EPS200: Atmospheric Chemistry Instructors: Daniel J. Jacob and Steven C. Wofsy Teaching Fellow: Helen M. Amos EPS 200 is intended as a “core” graduate course in atmospheric chemistry • Assumes no prior knowledge of atm chem • Suitable as “breadth” for students in other fields • complements other core course EPS208 (Physics of Climate) • broad survey of field, prepares for + complements more advanced courses: -EPS 236 Environmental Modeling -EPS 238 Spectroscopy and Radiative Transfer of Planetary Atmospheres -ES 267 Aerosol Science and Technology -ES 268 Environmental Chemical Kinetics BIG PROBLEMS IN ATMOSPHERIC CHEMISTRY Disasters Visibility Ozone layer Urban smog Regional smog Climate Point source Acid rain Biogeochemical cycles LOCAL < 100 km REGIONAL 100-1000 km GLOBAL > 1000 km GLOBAL OBSERVING SYSTEM FOR TROPOSPHERIC COMPOSITION Satellites Surface networks Chemical transport models (CTMs) Aircraft, ships CTMs solve coupled continuity equations for chemicals on global 3-D Eulerian grid: Emissions Transport Ci U Ci Pi (C) Li (C) Chemistry t Aerosol processes Deposition Dx ~100 km Dz ~ 1 km ATMOSPHERIC STRUCTURE AND TRANSPORT “SEA LEVEL” PRESSURE MAP (9/2/10, 23Z) SEA-LEVEL PRESSURE CAN’T VARY OVER MORE THAN A NARROW RANGE: 1013 ± 50 hPa Consider a pressure gradient at sea level operating on an elementary air parcel dxdydz: P(x) P(x+dx) Pressure-gradient force Vertical area dydz Acceleration dF ( P( x) P( x dx))dydz 1 dP dx For DP = 10 hPa over Dx = 100 km, ~ 10-2 m s-2 a 100 km/h wind in 3 h! Effect of wind is to transport air to area of lower pressure a dampen DP On mountains, however, the surface pressure is lower, and the pressure-gradient force along the Earth surface is balanced by gravity: P(z+Dz) P-gradient gravity P(z) aThis is why weather maps show “sea level” isobars; a The fictitious “sea-level” pressure at a mountain site assumes an air column to be present between the surface and sea level MASS ma OF THE ATMOSPHERE Mean pressure at Earth's surface: 984 hPa Radius of Earth: 6380 km ma 4 R2 PSurface g 5.13 10 kg 18 Total number of moles of air in atmosphere: ma Na 1.8 1020 moles Ma Mol. wt. of air: 29 g mole-1 = 0.029 kg mole-1 VERTICAL PROFILES OF PRESSURE AND TEMPERATURE Mean values for 30oN, March Stratopause Tropopause Barometric law (variation of pressure with altitude) • Consider elementary slab of atmosphere: P(z+dz) P(z) P( z ) P( z dz ) a gdz unit area PM a a RT Ideal gas law: dP Mag dz P RT dP a g dz hydrostatic equation Assume T = constant, integrate: P( z ) P(0)e z / H RT with scale height H 7.4 km (T 250 K) Mag Barometric law na ( z) na (0)e z / H P( z ) P( z H ) ; e P( z ) P( z 5km) 2 Application of barometric law: the sea-breeze effect ATMOSPHERIC TRANSPORT Forces in the atmosphere: • Gravity g • Pressure-gradient γp 1/ P • Coriolis c 2v sin to R of direction of motion (NH) or L (SH) • Friction γ f kv Equilibrium of forces: In vertical: barometric law p P In horizontal: geostrophic flow parallel to isobars v c P + DP In horizontal, near surface: flow tilted to region of low pressure p f v c P P + DP Air converges near the surface in low pressure centers, due to the modification of geostrophic flow under the influence of friction. Air diverges from high pressure centers. At altitude, the flows are reversed: divergence and convergence are associated with lows and highs respectively THE HADLEY CIRCULATION (1735): global sea breeze COLD HOT COLD Explains: • Intertropical Convergence Zone (ITCZ) • Wet tropics, dry poles •General direction of winds, easterly in the tropics and westerly at higher latitudes Hadley thought that air parcels would tend to keep a constant angular velocity. Meridional transport of air between Equator and poles results in strong winds in the longitudinal direction. …but this does not account for the Coriolis force correctly. TODAY’S GLOBAL INFRARED CLOUD MAP (intellicast.com) shows Intertropical Convergence Zone (ITCZ) as longitudinal band near Equator Today Bright colors indicate high cloud tops (low temperatures) TROPICAL HADLEY CELL • Easterly “trade winds” in the tropics at low altitudes • Subtropical anticyclones at about 30o latitude CLIMATOLOGICAL SURFACE WINDS AND PRESSURES (January) CLIMATOLOGICAL SURFACE WINDS AND PRESSURES (July) 500 hPa (~6 km) CLIMATOLOGICAL WINDS IN JANUARY: strong mid-latitude westerlies 500 hPa (~5 km) CLIMATOLOGICAL WINDS IN JULY mid-latitude westerlies are weaker in summer than winter ZONAL GEOSTROPHIC FLOW AND THERMAL WIND RELATION gz geopotential height 1 P 1 p y y a = latitude a = Earth radius = angular vel. of Earth P u f = 2sin (Coriolis parameter) z* H ln( p / po ) log-P coordinate H P + DP c 2u sin fu RTo scale height Mg fu Geostrophic balance: Thermal wind relation: f 1 a u R T z* aH y x ZONAL WIND: VARIATION WITH ALTITUDE follows thermal wind relation TIME SCALES FOR HORIZONTAL TRANSPORT (TROPOSPHERE) 1-2 months 2 weeks 1-2 months 1 year Illustrates long time scale for interhemispheric exchange Dust transport over the Pacific, April 21-25, 1998 • What is buoyancy? R. Husar TRANSPORT OF ASIAN DUST TO NORTH AMERICA Clear day Glen Canyon, AZ Mean April 2001 PM concentrations measured by MODIS April 16, 2001: Asian dust! GLOBAL TRANSPORT OF CARBON MONOXIDE (CO) Sources of CO: Incomplete combustion (fossil fuel, biofuel, biomass burning), oxidation of VOCs Sink of CO: atmospheric oxidation by OH radical (lifetime ~ 2 months) MOPITT satellite observations of CO concentrations at 500 hPa (~6 km) OBSERVATION OF CO FROM AIRS SATELLITE INSTRUMENT AIRS CO data at 500 hPa (W.W. McMillan) Averaging kernels for AIRS retrieval ATMOSPHERIC LAPSE RATE AND STABILITY “Lapse rate” = -dT/dz Consider an air parcel at z lifted to z+dz and released. It cools upon lifting (expansion). Assuming lifting to be adiabatic, the cooling follows the adiabatic lapse rate G : z stable G = 9.8 K km-1 g G dT / dz 9.8 K km-1 Cp z unstable inversion unstable What happens following release depends on the local lapse rate –dTATM/dz: ATM • -dTATM/dz > G e upward buoyancy amplifies (observed) initial perturbation: atmosphere is unstable • -dTATM/dz = G e zero buoyancy does not alter perturbation: atmosphere is neutral • -dTATM/dz < G e downward buoyancy relaxes T initial perturbation: atmosphere is stable • dTATM/dz > 0 (“inversion”): very stable The stability of the atmosphere against vertical mixing is solely determined by its lapse rate. WHAT DETERMINES THE LAPSE RATE OF THE ATMOSPHERE? • • An atmosphere left to evolve adiabatically from an initial state would eventually tend to neutral conditions (-dT/dz = G at equilibrium Solar heating of surface and radiative cooling from the atmosphere disrupts that equilibrium and produces an unstable atmosphere: z z ATM G z final G ATM T Initial equilibrium state: - dT/dz = G G initial T Solar heating of surface/radiative cooling of air: unstable atmosphere T buoyant motions relax unstable atmosphere back towards –dT/dz = G • Fast vertical mixing in an unstable atmosphere maintains the lapse rate to G. Observation of -dT/dz = G is sure indicator of an unstable atmosphere. IN CLOUDY AIR PARCEL, HEAT RELEASE FROM H2O CONDENSATION MODIFIES G Wet adiabatic lapse rate GW = 2-7 K km-1 z T RH “Latent” heat release as H2O condenses RH > 100%: Cloud forms GW 2-7 K km-1 G 9.8 K km-1 100% GW G 4 Altitude, km 3 cloud 2 boundary layer 1 0 -20 -10 0 10 Temperature, oC 20 30 SUBSIDENCE INVERSION typically 2 km altitude DIURNAL CYCLE OF SURFACE HEATING/COOLING: ventilation of urban pollution z PBL depth Subsidence inversion MIDDAY 1 km G Mixing depth 0 NIGHT MORNING T NIGHT MORNING AFTERNOON VERTICAL PROFILE OF TEMPERATURE Mean values for 30oN, March Altitude, km Radiative cooling (ch.7) - 3 K km-1 2 K km-1 Radiative heating: O3 + hn e O2 + O O + O2 + M e O3+M heat Radiative cooling (ch.7) - 6.5 K km-1 Latent heat release Surface heating LATITUDINAL STRUCTURE OF TROPOPAUSE REGION RADIATIVE-CONVECTIVE EQUILIBRIUM ATMOSPHERE BAROCLINIC INSTABILITY q3 > z q2 > q1 Buoyant vertical motion Is possible even when q / z 0 0 latitude Dominant mechanism for vertical motion in extratropics FIRST-ORDER PARAMETERIZATION OF TURBULENT FLUX • Observed mean turbulent dispersion of pollutants is nearGaussian eparameterize it by analogy with molecular diffusion: Instantaneous plume Time-averaged envelope z Near-Gaussian profile Source Turbulent flux = K z na C z <C> Turbulent diffusion coefficient • Typical values of Kz: 102 cm2s-1 (very stable) to 107 cm2 s-1 (very unstable); mean value for troposphere is ~ 105 cm2 s-1 • Same parameterization (with different Kx, Ky) is also applicable in horizontal direction but is less important (mean winds are stronger) TYPICAL TIME SCALES FOR VERTICAL MIXING • Estimate time Dt to travel Dz by turbulent diffusion: Dz Dt 2 2K z with K z 105 cm2s-1 tropopause (10 km) 10 years 5 km “planetary 2 km boundary layer” 0 km 1 month 1 week 1 day

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# 100 km Dz - Atmospheric Chemistry Modeling Group