STRATOSPHERIC OZONE DISTRIBUTION Marion Marchand CNRS-UPMC-IPSL

O 2 + UV-c -> O + O O + O 2 + M -> O 3 + M d[O 3 ]/dt = 2 J O2 [O 2 ] O 2 (molecules.cm -3 x ) Altitude (km) Photolysis coefficient (s -1 ) O 3 (molecules.cm -3 x ) J O2 (s -1 ) Altitude (km) SHAPE OF OZONE PROFILE UV

VMR (Volume Mixing Ratio)=[O 3 ]/[M] (-> better indicator of chemistry ) O3O3 [M] total air concentration (molec.cm -3 x ) [O 3 ] [M] [O 3 ] concentration (molec.cm -3 x )

Brewer (1949) quotes from Dobson et al. (1929): 'The only way in which we can reconcile the observed high ozone concentration in the Arctic in spring and the low concentration in the tropics, with the hypothesis that the ozone is formed by the action of the sunlight, would be to suppose a general slow poleward drift in the highest atmosphere with a slow descent of air near the poles. Such a current would carry the ozone formed in low latitudes to the poles and concentrate it there. If this were the case the ozone at the poles would be distributed through a moderate depth of atmosphere while that in low latitudes would all be high up.’ [SPARC] O 3 production from O 2 photolysis (molec.cm -3 )O 3 column (Dobson units) tropical minimum tropical maximum latitude month Altitude (km) Why?

(=O 3 concentration)

Brewer (1949) said `The observed distributions of water vapour can be explained by the existence of a circulation in which air enters the stratosphere at the equator, where it is dried by condensation, travels in the stratosphere to temperate and polar regions, and sinks into the troposphere.' [SPARC] Troposphere is humid but stratosphere is very dry. Why? very dry statosphere humid troposphere Temperature (°K) tropopause stratopause

GLOBAL DISTRIBUTION OF CHEMICAL TRACERS

hours days months dO 3 /dt~f(dynamics) dO 3 /dt~f(chemistry) months dO 3 /dt = f(chemistry + dynamics)

RESIDENCE TIME OF AIR IN STRATOSPHERE

CHEMISTRY OF OZONE: CONCEPTS

OXYGEN-ONLY CHEMISTRY ° First chemistry scheme proposed by Chapman in 1930, also called the ‘Chapman cycle’. ° The reactions are O 2 + h → O + O j O2 < 242 nm O + O 2 + M → O 3 + M k O+O2 O 3 + h → O + O 2 j O3 < 336 nm O + O 3 → O 2 + O 2 k O+O3

O 2 + h → O + O O + O 2 + M → O 3 + M O 3 + h → O + O 2 O + O 3 → O 2 + O 2 d[O 3 ]/dt = k O+O2.[O 2 ].[O].[M] -J O3.[O 3 ] -k O+O3.[O 3 ].[O] d[O]/dt = 2.J O2.[O 2 ] +J O3.[O 3 ] -k O+O2.[O 2 ].[O].[M] -k O+O3.[O 3 ].[O]

° k O+O2 et J O3 interconvert O 3 and O very rapidly, so introduce new species O x [=O + O 3 ] known as “odd- oxygen” d[O x ]/dt = d[O]/dt + d[O 3 ]/dt = 2.J O2.O k O+O3.O 3.O Production - Destruction If O x steady-state (i.e. d[O x ]/dt = 0), J O2.[O 2 ]= k O+O3.[O 3 ].[O] ODD OXYGEN CHEMICAL FAMILY fast slow

Mass balance for atomic oxygen O: d[O]/dt = 2.J O2.[O 2 ] +J O3.[O 3 ] -k O+O2.[O 2 ].[O].[M] -k O+O3.[O 3 ].[O] ° if O x steady-state (i.e. d[O x ]/dt = 0), J O2.[O 2 ]= k O+O3.[O 3 ].[O] -> d[O]/dt = J O2.[O 2 ] +J O3.[O 3 ] -k O+O2.[O 2 ].[O].[M] ° interconversion terms >> net chemical terms (chemical family approach): J O3.[O 3 ] >> J O2.[O 2 ] -> d[O]/dt ~ J O3.[O 3 ] -k O+O2.[O 2 ].[O].[M] ° lifetime of O = [O]/loss = 1/(k O+O2.[O 2 ].[M]) < 1 sec -> O steady-state (i.e. d[O]/dt = 0), [O]/[O 3 ] = J O3 / (k O+O2.[O 2 ].[M]) ° as expected, interconversion terms determine partitioning within chemical family

Volume mixing ratio Altitude (km) ° Chapman’s model of [O]/[O 3 ] validated against observations ° [O]/[O 3 ] O x [=O + O 3 ] ~ O 3 and d[O x ]/d= d[O 3 ]/dt O3O3 O

° if O x steady-state (i.e. d[O x ]/dt = 0), J O2.[O 2 ]= k O+O3.[O 3 ].[O] ° if O steady-state (i.e. d[O]/dt = 0), [O]/[O 3 ] = J O3 / (k O+O2.[O 2 ].[M]) ° using [O]=f([O 3 ]) expression in J O2.[O 2 ]= k O+O3.[O 3 ].[O], [O 3 ] = [O 2 ]. (k O+O2.[M] / k O+O3 ) 1/2. (J O2 / J O3 ) 1/2 with [O 2 ] =0.21 [M]

Calculated [O 3 ] from Chapman’s model is much too high compared to observations. Why? d[O 3 ]/dt = J O2.[O 2 ] - k O+O3.[O 3 ].[O] calculated observed

OZONE DESTROYING CATALYTIC CYCLES ° Bates and Nicolet introduced in 1950 the idea of ozone being destroyed via the following catalytic cycle: OH + O 3 → HO 2 + O 2 HO 2 + O → OH + O 2 net: O + O 3 → O 2 + O 2 ° NO 2 cycle in 1970 by Crutzen and also Johnston NO + O 3 → NO 2 + O 2 NO 2 + O → NO + O 2 net: O + O 3 → O 2 + O 2

° ClO cycle in 1974 by Stolarski and Cicerone. Cl + O 3 → ClO + O 2 ClO + O → Cl + O 2 net: O + O 3 → O 2 + O 2 ° A general form: X + O 3 → XO + O 2 fast XO + O → X + O 2 slow net: O + O 3 → O 2 + O 2 with X, the catalyst, being radical H, OH, NO, Cl or Br ° slow reaction is the limiting step in the cycle, d[O 3 ]/dt ~ - 2.k XO+O.[XO].[O]

d[O 3 ]/dt ~ +2.J O2.[O 2 ] -2.k O+O3.[O 3 ]. [O] -2.k HO2+O3. [HO 2 ]. [O 3 ] -2.k NO2+O. [NO 2 ]. [O] -2.k HO2+O3. [ClO]. [O] Catalyst: H, OH, NO, Cl, Br,..

STRATOSPHERIC SOURCE GASES The key stratospheric source gases are long-lived in the troposphere, and hence, once emitted at the surface, they can reach the stratosphere ° Stratospheric hydrogen radicals (OH, HO 2 ) originate mostly from H 2 O injected from the troposphere and from the in-situ oxidation of (natural and anthropogenic) CH 4 by, O( 1 D) + H 2 O → OH + OH O( 1 D) + CH 4 → OH + CH 3 --> more oxidation, more OH ° Most of the stratospheric nitrogen oxide radicals (NO 2, NO) originates from N 2 O oxidised in the stratosphere via the following reaction, O( 1 D) + N 2 O → NO + NO

STRATOSPHERIC SOURCE GASES ° Stratospheric chlorine radicals (Cl, ClO) originate mostly from CFCs that are photolysed by UV radiation following, CF x Cl y + h → Cl + CF x Cl (y-1) --> more oxidation, more Cl

° Sources of stratospheric chlorine radicals (Cl, ClO) natural anthropogenic

RESERVOIR SPECIES ° Up to now, we considered the different catalytic cycles independently. But radicals from one chemical family can interact with radicals from another family. ° Reactions between radicals lead to formation of species with longer lifetimes, much less reactive, called reservoirs, e.g. ClO + HO 2 → HOCl + O 2 HO 2 + NO 2 + M → HO 2 NO 2 + M ClO + NO 2 + M → ClONO 2 + M OH + NO 2 + M → HNO 3 + M NO 3 + NO 2 + M → N 2 O 5 + M ° Reservoir species can be dissociated back rather quickly to release ozone destroying radicals

CHEMICAL MODEL O 3 BUDGET : Production - Destruction d[O 3 ]/dt = 2 J O2 [O 2 ] -  2 k[XO][O] with X=O 2, OH, NO, Cl NO x (N 2 O) HO x (CH 4,H 2 O) ClO x (CFCs) production/destruction rate (molec.cm -3.s -1 ) O/O 3 cycle  O3=f(  X, altitude)

ozone abundance (mPa) ° complete destruction between 14 and 22 km ° - d[O 3 ]/dt ~ 2 %/day

PSC clouds formed at T < ~193 K by co-condensation of HNO 3 and H 2 O

Heterogeneous chemistry: ° chlorine activation (reservoirs species -> into chlorine radicals) ° O 3 destruction -> speed up reactions that are are very slow or non-existent in the gas-phase Ice or HNO 3 /H 2 O PSCs

CATALYTIC CYCLES OF POLAR OZONE DESTRUCTION ° BUT observed loss rates (- d[O 3 ]/dt) ~ 2 %/day -> [O] is much too low (not enough sunlight during early spring) for ClO+O cycle to account for observed loss rates. ° Cl 2 O 2 cycle in 1987 by Molina and Molina. ClO + ClO + M  Cl 2 O 2 + M equilibrium Cl 2 O 2 + h → Cl + ClOO ClOO + M → Cl + O 2 + M 2 x ( Cl + O 3 → ClO + O 2 ) -> little sunlight is required (fast J Cl2O2 ) and more efficient at low temperatures because it slows down thermal decomposition of Cl 2 O 2 (Cl 2 O 2 → ClO + ClO)

° ClO-BrO cycle ClO + BrO → Cl + Br + O 2 slow Cl + O 3 → ClO + O 2 fast Br + O 3 → BrO + O 2 fast net: O + O 3 → O 2 + O 2 -> little sunlight is required because the cycle does not involve atomic oxygen O ° Polar ozone loss rate: d[O 3 ]/dt ~ -2.J Cl2O2. [Cl 2 O 2 ] -2.k ClO+BrO.ClO.BrO

high chlorine loading and very cold/ isolated polar vortex

INTERACTIONS OZONE-CLIMAT

Effet de Serre = Bilan au niveau de la Terre + effet de Serre: Bilan au niveau de la couche (loi Kirchhoff): a=2 Bilan au niveau de la Terre sans effet de Serre (loi de stefan): = Ts = 255 K seulement ! + = _2_2

Stratospheric cooling / heating rate (K/day) Heating: O 3 + UV → O + O 2 O + O 2 + M → O 3 + M (  Q) -> dT/dz > 0 Cooling: mainly CO 2 +/- IR H 2 O and O 3 significant Temperature tropopause stratopause

O 3 and T strongly coupled: d(O 3 )/dt > 0 -> d(T)/dt > 0 -> attenue le d(O 3 )/dt > 0 d(T)/dt > 0 -> d(O 3 )/dt attenue le d(T)/dt > 0 ~ 45 km Relation-2 O3-T: O + O 3 → O 2 + O 2 k = 2.e-11 *exp( / T) Quand T diminue => ralentissement destruction O3 => d(T)/dt > 0 -> d(O 3 )/dt < 0 Relation-1 O3 - T : O 3 + UV → O + O 2 O + O 2 + M → O 3 + M (  Q) => d(O 3 )/dt > 0 -> d(T)/dt > 0

total Ozone changes Greenhouse gas changes

Diminution O3-strato modifie l’équilibre radiatif à la tropopause  effet sur la T au sol : -Augmentation de la pénétration des UV dans le système surface-troposphère (Forçage positif) -Sans ajustement de T: réduction des émissions IR de corp noir vers la surface (Forçage négatif) -Avec ajustement de T: diminution de l’absorption du rayonnement UV par l’O3 => refroidissement de la stratosphère => réduction de l’émission des corps noirs vers le sol (Forçage négatif) Wm-2

COMPLEXITY OF INTERACTIONS CTM + GCM = CCM (Chemistry-Climate Model) -> predict the future evolution of ozone layer DYNAMICS (T, winds) RADIATION CHEMISTRY CO 2 O 3, CH 4 CCM

Future O 3 = f(CFCs, greenhouse gases) Montreal process: chlorine loading -> ozone layer Kyoto process greenhouse gases -> climate