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 LOCAL < 100 km REGIONAL km GLOBAL > 1000 km Urban smog Point source Disasters Visibility Regional smog Acid rain Ozone layer Climate Biogeochemical cycles

GLOBAL OBSERVING SYSTEM FOR TROPOSPHERIC COMPOSITION Satellites CTMs solve coupled continuity equations for chemicals on global 3-D Eulerian grid: Surface networks Aircraft, ships Chemical transport models (CTMs) Emissions Transport Chemistry Aerosol processes Deposition  x ~100 km  z ~ 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) Vertical area dydz Pressure-gradient force Acceleration For  P = 10 hPa over  x = 100 km,  ~ m s -2  100 km/h wind in 3 h! Effect of wind is to transport air to area of lower pressure  dampen  P On mountains, however, the surface pressure is lower, and the pressure-gradient force along the Earth surface is balanced by gravity: P(z) P(z+  z) P-gradient gravity  This is why weather maps show “sea level” isobars;  The fictitious “sea-level” pressure at a mountain site assumes an air column to be present between the surface and sea level

MASS m a OF THE ATMOSPHERE Radius of Earth: 6380 km Mean pressure at Earth's surface: 984 hPa Total number of moles of air in atmosphere: Mol. wt. of air: 29 g mole -1 = kg mole -1

VERTICAL PROFILES OF PRESSURE AND TEMPERATURE Mean values for 30 o N, March Tropopause Stratopause

Barometric law (variation of pressure with altitude) Consider elementary slab of atmosphere: P(z) P(z+dz) hydrostatic equation Ideal gas law: Assume T = constant, integrate: Barometric law unit area

Application of barometric law: the sea-breeze effect

ATMOSPHERIC TRANSPORT Forces in the atmosphere: Gravity Pressure-gradient Coriolis Friction to R of direction of motion (NH) or L (SH) Equilibrium of forces: In vertical: barometric law In horizontal: geostrophic flow parallel to isobars P P +  P pp cc v In horizontal, near surface: flow tilted to region of low pressure P P +  P cc v ff pp

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 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) Today Bright colors indicate high cloud tops (low temperatures) shows Intertropical Convergence Zone (ITCZ) as longitudinal band near Equator

TROPICAL HADLEY CELL Easterly “trade winds” in the tropics at low altitudes Subtropical anticyclones at about 30 o 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 Geostrophic balance: P P +  P u x y Thermal wind relation:

ZONAL WIND: VARIATION WITH ALTITUDE follows thermal wind relation

TIME SCALES FOR HORIZONTAL TRANSPORT (TROPOSPHERE) 2 weeks 1-2 months 1 year

Illustrates long time scale for interhemispheric exchange

Dust transport over the Pacific, April 21-25, 1998 R. Husar What is buoyancy?

TRANSPORT OF ASIAN DUST TO NORTH AMERICA TRANSPORT OF ASIAN DUST TO NORTH AMERICA Glen Canyon, AZ Clear dayApril 16, 2001: Asian dust! Mean April 2001 PM concentrations measured by MODIS

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 T z  = 9.8 K km -1 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  : z “Lapse rate” = -dT/dz ATM (observed) What happens following release depends on the local lapse rate –dT ATM /dz: -dT ATM /dz >   upward buoyancy amplifies initial perturbation: atmosphere is unstable -dT ATM /dz =   zero buoyancy does not alter perturbation: atmosphere is neutral -dT ATM /dz <   downward buoyancy relaxes initial perturbation: atmosphere is stable dT ATM /dz > 0 (“inversion”): very stable unstable inversion unstable 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 =  at equilibrium Solar heating of surface and radiative cooling from the atmosphere disrupts that equilibrium and produces an unstable atmosphere: Initial equilibrium state: - dT/dz =  z T z T Solar heating of surface/radiative cooling of air: unstable atmosphere ATM   ATM z T initial final  buoyant motions relax unstable atmosphere back towards –dT/dz =  Fast vertical mixing in an unstable atmosphere maintains the lapse rate to  Observation of -dT/dz =  is sure indicator of an unstable atmosphere.

IN CLOUDY AIR PARCEL, HEAT RELEASE FROM H 2 O CONDENSATION MODIFIES  RH > 100%: Cloud forms “Latent” heat release as H 2 O condenses  9.8 K km -1  W  2-7 K km -1 RH 100% T z  WW Wet adiabatic lapse rate  W = 2-7 K km -1

Temperature, o C Altitude, km cloud boundary layer

SUBSIDENCE INVERSION typically 2 km altitude

DIURNAL CYCLE OF SURFACE HEATING/COOLING: ventilation of urban pollution z T 0 1 km MIDDAY NIGHT MORNING Mixing depth Subsidence inversion NIGHTMORNINGAFTERNOON  PBL depth

VERTICAL PROFILE OF TEMPERATURE Mean values for 30 o N, March Altitude, km Surface heating Latent heat release Radiative cooling (ch.7) K km -1 2 K km K km -1 Radiative cooling (ch.7) Radiative heating: O 3 + h   O 2 + O O + O 2 + M  O 3 +M heat

LATITUDINAL STRUCTURE OF TROPOPAUSE REGION

RADIATIVE-CONVECTIVE EQUILIBRIUM ATMOSPHERE

BAROCLINIC INSTABILITY z latitude0    > > Buoyant vertical motion Is possible even when Dominant mechanism for vertical motion in extratropics

FIRST-ORDER PARAMETERIZATION OF TURBULENT FLUX Observed mean turbulent dispersion of pollutants is near- Gaussian  parameterize it by analogy with molecular diffusion: Source Instantaneous plume Time-averaged envelope z Near-Gaussian profile Typical values of K z : 10 2 cm 2 s -1 (very stable) to 10 7 cm 2 s -1 (very unstable); mean value for troposphere is ~ 10 5 cm 2 s -1 Same parameterization (with different K x, K y ) is also applicable in horizontal direction but is less important (mean winds are stronger) Turbulent diffusion coefficient

TYPICAL TIME SCALES FOR VERTICAL MIXING Estimate time  t to travel  z by turbulent diffusion: 0 km 2 km 1 day “planetary boundary layer” tropopause 5 km (10 km) 1 week 1 month 10 years