The Atmosphere:Part 4: Moist convection

• Composition / Structure• Radiative transfer

• Vertical and latitudinal heat transport• Atmospheric circulation• Climate modeling

Suggested further reading:

Hartmann, Global Physical Climatology (Academic Press, 1994)

Radiative-convective equilibrium(unsaturated)

Better, but:

• surface still too cold

• tropopause still too warm

Moist convection

Above a thin boundary layer, most atmospheric convection involves phase change of water: condensation releases latent heat

When saturation occurs …..

• Heterogeneous Nucleation

• Supersaturations very small in atmosphere – condensation very fast

• Drop size distribution sensitive to sizedistribution of cloud condensation nuclei

Formation of precipitation(how to produce droplets big enough to fall?)

• Bergeron-Findeisen Process(rapid transfer of moisture from liquid to solid condensate)

• Stochastic coalescence (sensitive to drop size distributions)

• Strongly nonlinear function of cloud water concentration

• Time scale of precipitation formation ~10-30 minutes

— little support for overriding importance of ice nucleation in general

Formation of precipitation(how to produce droplets big enough to fall?)

• Bergeron-Findeisen Process(rapid transfer of moisture from liquid to solid condensate)

• Stochastic coalescence (sensitive to drop size distributions)

• Strongly nonlinear function of cloud water concentration

• Time scale of precipitation formation ~10-30 minutes

Moist variables and thermodynamics

e — vapor pressure of water [hPa]

es(T) — saturation vapor pressure of water [hPa]

q — specific humidity = (mass vapor)/(mass air) [g/kg]

qs — saturation specific humidity [g/kg]

U=q/qs — relative humidity [%]

Clausius-Clapeyron:

(assuming es<<p)

d lnesdT L

RT2

→ es exp − LRT

q ep ,

mvmair

0.622

Destabilization by condensation in saturated air

∂T∂z −Γ

s cp ln

ds cpdT Γdz dQT − L dq

T

dq −dqsp, T − p desT − pdesdT dT

dT Γdz − LcpT dq L

cpTpdesdT

dT

dTdz

−Γm

Γm Γ 1 − LcpTp

desTdT

−1

~If the parcel is saturated, q=qs,

Гm ranges from 3 K/km (moist surface tropical air) to 10 K/km (cold air, e.g. near tropopause); typical value 7 K/km.

Destabilization by condensation in saturated air

∂T∂z −Γ

s cp ln

ds cpdT Γdz dQT − L dq

T

dq −dqsp, T − p desT − pdesdT dT

dT Γdz − LcpT dq L

cpTpdesdT

dT

dTdz

−Γm

Γm Γ 1 − LcpTp

desTdT

−1

Moist adiabatic process:

dQm cpdT g dz L dq 0

Qm cpT gz Lq

moist static energy is conserved

expect uniform Q in convectively adjusted state

Moist radiative-convective equilibrium(Manabe & Strickler 1964)

close to typical observed midlatitudeprofile

Moist radiative-convective equilibriumRoles of various absorbers

Where does convection occur?

tropical deep convection: cold cloud tops

Net outgoing longwave radiation (DJF) (measured from space: Wm-2)

convective clouds not common in desert

belts: radiation from warm low levels

less deep extratropical convective and non-convective clouds

Where does convection occur?

Climatological sea surface temperature

Deep convection over equatorial continents and warmest water

Calculated rad-con equilibrium Tvs. observed T

near-equatorial lapse rate maintained near neutral stability by moist convection

Calculated rad-con equilibrium Tvs. observed T

pole-to-equator temperature contrast too big in equilibrium state (especially in winter)

Zonally averaged net radiation

Diurnally-averaged radiation

IR

solar

Local radiative equilibrium at all latitudes

Zonally averaged net radiation

Diurnally-averaged radiation

Observed radiative budget

Implied energy transport: requires fluid motions to effect the implied heat transport

Roles of atmosphere and ocean

net

ocean

atmosphere

Trenberth & Caron (2001)

Radiative effects of clouds

Low clouds cool:

• increase albedo

• radiate at near-surface T

High clouds warm:

• mostly thin — little effect on albedo

• radiate at low T — weakens IR cooling

Aerosols

Sea salt and dust — most mass but few in number, so less important

Sulfate — small but large in number. Biogenic (via DMS) and human-induced (via SO2)

Volcanic aerosols in the stratosphere

Aerosols: direct effect

Aerosols: indirect effect