EPS200: Atmospheric ChemistryEPS200: Atmospheric Chemistry

Instructors: Daniel J. Jacob and Steven C. WofsyTeaching 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 CHEMISTRYBIG PROBLEMS IN ATMOSPHERIC CHEMISTRY

LOCAL < 100 km

REGIONAL100-1000 km

GLOBAL > 1000 km

Urban smog

Point source

Disasters Visibility

Regional smog

Acid rain

Ozonelayer

Climate

Biogeochemical cycles

GLOBAL OBSERVING SYSTEM FOR TROPOSPHERIC COMPOSITIONGLOBAL OBSERVING SYSTEM FOR TROPOSPHERIC COMPOSITION

Satellites

CTMs solve coupled continuity equations for chemicals on global 3-D Eulerian grid:

Surface networks

Aircraft,ships

Chemical transportmodels (CTMs)

( ) ( )ii i i

CC P L

t

U C C

EmissionsTransportChemistry

Aerosol processesDeposition

x ~100 kmz ~ 1 km

ATMOSPHERIC STRUCTURE AND TRANSPORTATMOSPHERIC STRUCTURE AND TRANSPORT

““SEA LEVEL” PRESSURE MAP (9/2/10, 23Z)SEA LEVEL” PRESSURE MAP (9/2/10, 23Z)

SEA-LEVEL PRESSURE CAN’T VARY OVER MORE SEA-LEVEL PRESSURE CAN’T VARY OVER MORE THAN A NARROW RANGE: 1013 ± 50 hPaTHAN 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 ( ( ) ( ))d P x P x dx dydz F

Acceleration 1 dP

dx

For P = 10 hPa over x = 100 km, ~ 10-2 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 MASS mma a OF THE ATMOSPHEREOF THE ATMOSPHERE

Radius of Earth:6380 km

Mean pressure at Earth's surface:984 hPa

Total number of moles of air in atmosphere:

20 1.8 10 molesaa

a

mN

M

Mol. wt. of air: 29 g mole-1 = 0.029 kg mole-1

2184

5.13 10 kgSurfacea

R Pm

g

VERTICAL PROFILES OF PRESSURE AND TEMPERATUREVERTICAL PROFILES OF PRESSURE AND TEMPERATUREMean values for 30Mean values for 30ooN, MarchN, March

Tropopause

Stratopause

Barometric law (variation of pressure with altitude)Barometric law (variation of pressure with altitude)

• Consider elementary slab of atmosphere:

P(z)

P(z+dz)( ) ( ) a a

dPP z P z dz gdz g

dz

hydrostaticequation

Ideal gas law:a a

a

PM dP M gdz

RT P RT

Assume T = constant, integrate:

/( ) (0) with 7.4 km ( 250 K) z H

a

RTP z P e H T

M g scale height

Barometric law

( ) ( )( ) ; ( 5km)

2

P z P zP z H P z

e /( ) (0) z H

a an z n e

unit area

Application of barometric law: the sea-breeze effectApplication of barometric law: the sea-breeze effect

ATMOSPHERIC TRANSPORTATMOSPHERIC TRANSPORT

Forces in the atmosphere:

• Gravity • Pressure-gradient• Coriolis • Friction

g 1/ P pγ

2 sinc v to R of direction of motion (NH) or L (SH)kfγ v

Equilibrium of forces:

In vertical: barometric law

In horizontal: geostrophic flow parallel to isobars P

P + P

p

c

v

In horizontal, near surface: flow tilted to region of low pressure

P

P + Pc

vf

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 breezeTHE HADLEY CIRCULATION (1735): global sea breeze

HOT

COLD

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’S GLOBAL INFRARED CLOUD MAP (intellicast.com)

TodayBright colors indicate high cloud tops (low temperatures)

shows Intertropical Convergence Zone (ITCZ) as longitudinal band near Equator

TROPICAL HADLEY CELLTROPICAL HADLEY CELL

• Easterly “trade winds” in the tropics at low altitudes• Subtropical anticyclones at about 30o latitude

CLIMATOLOGICAL SURFACE WINDS AND PRESSURESCLIMATOLOGICAL SURFACE WINDS AND PRESSURES(January)(January)

CLIMATOLOGICAL SURFACE WINDS AND PRESSURESCLIMATOLOGICAL SURFACE WINDS AND PRESSURES(July)(July)

500 hPa (~6 km) CLIMATOLOGICAL WINDS IN JANUARY:500 hPa (~6 km) CLIMATOLOGICAL WINDS IN JANUARY:strong mid-latitude westerliesstrong mid-latitude westerlies

500 hPa (~5 km) CLIMATOLOGICAL WINDS IN JULY500 hPa (~5 km) CLIMATOLOGICAL WINDS IN JULYmid-latitude westerlies are weaker in summer than wintermid-latitude westerlies are weaker in summer than winter

ZONAL GEOSTROPHIC FLOW AND THERMAL WIND RELATIONZONAL GEOSTROPHIC FLOW AND THERMAL WIND RELATION

Geostrophic balance:

P

P + Pu

x

y

1 1p

P

y y a

2 sinc u fu *

geopotential height

= latitude

= Earth radius

= angular vel. of Earth

= 2 sin (Coriolis parameter)

ln( / ) log-P coordinate

scale height

o

o

gz

a

f

z H p p

RTH

Mg

1fu

a

*

u R Tfz aH

Thermal wind relation:

ZONAL WIND: VARIATION WITH ALTITUDEZONAL WIND: VARIATION WITH ALTITUDEfollows thermal wind relation

TIME SCALES FOR HORIZONTAL TRANSPORTTIME SCALES FOR HORIZONTAL TRANSPORT(TROPOSPHERE)(TROPOSPHERE)

2 weeks1-2 months

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

GlenCanyon, AZ

Clear day April 16, 2001: Asian dust!

Mean April 2001PM concentrationsmeasured by MODIS

GLOBAL TRANSPORT OF CARBON MONOXIDE (CO) 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 ofCO concentrations at 500 hPa (~6 km)

OBSERVATION OF CO FROM AIRS SATELLITE INSTRUMENTOBSERVATION OF CO FROM AIRS SATELLITE INSTRUMENT

AIRS CO data at 500 hPa (W.W. McMillan)

Averaging kernelsfor AIRS retrieval

ATMOSPHERIC LAPSE RATE AND STABILITYATMOSPHERIC 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

-1/ 9.8 K kmp

gdT dz

C

ATM(observed)

What happens following release depends on the local lapse rate –dTATM/dz:• -dTATM/dz > upward buoyancy amplifies initial perturbation: atmosphere is unstable• -dTATM/dz = zero buoyancy does not alter perturbation: atmosphere is neutral• -dTATM/dz < downward buoyancy relaxes initial perturbation: atmosphere is stable• dTATM/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 WHAT DETERMINES THE LAPSE RATE OF THE ATMOSPHERE?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 equilibriumstate: - dT/dz =

z

T

z

T

Solar heating ofsurface/radiative cooling of air: unstable atmosphere

ATM

ATM

z

Tinitial

final

buoyant motions relaxunstable 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 IN CLOUDY AIR PARCEL, HEAT RELEASE FROM HH22O CONDENSATION MODIFIES O CONDENSATION MODIFIES

RH > 100%:Cloud forms

“Latent” heat releaseas H2O condenses

9.8 K km-1

W2-7 K km-1

RH

100%

T

z

W

Wet adiabatic lapse rate W = 2-7 K km-1

-20 -10 0 10 20 30 Temperature, oC

0

1

4

2

3

Alt

itu

de,

km

cloud

boundarylayer

SUBSIDENCE INVERSIONSUBSIDENCE INVERSION

typically 2 km altitude

DIURNAL CYCLE OF SURFACE HEATING/COOLING:DIURNAL CYCLE OF SURFACE HEATING/COOLING:ventilation of urban pollutionventilation of urban pollution

z

T0

1 km

MIDDAY

NIGHT

MORNING

Mixingdepth

Subsidenceinversion

NIGHT MORNING AFTERNOON

PBLdepth

VERTICAL PROFILE OF TEMPERATUREVERTICAL PROFILE OF TEMPERATUREMean values for 30Mean values for 30ooN, MarchN, March

Alt

itu

de,

km

Surface heating

Latent heat releaseRadiativecooling (ch.7) - 6.5 K km-1

2 K km-1

- 3 K km-1Radiativecooling (ch.7)

Radiative heating:O3 + hO2 + OO + O2 + M O3+M

heat

LATITUDINAL STRUCTURE OF TROPOPAUSE REGIONLATITUDINAL STRUCTURE OF TROPOPAUSE REGION

RADIATIVE-CONVECTIVE EQUILIBRIUM ATMOSPHERERADIATIVE-CONVECTIVE EQUILIBRIUM ATMOSPHERE

BAROCLINIC INSTABILITYBAROCLINIC INSTABILITY

z

latitude0

>>

Buoyant vertical motionIs possible even when

/ 0z

Dominant mechanism for vertical motion in extratropics

FIRST-ORDER PARAMETERIZATION OF TURBULENT FLUXFIRST-ORDER PARAMETERIZATION OF TURBULENT FLUX

• Observed mean turbulent dispersion of pollutants is near-Gaussian parameterize it by analogy with molecular diffusion:

Source

Instantaneousplume

Time-averagedenvelope

<C>

z

Turbulent flux = aznK

C

z

Near-Gaussianprofile

• 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)

Turbulent diffusioncoefficient

TYPICAL TIME SCALES FOR VERTICAL MIXINGTYPICAL TIME SCALES FOR VERTICAL MIXING

• Estimate time t to travel z by turbulent diffusion:

2

5 2 -1 with 10 cm s2 z

z

zt K

K

0 km

2 km

1 day“planetaryboundary layer”

tropopause

5 km

(10 km)

1 week

1 month

10 years