Lecture 16 Atmospheric chemistry by zrn20302

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									              Lecture 16: Atmospheric chemistry
• Questions
   – How do solar forcing, radiative and convective transfer set
     the vertical temperature structure of the atmosphere, the
     latitudinal heat transport by the atmosphere, and the global
     wind patterns that drive ocean circulation?
   – How does the greenhouse effect work?
   – What’s up with the ozone layer?
• Tools
   – Gas phase chemistry, radiative and convective heat transfer,
     box models, photochemistry, etc.
• Reading:
   – Not well-treated in either Albarède or Press et al., but some
     issues are raised in Press et al. chapter 23
   – A good short book is Daniel Jacob, Introduction to
     Atmospheric Chemistry


                                                                     1
                      Atmospheric structure: 0-D
• Radiative forcing: the atmosphere is heated from above by UV absorption in
  stratosphere and from below by IR absorption in troposphere. Most sunlight
  (visible peak) gets through to the ground. A significant fraction (~75%) of the
  IR is absorbed and re-radiated at lower temperature.




                     Incoming                      Outgoing
                     radiation                     radiation




                                                                                    2
                    Atmospheric structure: 0-D
• Radiative balance: incoming radiation = outgoing radiation
   – Incoming radiation = FSpr2(1- a)
       • Solar flux at 1 AU, FS = 1380 W/m2
       • Area receiving sunlight is area of Earth projected as a
         disk, pr2, where r = 6471 km.
       • Albedo of earth a ~ 0.3 (where aice~1)
   – Outgoing radiation = 4pr2sTE4
       • Area radiating is surface area of sphere, 4pr2
       • TE is the effective blackbody temperature, s is the Stefan-
         Boltzmann constant
   – So TE = [FS(1- a) / 4s]1/4 = 255 K = –18 °C
• So if the Surface temperature of the Earth were the effective
  radiating temperature (i.e., no atmosphere), all water would be
  frozen.
   – To raise TE to 273 K by lowering albedo alone would require
     a ≤ 0.08!

                                                                       3
           Atmospheric structure: Greenhouse effect I
• Now imagine an atmospheric
  layer that is transparent to
  incoming solar radiation but
  absorbs a fraction f of outgoing
  infrared radiation.
• Now we write two independent
  radiative balance equations, for
  the surface at temperature To
  and for the absorbing layer at T1
  – Per unit area, FS (1- a)/4 = (1- f)sTo4 + fsT14 from space
  – Per unit area, 2fsT14 = fsTo4 at absorbing layer (top and bottom both
     radiate, hence the 2; we used Kirchhoff’s law el = al and a greybody assumption)
   – So To = [FS(1- a) / 4s(1-f/2)]1/4 and T1= To 2–1/4
• Hence actual mean ground temperature To = 288 K for the earth
  implies f = 0.77 (in which case T1 = 242 K). The maximum effect
  from a single layer would be at f = 1 and To = 304 K = 31 °C (in
  which case, naturally T1 = 255 K). Of course there could be
  different absorbing layers at different wavelengths, etc.
                                                                                        4
                Atmospheric structure: Greenhouse effect II
• Here is an actual outgoing radiation spectrum measured over Africa at noon. The
  ground is radiating at 320 K in the non-absorbing atmospheric window. The
  tropopause (where CO2 becomes optically thin) is radiating at ~215 K, the lower
  troposphere is radiating at ~270 K (H2O is thin above ~5 km). The stratosphere is
  radiating at 280 K (where O3 becomes optically thin)


                                                   Of course, this is neither a
                                                   steady-state nor a 0-dimensional
                                                   situation, but in some sense the
                                                   ground-atmosphere system must
                                                   adjust itself to match the integral
                                                   under this curve to incoming
                                                   solar radiation




                                                                                         5
Atmospheric Structure: 1-D
• Pressure structure
   – Hydrostatic equilibrium
     between pressure gradient
     and gravity
                 dP
                     -g
                 dz
   – Ideal gas law (Ma = 28.96
     g/mol)           PM
                         a
                       RT
       dP    gMa
          -     dz
        P     RT
   – Or, assuming constant T:
                                  – Let H = RT/gMa be the scale height of
 P(z)  P(0)exp- gMa z          the atmosphere (7.4 km at T=250 K):
                RT                                       - z/ H
                                            P(z)  P(0)e
   – The logP-z curve in the      – Thus, e.g. a supersonic jet flying at 20
     figure is not quite linear     km is a full scale height above a normal
     because the temperature is     jet flying at 12 km and sees ~1/e times
     not actually constant          the air density in its path                6
Atmospheric Structure: 1-D
• Temperature structure
   – There are three reversals in
     the average temperature
     profile of the atmosphere
     that divide it into four
     layers:
   – The Thermosphere, above
     ~80 km (not shown in
     figure), gets very hot due to
     UV absorption by O2, but
     the density is so low it
     hardly matters                                                          9.8 K/km
   – The Mesosphere is heated
     from below and has
     decreasing T with altitude
   – The stratosphere is heated
     from above by UV                        – Convective stability is defined by the
     absorption by ozone. It is                 temperature gradient relative to the
     stably stratified.                         adiabatic lapse rate. For dry air:
   – The troposphere is heated by - T   T  P   TV a g  g  9.8 K / km
     IR absorption by CO2 and        z S P S z S     Cp V Cp
     H2O and may become
     convectively unstable. a = 1/T for ideal gas • For saturated air, the moist lapse
                                                             rate is more like 6 K/km 7
                     Atmospheric structure: 2-D                        8

• The Earth is unevenly heated by sunlight: the equator receives
  much more radiation per unit area than the poles
• It is the job of the atmosphere and oceans to try to eliminate the
  resulting temperature gradient by zonal heat transport
• The resulting transport is of two types: ocean transport is
  dominantly sensible heat transport (advection of warm water
  polewards), atmospheric transport is dominantly latent heat
  transport (low-latitude evaporation, high-latitude condensation)
                   Atmospheric structure: 2-D




• Total zonal heat transport is obtained from radiative balance
  calculations based on solar forcing and measured outgoing IR
  as a function of latitude (see Problem Set 6)
• Atmospheric heat transport is obtained from Radiosonde data
  that give abundant regular measurements of temperature, winds,
  and humidity
• Oceanic heat transport is obtained by difference, but shows
  important features such as Western Boundary currents in North
                                                                   9
                   Atmospheric structure: 3-D
• In the absence of Coriolis force, solar forcing would drive
  single Hadley cells in each hemisphere, which we can
  understand using the “sea-breeze circulation”




                                                                10
                         Atmospheric structure: 3-D
• But by ±30° latitude, the Coriolis force gets strong enough to break up the
  Hadley circulation, resulting in subtropical oceanic gyres, tropical rainfall,
  the 30° desert band, trade winds, etc.




                                                                         Remember the
                                                                         geostrophic
                                                                         equation?




                                                                                   11
                         Atmospheric structure: 3-D
• But by ±30° latitude, the Coriolis force gets strong enough to break up the
  Hadley circulation, resulting in subtropical oceanic gyres, tropical rainfall,
  the 30° desert band, trade winds, etc.




                                                                         Remember the
                                                                         geostrophic
                                                                         equation?




                                                                                   12
                 Bulk chemistry of atmosphere
• To first order, the modern atmosphere originated by degassing
  of volatile compounds from the earth’s interior. This process
  continues, as demonstrated for example by the 3He flux at mid-
  ocean ridges




                                                                   13
                   Bulk chemistry of atmosphere
• What comes out of the Earth: CO2, H2O, S, N2, noble gases
• What is now in the atmosphere:
  – 78.08% N2
  – 20.05% O2
  – 0.9% Ar
  – 275 380 ppm CO2
  – 0.0005% He
  – 0.00005% H2
• Why are they different?
  – H2O condenses. CO2 dissolves in oceans (60x more than
    atmosphere) and precipitates as carbonates.
  – Noble gas in atmosphere is dominantly radiogenic (40Ar, 4He)
  – H2 is lost from exosphere (He/H2 ratio ~ 10 is 1000x
    primordial ratio)
  – O2 is produced and maintained by biology

                                                                   14
                  Geochemical cycles: Nitrogen
• Here are the basic elements from which we might construct a
  box model to understand the cycling of Nitrogen in the surface
  reservoirs of the Earth:




                                                                   15
                     Geochemical cycles: Nitrogen
• Here is a steady-state quantification of the N box model:


         t = 13 Ma
                                            t = 0.03 a




                                 t = 27 a                     t = 0.6 a

           t=4a


                                                              t = 50 a


                                                         t = 200 Ma!




                                                                          16
              Geochemical cycles: Oxygen and Carbon
• To make atmospheric oxygen, it is not enough to have
  photosynthesis, because respiration and decay of organic carbon
  take the oxygen back to CO2.
• Rather, each mole of oxygen in the atmosphere must be
  compensated by a mole of buried organic C in sediments
• But the total inventory of sedimentary organic C, about 107 Pg, is
  enough to account for 30 times the atmospheric inventory of O2!
   – Think about this next time you burn fossil fuel, but don’t
     think too hard…the industrial increase in CO2 from 280 to
     380 ppm represents a decrease of O2 from 20 to 19.98%
• The balance is accounted for by burial and storage of SO42- and
  Fe2O3, since the mantle provides mostly S2- and FeO.




                                                                       17
                         Geochemical cycles: Carbon
• The important greenhouse gases are CO2, CH4, and H2O (but H2O is a passive
  amplifier, not a cause), so global climate is intimately tied to the carbon cycle




                      About this time Wally Broecker sent an alarmist
                      letter to President Nixon about global cooling




                                                                                      18
                     Geochemical cycles: Carbon
• Proxy records allow longer reconstructions than instrumental data...




                                                                     19
        Carbon
• We have accurate
  measurements of the
  increase in atmospheric
  CO2 concentrations.
• We can estimate the
  effect on climate forcing.




                               • CO2 is the biggest
                                 climate forcing, but
                                 many others are
                                 significant. This is the
                                 2001 assessment by the
                                 IPCC

                                                       20
                              Carbon
• We also know from economic records the total amount of fossil
  fuel burned, and only about half the resulting CO2 has
  accumulated in the atmosphere…where is the rest?
                                        • Taken up by ocean and by
                                          terrestrial biosphere, but
                                          how much of each?
                                        • One good way to tell is
                                          from simultaneous high-
                                          precision data on O2 and
                                          CO2
                                        • Fossil fuel, a mix of coal,
                                          gas, and oil, consumes
                                          1.38 mol O2 for every 1
                                          mol CO2 released
                                        • Land uptake by
                                          photosynthesis is 1:1,
                                          CO2+H2O=CH2O+O2
                                        • Ocean uptake by solubility
                                          and pH adjustment has no
                                          effect on O2
                                        • Ocean warming lowers O2
                                          solubility, though            21
              Stratospheric ozone: production and loss
• The existence of ozone in the stratosphere determines the
  temperature structure of the upper atmosphere and, by the way,
  is essential for life at the Earth’s surface.
• It is therefore worthwhile to understand the chemical kinetics of
  production and loss and the effects of anthropogenic gases.

                                                  O3 photolysis

                                          O2 photolysis
            Stratospheric ozone: production and loss
• Production of ozone in the stratosphere is well understood; the
  mechanism was defined by Chapman in 1930:
                   k1                  (activation reaction, l ≤ 240 nm,
(1)   O2  h  O  O
                                     both oxygens are O(3P) 2 p2 2 p1 2p1 )
                                                                  x   y z


                         k2                    (M is any 3rd body; reactions
(2)               
      O  O2  M  O3  M                   2 and 3 make a rapid cycle
                                               that defines the odd oxygen
                   k3                          or Ox family, l ≤ 320 nm)
(3)
      O3  h  O2  O
               
                                      (This O is O(1D) until some collision)
                  k4                               2    2    0
                                                2 px 2 py 2 pz
(4)           
      O3  O  2O2                  (quenching reaction)

                                           We will see that k2 and k3 are
                                           much greater than k1 and k4,
                                           so that at steady state there is
                                           a significant abundance of Ox
                                           and it does not matter
                                           whether it is O3 or O.
            Stratospheric ozone: production and loss
• Steady-state solution for ozone abundance:
   – Ox steady state means setting rate of reaction 2 equal to 3:
                                [O]     k3
                                           2
                               [O3 ] k2CO na
                                          2

       where CO2 is the mixing ratio of O2 (0.2) and na is the
         number density of all air molecules (altitude dependent)
    – Then steady-state for entry and exit to Ox cycle means
      setting rate of reaction 1 equal to reaction 4:
                                               1
                        k1CO2 na  k1k2 2
              [O3 ]                      
                                                      3
                                                 CO2 na 2
                          k4 [O]     k3k 4 
Note: the photolysis rate constants k1 and k3 include a term for the ultraviolet
flux, so they increase upwards as the column depth of O2 and O3 above z
decreases. On the other hand, the number density of the atmosphere falls off
exponentially with increasing altitude.
            Stratospheric ozone: production and loss
                                              Up here no
                                              O2 to react




                                                    Down here
                                                    no UV flux




• Steady-state abundance of O3 depends on product of k1 and
  na3/2, so there is a maximum at ~30 km. The general shape of
  the prediction is a good match to abundance data.
• But the Chapman mechanism predicts a factor of 2 too much
  O3…the source is certain so there must be another sink!
               Stratospheric ozone: production and loss
• The missing sinks for ozone come from catalytic loss cycles, i.e.
  reaction cycles where the ozone destruction agent is regenerated and
  can destroy many ozone molecules before it exits the cycle
• Good catalysts are generally radical species with an odd number of
  electrons such as the hydroxyl radical OH (9 e–)
• The OH loss cycle must be initiated by O(1D), normally produced
  by k3 photolysis of O3:

                  k3                1
  O3  h  O2  O( D)
                                          Activation steps: removes
                                            one Ox, makes 2 OH,
                 1
                
   H2 O  O( D) 2OH                     requires deep UV and H2O

           
   OH  O3 HO2  O2                     Catalytic cycle: net
                                            reaction is 2O3 -> 3O2
                                            (OH and OH2 are the HOx
            
   HO2  O3 OH  2O2                    radical family)


           
  OH  HO2 H2O  O2                     Termination step, slow
               Stratospheric ozone: production and loss
• The OH loss cycle is efficient in principle but does not account for enough
  O3 loss in the middle and upper stratosphere
   – Limited at low altitude by low UV flux
   – Limited at high altitude by very low H2O mixing ratio
• A more important (but more complicated) natural catalytic loss cycle (whose
  discovery earned Paul Crutzen a Nobel prize) is the NOx radical system:
  NOx: radical
  species NO and
  NO2
  NOy: nonradical
  reservoir species
  N2O5 and HNO3
              Stratospheric ozone: production and loss
  • When reaction of NO with O3 produces NO2, it has several
    possible fates:
     – Photolysis cycles it back to NO with no net effect
     – Reaction with O catalytically destroys two Ox species
     – Reaction with OH radical or O3 inactivates one NOx
  NOx cycle: no net effect, but           Catalytic cycle branch: net
  rapidly cycles NO & NO2                 reaction consumes 2 Ox

        
NO  O3 NO2  O2                           
                                      NO  O3 NO2  O2
                                               k3         1
NO2  h NO  O
                                    O3  h  O2  O( D)
                                               
                   k2                           
                                      NO2  O NO  O2
            
O  O2  M  O3  M
               
      NO2  OH HNO3 Termination step, daytime
NO2  O3 NO3  O2 NO3  NO2  M N2O5  M
                                               Termination step,
                                                      nighttime
             Stratospheric ozone: production and loss
• Because N2O from the biosphere is stable and non-condensable,
  it reaches upper stratosphere and meets enough O(1D) to form
  NOx and initiate O3-loss catalysis
• The other O3-loss mechanism is mostly anthropogenic and
  involves sources of Cl and Br stable enough to reach
  stratosphere




                                Together, the Chapman source roughly
                                balances these four loss mechanisms and
                                explains the O3 abundance at all heights
                                in the normal stratosphere: Chapman
                                (O3+O), HOx, NOx, and ClOx
       Polar Stratospheric ozone: the Antarctic Ozone Hole
• The total disappearance of the ozone layer in the mid-
  stratosphere over Antarctica provides a challenge to the
  standard gas-phase theory of ozone balance, since in winter
  there is not enough light to drive the HOx, NOx, or ClOx losses
                                       October
                                       2000




                                       October
                                       2002: ?
Polar Stratospheric ozone: the Antarctic Ozone Hole
Polar Stratospheric ozone: the Antarctic Ozone Hole
                             The story is complicated but here is
                             its essence:
                             1) When temperature drops below
                             197 K Polar Stratospheric Clouds
                             (PSC) of HNO3•3H2O can form even
                             though H2O is very rare.
                             2) PSC surfaces provide rapid total
                             conversion of inactive Cl species
                             HCl and ClNO3 to active ClOx and
                             HNO3.
                              3) When temperatures rise again in
                             September, the HNO3 would
                             scavenge all the ClOx back to
                             ClNO3, except that the PSC particles
                             grow big enough to sediment out of
                             the stratosphere, removing HNO3
                             and leaving behind active ClOx.
                             4) When light returns in Southern
                             Spring, at high ClO concentrations a
                             catalytic photolysis mechanism can
                             run that consumes O3 without O(1D).
                             (More Nobel-quality chemistry, this time to
                             Molina and Rowland)
                         Tropospheric Ozone
• Yes, the air in Pasadena really is getting better!




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