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CHAPTER 4: 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. g p. P. v. P + D P.

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chapter 4 atmospheric transport
CHAPTER 4: 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

gp

P

v

P + DP

gc

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

gp

P

v

gf

P + DP

gc

slide2

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

COLD

THE HADLEY CIRCULATION (1735): global sea breeze

Explains:

  • Intertropical Convergence Zone (ITCZ)
  • Wet tropics, dry poles
  • Easterly trade winds in the tropics

But… Meridional transport of air between Equator and poles results in strong winds in the longitudinal direction because of conservation of angular momentum; this results eventually in unstable conditions.

tropical hadley cell
TROPICAL HADLEY CELL
  • Easterly “trade winds” in the tropics at low altitudes
  • Subtropical anticyclones at about 30o latitude
today s global cloud map
TODAY’S GLOBAL CLOUD MAP

Bright colors indicate the tallest clouds

Intellicast.com

time scales for horizontal transport troposphere
TIME SCALES FOR HORIZONTAL TRANSPORT(TROPOSPHERE)

1-2 months

2 weeks

1-2 months

1 year

vertical transport buoyancy
VERTICAL TRANSPORT: BUOYANCY

Consider an object (density ρ) immersed in a fluid (density ρ’):

z+Dz

Object (r)

Fluid (r’)

P(z) > P(z+Δz) epressure-gradient force

on object directed upward

z

g

Buoyancy acceleration (upward) :

so ρ↑ as T↓

For air,

Barometric law assumes T = T’ e gb = 0(zero buoyancy)

T

T’ produces buoyant acceleration upward or downward

atmospheric lapse rate and stability

“Lapse rate” = -dT/dz

ATMOSPHERIC LAPSE RATE AND STABILITY

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 rateG :

z

G = 9.8 K km-1

stable

z

unstable

What happens following release depends on the local lapse rate –dTATM/dz:

  • -dTATM/dz > Ge upward buoyancy amplifies initial perturbation: atmosphere is unstable
  • -dTATM/dz = Ge zero buoyancy does not alter perturbation: atmosphere is neutral
  • -dTATM/dz < Ge downward buoyancy relaxes initial perturbation: atmosphere is stable
  • dTATM/dz > 0 (“inversion”): very stable

ATM

(observed)

inversion

unstable

T

The stability of the atmosphere against vertical mixing is solely determined by its lapse rate.

temperature sounding at chatham ma today feb 14 2013 at 12z 7 am
TEMPERATURE SOUNDING AT CHATHAM, MAToday (Feb 14, 2013) at 12Z (7 am)

Dew point

Temperature

Aadiabatic

lapse rate

weather.unisys.com

what determines the lapse rate of the atmosphere
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

z

final

G

ATM

G

ATM

initial

G

T

T

T

Initial equilibrium

state: - dT/dz = G

Solar heating of

surface/radiative cooling of air: unstable atmosphere

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 = Gis sure indicator of an unstable atmosphere.

in cloudy air parcel heat release from h 2 o condensation modifies g
IN CLOUDY AIR PARCEL, HEAT RELEASE FROM H2O CONDENSATION MODIFIES G

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

z

T

RH

100%

GW

“Latent” heat release

as H2O condenses

GW = 2-7 K km-1

RH > 100%:

Cloud forms

G

G = 9.8 K km-1

slide15

4

3

Altitude, km

2

1

0

-20 -10 0 10 20 30

Temperature, oC

cloud

boundary

layer

subsidence inversion
SUBSIDENCE INVERSION

typically

2 km altitude

diurnal cycle of surface heating cooling ventilation of urban pollution
DIURNAL CYCLE OF SURFACE HEATING/COOLING:ventilation of urban pollution

z

Subsidence

inversion

Planetary

Boundary

Layer (PBL)

depth

MIDDAY

1 km

G

Mixing

depth

NIGHT

0

MORNING

T

NIGHT

MORNING

AFTERNOON

vertical profile of temperature mean values for 30 o n march
VERTICAL PROFILE OF TEMPERATUREMean values for 30oN, March

Radiative

cooling (ch.7)

- 3 K km-1

Altitude, km

+2 K km-1

Radiative heating:

O3 + hn →O2 + O

O + O2 + M →O3+M

heat

Radiative

cooling (ch.7)

Latent heat release

- 6.5 K km-1

Surface heating

typical time scales for vertical mixing
TYPICAL TIME SCALES FOR VERTICAL MIXING

tropopause

(10 km)

10 years

1 month

2 km

planetary

boundary layer

1 day

0 km