight, lightening, solar ions, etc.) Ozone formation: O + O2 + M  O3 + M Water Photolysis: H2O + hν  H + OH Carbon monoxide destruction: CO + OH  CO2 + H Methane photolysis: CH4 + hν  CH3 + H

10. Chemical Reactions In situ formation is a major source of many important trace gases such as ozone (O3) and a major loss mechanism for other gases such as methane (CH4). Much of atmospheric chemistry is initiated and driven by the absorption of solar UV radiation (λ < 3000 Angstroms = 300nm). If the photon energy (hν) is high enough, it will break apart molecules into more reactive components (e.g. radicals). Example: Formation of the Hydroxyl Radical (OH) Above ~60 km altitude: H2O + hν H + OH Below ~60 km: O3 + hν  O2 + O* (λ < 3200 Angstrom) O* + H2O  2 OH Net result: O3 + H2O + hν O2 + 2OH

11. 3. Important Tropospheric Trace Gases The hydroxyl radical OH is powerful oxidizing agent that reacts with most trace gases containing H, C, N, O, & S. Reactions with OH control the loss of numerous atmospheric pollutants and control the buildup of some GH gases. Because of its role in removing many pollutants, OH has been called the atmosphere’s detergent. The dominant loss of OH is via: OH + CO  CO2 + H OH + CH4  CH3 + H2O OH is so intensively reactive that its steady state abundance is extremely low, of order 10-4 ppbv, or mixing ratio ~10-13. Only in the past decade have techniques been developed to measure such low abundances.

12. Important Hydroxyl (OH) as an Oxidant

13. Carbon Monoxide (CO) CO is produced by the oxidation of CH4 or Non-methane hydro-carbons (NMHC) such as isoprene. Fossil fuel combustion and the burning of biomass are also very significant. The loss of CO CO + OH  CO2 + H is also the dominant loss process of OH in non-urban and non-forested areas. So the distribution of OH is sometimes inferredby the abundances of CO. CO has a seasonal cycle, with high abundances during winter when OH concentrations are low, and low abundances in spring and summer when OH concentrations are high.

14. Carbon Monoxide (CO)

15. Types of Atmospheric Components

16. Atmospheric Ozone (O3) Ozone Vertical Distribution Ozone “layer” • Approximately 90% of Earth’s ozone is in • the stratosphere. The Ozone “layer” absorbs • solar UV radiation that can produce cellular • damage. • However, ozone in direct contact with cells • is a poison because of its oxidizing nature. Ozone – O3 Source: O2 + photon  O + O (photolysis) O + O2  O3 (chemical reaction)

17. Ozone Spatial Variation: Altitude vs. Latitude Ozone “layer”

18. Ozone Seasonal Variation: Latitude vs. Month of Year

19. Seasonal Variation of Tropospheric Ozone

20. Ozone: Human Influences

21. The Basics of Chemical Kinetics 2-Body Chemical Reaction: A + B  C + D (k = rate coefficient, units: cm3 s-1) A, B: Reactants C, D: Products Reaction Rate: Rate = k [A] [B] (molecules cm-3 s-1) where [X] denotes the concentration of reactant X (cm-3) Homogeneous Chemistry: Reaction occurs entirely in gas phase. If reaction occurs in gas phase, but products later condense, it is still considered a homogeneous gas phase reaction. Heterogeneous Chemistry: Reaction occurs on surface or in liquid background. Chemistry occurring inside liquid aerosols are considered heterogeneous. Photolysis: A + hν  B + C (J photolysis rate coefficient, units: s-1) Where hν denotes a photon of frequency ν, and energy E = hν. Photolysis Rate: Rate = J [A] (molecules cm-3 s-1)

22. Basics - continued Sometimes a reaction of the form A + B  C + D is extremely slow because of the necessity of conserving both energy and momentum during the reaction. In those cases a nearby atom or molecule can act as a “third body” to speed up the reaction, but without being consumed during the reaction. The rate of the reaction in that case is dependent upon the density of the third body M which we denote, by convention, as [M]. 3-Body Chemical Reaction: A + B + M  C + M (k = rate coefficient units: cm6 s-1) 3-Body Rate: k [A] [B] [M] (cm-3 s-1) An example of a 3-body reaction is the stratospheric formation of ozone (O3): O + O2 + M  O3 + M where M may be a background N2 molecule, or even another O2 molecule.

23. Photolysis Lifetime The chemical lifetime is the timescale for chemical process to change the abundance of a certain species by a factor of e-1. First example: photolysisA + hν  B + C Photolysis Rate:J [A] (cm-3 s-1) The rate of change of the concentration of A is then, since photolysis is a loss of A: d[A]/dt = -J[A] Which gives d[A]/[A] = - J dt This can be integrated from starting concentration [A]o to [A] as time goes from 0 to t. [A](t) = [A]o e-Jt = [A]o e-t/τ So the chemical loss (1/e) time scale is τ = 1/J, which has units of time.

24. Chemical Reaction Lifetime Second example: 2-body reactionA + B  C + D Reaction Rate: Rate = k [A] [B] (cm-3 s-1) The rate of change of the concentration of A is then: d[A]/dt = -k [A] [B] Which gives d[A]/[A] = - k [B] dt If we assume [B] is constant, then the solution for [A](t) is: [A](t) = [A]o e-k[B]t = [A]o e-t/τ So the chemical loss timescale is τ = 1/k[B] But every time an A is destroyed, a B and a C are produced, so -d[A]/dt = d[B]/dt = d[C]/dt

25. Stratospheric Ozone Production of “active” oxygen: O2 + hν O + O O3 source: O + O2 + M  O3 + M O3 loss: O3 + hν  O2 + O Net result: nothing However, the absorption of solar UV radiation leads to atmospheric heating, even though it is a chemical “do nothing” cycle. The loss of “active” oxygen is accomplished by catalytic cycles of the form H + O3  OH + O2 OH + O  H + O2 Net result: O + O3  2 O2 Ozone “layer” General catalytic cycle: X + O3 XO +O2 XO + O  X + O2 Net result: O + O3  2O2 Where X = H, OH, Cl, N, NO, Br, etc.

26. Tropospheric Ozone Let [Y] denote the number density of gas Y: Catalytic Cycle: NO2 + hν NO + O Rate Coefficient J; Rate = J[NO2] O + O2 + M  O3 +M Rate Coefficient k1 Rate = k1[O][O2]M O3 + NO  NO2 + O2 Rate Coefficient k2 Rate = k2[O3][NO] Net: “do nothing” cycle These reactions are extremely fast. So the rate of the last reaction is the same as that of the first reaction. Thus we can write: J[NO2] = k2[O3][NO] and solve for the abundance of [O3] [O3] = (J/k2) ([NO2]/[NO])Photostationary State Relation

27. Hydrogen Radicals and Ozone Chemistry [O3] = (J/k2) ([NO2]/[NO])Photostationary State Relation The O3 abundance predicted by this cycle is far below observed values. The reason is that the OH radicals reduce the abundance of NO and increase the abundance of NO2 via OH + CO + O2 HO2 + CO2 HO2 + NO  OH + NO2 The loss of OH: OH + CO  CO2 + H (still leaves a radical) OH + HO2  H2O + O2 (removes radicals)

28. Methane and Photochemical Smog CH4 – methane: biomass burning, cows, termites, ride paddies, etc. Oxidation: OH + CH4 H2O + CH3 OH + CH3CHO  H2O + CH3CO etc Net result: CH4 + 2O2 + 2NO  H2O + 2NO2 + HCHO (formaldehyde) CH3CO + …  CH3COO2 + NO2  CH3COO2NO2 (PAN) Formaldehyde is an eye irritant. Peroxyacetyl nitrate (PAN) is an important component of photochemical smog. Most of the complex hydrocarbons condense onto aerosols, there initiating even more complex liquid phase chemistry.

29. Photochemical Smog • Diurnal Variation: • Hydrocarbons, NOx build up in the morning rush hour. • Ozone, PAN, aldehydes form during the day and build up until the afternoon.

30. Sulfur Gases and Acid Rain Sulfur is assimilated by living organisms and is released as various gases as an end product of metabolism. Some important sulfur gases are: H2S – hydrogen sulfide: blue-green algae in marshlands, soils, ocean, volcanoes CH3SCH3 – dimethyl sulfide (DMS): from seaweed and phytoplankton SO2 – sulfur dioxide: oxidation of DMS and H2S, volcanoes, biomass burning COS – carbonyl sulfide: biomass decay, fossil fuel, oxidation of CS2 CS2 – carbon disulfide: biomass decay & burning, fossil fuel OH + H2S  H2O + HS HS + O3  HSO + O2 HS + NO2  HSO + NO HSO + O3  HSO2 + O2 HSO2 + O2  HO2 + SO2 OH + SO2 + M  HOSO2 + M HOSO2 + O2  HO2 + SO3 Sulfuric Acid SO3 + H2O  H2SO4 (absorbed in aerosols, deposition)

31. 4. Tropospheric Aerosols Atmospheric aerosols are suspensions of small solid and/or liquid particles (excluding water cloud particles) in air that have negligible terminal fall speeds.

32. What are Aerosols? • Atmospheric aerosols (or particulate matter) are solid or liquid particles or both suspended in air with diameters between about 0.002 μm to about 100 μm. • Aerosol particles vary greatly in size, source, chemical composition, amount and distribution in space and time, and how long they survive in the atmosphere. • Primary atmospheric aerosols are particulates that emitted directly into the atmosphere (for instance, sea-salt, mineral aerosols (or dust), volcanic dust, smoke and soot, some organics). • Secondary atmospheric aerosols are particulates that formed in the atmosphere by gas-to-particles conversion processes (for instance, sulfates, nitrates, some organics).

33. Why are Aerosols Important? A significant fraction of atmospheric aerosols are anthropogenic. • Importance of aerosols: • heterogeneous chemistry • air quality and human health • visibility reduction • acid deposition • cloud formation • climate and climate change

34. Atmospheric Aerosols in the Upper Atmosphere http://www.misu.su.se/~gumbel/norfa/maa_scheme.gif

35. Atmospheric Particles Size Distribution Diameter and radius of a particle are both used to characterized its size. If a particle is non-spherical, its equivalent radius is used. There are several ways to define particle equivalent radius (for instance, aerodynamic equivalent radius, which is radius of a sphere that experiences the same resistance to motion as the non-spherical particle). http://jan.ucc.nau.edu/~doetqp-p/courses/env440/env440_2/lectures/lec35/Fig35_10.gif

36. Marine Aerosols http://www.atmos.washington.edu/~beckya/ALKCartoon.gif

37. Surface Bubbles

38. Sources of Atmospheric Particulates

39. In Situ Sources of Atmospheric Particulates (in Tg/year)

40. Number Distribution of Tropospheric Aerosols

42. Number Distributions of Tropospheric Aerosols

43. 5. Air Pollution • In Urban and industrialized locations, anthropogenic emissions lead to large concentrations of undesirable chemical species which can adversely affect air quality, visibility, and pose threats to human health. • Severe air pollution episodes occur when the rates of emissions orformation of pollutants greatly exceed the rates at which the pollutants are dispersed by winds, chemical reactions, or deposition. • Severe air pollution episodes tend to occur in association with extended periods of light winds and strong static stability. Emissions from Fossil Fuel Combustion: CO, CO2, NOx (radicals of N), SO2, hydrocarbons, etc.

44. Urban, Continental, and Marine Particulates

45. Anthropogenic Greenhouse Gases

46. Nitrogen-containing Gases

47. Sulfur-containing gases

48. 7. Stratospheric ChemistryOzone Vertical Distribution • It forms a protective shield that reduces the intensity of UV radiation (with wavelengths between 0.23 and 0.32 μm) from the sun that reaches the earth’s surface. • Because of the absorption of UV radiation by O3, it determines the vertical profile of temperature in the stratosphere. • It is involved in many stratospheric chemical reactions.

49. The Ozone Layer The ozone layer, or ozonosphere layer (rarely used term), is that part of the Earth’s atmosphere which contains relatively high concentrations of ozone (O3). "Relatively high" means a few parts per million - much higher than the concentrations in the lower atmosphere but still small compared to the main components of the atmosphere. The ozone layer was discovered in 1913 by the French physicists Charles Fabry and Henri Buisson. Its properties were explored in detail by the British meteorologist G. M. B. Dobson, who developed a simple spectrophotometer that could be used to measure stratospheric ozone from the ground. Between 1928 and 1958 Dobson established a worldwide network of ozone monitoring stations which continues to operate today. The "Dobson unit", a convenient measure of the total amount of ozone in a column overhead, is named in his honor. http://www.search.com/reference/Ozone_layer

50. Solar UV Radiation and Ozone http://en.wikipedia.org/wiki/Image:Ozone_altitude_UV_graph.jpg