radiative transfer in the atmosphere - ssec, uw-madison · 2011. 9. 19. · radiative transfer in...
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Radiative Transfer in the
Atmosphere
Lectures in Brienza
19 Sep 2011
Paul Menzel
UW/CIMSS/AOS
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Outline
Radiation Definitions
Planck Function
Emission, Absorption, Scattering
Radiative Transfer Equation
Satellite Derived Met Parameters
Microwave Considerations
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Spectral Characteristics of Energy Sources and Sensing Systems
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Terminology of radiant energy
Energy from
the Earth Atmosphere
over time is
Flux
which strikes the detector area
Irradiance
at a given wavelength interval
Monochromatic
Irradiance
over a solid angle on the Earth
Radiance observed by
satellite radiometer
is described by
can be inverted to
The Planck function
Brightness temperature 4
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Definitions of Radiation __________________________________________________________________
QUANTITY SYMBOL UNITS
__________________________________________________________________
Energy dQ Joules
Flux dQ/dt Joules/sec = Watts
Irradiance dQ/dt/dA Watts/meter2
Monochromatic dQ/dt/dA/d W/m2/micron
Irradiance
or
dQ/dt/dA/d W/m2/cm-1
Radiance dQ/dt/dA/d/d W/m2/micron/ster
or
dQ/dt/dA/d/d W/m2/cm-1/ster
__________________________________________________________________
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Telescope Radiative Power Capture
proportional to A Spectral Power radiated from A2 to A1 = L(λ) A11 W/m
A1
A2
R
1 = A2 / R2
{ Note: A1 A2 / R2 = A11 = A22 }
Radiance from surface
= L(λ) W /m2 /sr /m
Instrument collection area
Earth pixel
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Using wavelengths
c2/λT
Planck’s Law B(λ,T) = c1 / λ5 / [e -1] (mW/m2/ster/cm)
where λ = wavelengths in cm
T = temperature of emitting surface (deg K)
c1 = 1.191044 x 10-5 (mW/m2/ster/cm-4)
c2 = 1.438769 (cm deg K)
Wien's Law dB(λmax,T) / dλ = 0 where λ(max) = .2897/T
indicates peak of Planck function curve shifts to shorter wavelengths (greater wavenumbers)
with temperature increase. Note B(λmax,T) ~ T5.
Stefan-Boltzmann Law E = B(λ,T) dλ = T4, where = 5.67 x 10-8 W/m2/deg4.
o
states that irradiance of a black body (area under Planck curve) is proportional to T4 . Brightness Temperature
c 1
T = c2 / [λ ln( _____ + 1)] is determined by inverting Planck function
λ5Bλ
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Spectral Distribution of Energy Radiated
from Blackbodies at Various Temperatures
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Bλ/Bλmax
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Area / 3 2x Area /3
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BT4 – BT11
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Observed BT at 4 micron
Range BT[250, 335]
Range R[0.2, 1.7]
Values over land
Larger than over water
Reflected Solar everywhere
Stronger over Sunglint
Window Channel:
•little atmospheric absorption
•surface features clearly visible
Clouds are cold
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Observed BT at 11 micron
Range BT [220, 320]
Range R [2.1, 12.4]
Values over land
Larger than over water
Undetectable Reflected Solar
Even over Sunglint
Window Channel:
•little atmospheric absorption
•surface features clearly visible
Clouds are cold
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2
1
B (λ, T) / B (λ, 273K)
200 250 300
Temperature (K)
4μm
6.7μm
10μm
15μm
microwave
Temperature Sensitivity of B(λ,T) for typical earth temperatures
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∆B/B= ∆T/T
Integrating the Temperature Sensitivity Equation
Between Tref and T (Bref and B):
B=Bref(T/Tref)
Where =c2/Tref (in wavenumber space)
(Approximation of) B as function of and T
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B=Bref(T/Tref)
B=(Bref/ Tref) T
B T
The temperature sensitivity indicates the power to which the Planck radiance
depends on temperature, since B proportional to T satisfies the equation. For
infrared wavelengths,
= c2/T = c2/T.
__________________________________________________________________
Wavenumber Typical Scene Temperature
Temperature Sensitivity
900 300 4.32
2500 300 11.99
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Non-Homogeneous FOV
N
1-N
Tcold=220 K
Thot=300 K
B=N*B(Tcold)+(1-N)*B(Thot)
BT=N*Tcold+(1-N)*Thot
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For NON-UNIFORM FOVs:
Bobs=NBcold+(1-N)Bhot
Bobs=N Bref(Tcold/Tref)+ (1-N) Bref(Thot/Tref)
Bobs= Bref(1/Tref) (N Tcold
+ (1-N)Thot)
For N=.5
Bobs/Bref=.5 (1/Tref) ( Tcold
+ Thot)
Bobs/Bref=.5 (1/TrefTcold) (1+ (Thot/ Tcold)
)
The greater the more predominant the hot term
At 4 µm (=12) the hot term more dominating than at 11 µm (=4)
N
1-N Tcold
Thot
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Cloud edges and broken clouds appear different in 11 and 4 um images.
T(11)**4=(1-N)*Tclr**4+N*Tcld**4~(1-N)*300**4+N*200**4
T(4)**12=(1-N)*Tclr**12+N*Tcld**12~(1-N)*300**12+N*200**12
Cold part of pixel has more influence for B(11) than B(4)
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Table 6.1 Longwave and Shortwave Window Planck Radiances (mW/m**2/ster/cm-1) and Brightness Temperatures (degrees K) as a function of Fractional Cloud Amount (for cloud of 220 K and surface of 300 K) using B(T) = (1-N)*B(Tsfc) + N*B(Tcld).
Cloud
Fraction N
Longwave Window Rad Temp
Shortwave Window Rad Temp
Ts -T1
1.0
23.5
220
.005
220
0
.8
42.0
244
.114
267
23
.6
60.5
261
.223
280
19
.4
79.0
276
.332
289
13
.2
97.5
289
.441
295
6
.0
116.0
300
.550
300
0
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1.0 0.8 0.6 0.4 0.2 0.0 N
T
300
280
260
240
220
BT4
BT11
SW and LW BTs for different cloud amounts
when Tcld=220 and Tsfc=300
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Using wavenumbers
c2/T
Planck’s Law B(,T) = c13 / [e -1] (mW/m2/ster/cm-1)
where = # wavelengths in one centimeter (cm-1)
T = temperature of emitting surface (deg K)
c1 = 1.191044 x 10-5 (mW/m2/ster/cm-4)
c2 = 1.438769 (cm deg K)
Wien's Law dB(max,T) / d = 0 where max) = 1.95T
indicates peak of Planck function curve shifts to shorter wavelengths (greater wavenumbers)
with temperature increase.
Stefan-Boltzmann Law E = B(,T) d = T4, where = 5.67 x 10-8 W/m2/deg4.
o
states that irradiance of a black body (area under Planck curve) is proportional to T4 . Brightness Temperature
c13
T = c2/[ln(______ + 1)] is determined by inverting Planck function
B
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Using wavenumbers Using wavelengths
c2/T c2 /T
B(,T) = c13 / [e -1] B(,T) = c1
/{ 5 [e -1] }
(mW/m2/ster/cm-1) (mW/m2/ster/m)
(max in cm-1) = 1.95T (max in cm)T = 0.2897
B(max,T) ~ T**3. B( max,T) ~ T**5.
E = B(,T) d = T4, E = B(,T) d = T4,
o o
c13 c1
T = c2/[ln(______ + 1)] T = c2/[ ln(______ + 1)]
B 5 B
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Incoming solar radiation (mostly visible) drives the earth-atmosphere (which emits
infrared).
Over the annual cycle, the incoming solar energy that makes it to the earth surface
(about 50 %) is balanced by the outgoing thermal infrared energy emitted through
the atmosphere.
The atmosphere transmits, absorbs (by H2O, O2, O3, dust) reflects (by clouds), and
scatters (by aerosols) incoming visible; the earth surface absorbs and reflects the
transmitted visible. Atmospheric H2O, CO2, and O3 selectively transmit or absorb
the outgoing infrared radiation. The outgoing microwave is primarily affected by
H2O and O2.
Solar (visible) and Earth emitted (infrared) energy
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Visible
(Reflective Bands)
Infrared
(Emissive Bands)
Radiative Transfer Equation
in the IR
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Relevant Material in Applications of Meteorological Satellites
CHAPTER 2 - NATURE OF RADIATION
2.1 Remote Sensing of Radiation 2-1
2.2 Basic Units 2-1
2.3 Definitions of Radiation 2-2
2.5 Related Derivations 2-5
CHAPTER 3 - ABSORPTION, EMISSION, REFLECTION, AND SCATTERING
3.1 Absorption and Emission 3-1
3.2 Conservation of Energy 3-1
3.3 Planetary Albedo 3-2
3.4 Selective Absorption and Emission 3-2
3.7 Summary of Interactions between Radiation and Matter 3-6
3.8 Beer's Law and Schwarzchild's Equation 3-7
3.9 Atmospheric Scattering 3-9
3.10 The Solar Spectrum 3-11
3.11 Composition of the Earth's Atmosphere 3-11
3.12 Atmospheric Absorption and Emission of Solar Radiation 3-11
3.13 Atmospheric Absorption and Emission of Thermal Radiation 3-12
3.14 Atmospheric Absorption Bands in the IR Spectrum 3-13
3.15 Atmospheric Absorption Bands in the Microwave Spectrum 3-14
3.16 Remote Sensing Regions 3-14
CHAPTER 5 - THE RADIATIVE TRANSFER EQUATION (RTE)
5.1 Derivation of RTE 5-1
5.10 Microwave Form of RTE 5-28
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Emission, Absorption, Reflection, and Scattering
Blackbody radiation B represents the upper limit to the amount of radiation that a real
substance may emit at a given temperature for a given wavelength.
Emissivity is defined as the fraction of emitted radiation R to Blackbody radiation,
= R /B .
In a medium at thermal equilibrium, what is absorbed is emitted (what goes in comes out) so
a = .
Thus, materials which are strong absorbers at a given wavelength are also strong emitters at
that wavelength; similarly weak absorbers are weak emitters.
If a, r, and represent the fractional absorption, reflectance, and transmittance,
respectively, then conservation of energy says
a + r + = 1 .
For a blackbody a = 1, it follows that r = 0 and = 0 for blackbody radiation. Also, for a
perfect window = 1, a = 0 and r = 0. For any opaque surface = 0, so radiation is either
absorbed or reflected a + r = 1.
At any wavelength, strong reflectors are weak absorbers (i.e., snow at visible wavelengths),
and weak reflectors are strong absorbers (i.e., asphalt at visible wavelengths). 30
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Planetary Albedo
Planetary albedo is defined as the fraction of the total incident solar
irradiance, S, that is reflected back into space. Radiation balance then
requires that the absorbed solar irradiance is given by
E = (1 - A) S/4.
The factor of one-fourth arises because the cross sectional area of the earth
disc to solar radiation, r2, is one-fourth the earth radiating surface, 4r2.
Thus recalling that S = 1380 Wm-2, if the earth albedo is 30 percent,
then E = 241 Wm-2. For radiative equilibrium E=σT4 or T=255K
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Selective Absorption and Transmission Assume that the earth behaves like a blackbody and that the atmosphere has an absorptivity
aS for incoming solar radiation and aL for outgoing longwave radiation. Let Ya be the
irradiance emitted by the atmosphere (both upward and downward); Ys the irradiance emitted
from the earth's surface; and E the solar irradiance absorbed by the earth-atmosphere system.
Then, radiative equilibrium requires E - (1-aL) Ys - Ya = 0 , at top of atm,
(1-aS) E - Ys + Ya = 0 , at sfc. Solving yields (2-aS)
Ys = E , and
(2-aL)
(2-aL) - (1-aL)(2-aS)
Ya = E .
(2-aL) Since aL > aS, the irradiance and hence the radiative equilibrium temperature at the earth
surface is increased by the presence of the atmosphere. With aL = 0.8 and aS = 0.1 and
E = 241 Wm-2, Stefans Law yields a blackbody temperature at the surface of 286 K, in
contrast to 255 K for an atmospheric absorptance independent of wavelength (aS = aL).
The atmospheric temperature in this example is 245 K.
E (1-al) Ys Ya
top of the atmosphere
(1-as) E Ys Ya
earth surface.
Incoming Outgoing IR
solar
(2-aS)
Ys = E = Ts4
(2-aL)
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Transmittance Transmission through an absorbing medium for a given wavelength is governed by
the number of intervening absorbing molecules (path length u) and their absorbing
power (k) at that wavelength. Beer’s law indicates that transmittance decays
exponentially with increasing path length - k u (z)
(z ) = e where the path length is given by u (z) = dz .
z k u is a measure of the cumulative depletion that the beam of radiation has
experienced as a result of its passage through the layer and is often called the optical
depth . Realizing that the hydrostatic equation implies g dz = - q dp where q is the mixing ratio and is the density of the atmosphere, then p - k u (p)
u (p) = q g-1 dp and (p o ) = e .
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Spectral Characteristics of
Atmospheric Transmission and Sensing Systems
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Relative Effects of Radiative Processes
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0.01 0.1 1.0 10.0
There are 3 modes :
- « nucleation »: radius is
between 0.002 and 0.05 m.
They result from combustion
processes, photo-chemical
reactions, etc.
- « accumulation »: radius is
between 0.05 m and 0.5 m.
Coagulation processes.
- « coarse »: larger than 1 m.
From mechanical processes like
aeolian erosion.
« fine » particles (nucleation and
accumulation) result from anthropogenic
activities, coarse particles come from
natural processes.
Aerosol Size Distribution
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Scattering of early morning sun light from haze
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Measurements in the Solar Reflected Spectrum
across the region covered by AVIRIS
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AVIRIS Movie #1
AVIRIS Image - Linden CA 20-Aug-1992
224 Spectral Bands: 0.4 - 2.5 m
Pixel: 20m x 20m Scene: 10km x 10km
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AVIRIS Movie #2
AVIRIS Image - Porto Nacional, Brazil
20-Aug-1995
224 Spectral Bands: 0.4 - 2.5 m
Pixel: 20m x 20m Scene: 10km x 10km
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Relevant Material in Applications of Meteorological Satellites
CHAPTER 2 - NATURE OF RADIATION
2.1 Remote Sensing of Radiation 2-1
2.2 Basic Units 2-1
2.3 Definitions of Radiation 2-2
2.5 Related Derivations 2-5
CHAPTER 3 - ABSORPTION, EMISSION, REFLECTION, AND SCATTERING
3.1 Absorption and Emission 3-1
3.2 Conservation of Energy 3-1
3.3 Planetary Albedo 3-2
3.4 Selective Absorption and Emission 3-2
3.7 Summary of Interactions between Radiation and Matter 3-6
3.8 Beer's Law and Schwarzchild's Equation 3-7
3.9 Atmospheric Scattering 3-9
3.10 The Solar Spectrum 3-11
3.11 Composition of the Earth's Atmosphere 3-11
3.12 Atmospheric Absorption and Emission of Solar Radiation 3-11
3.13 Atmospheric Absorption and Emission of Thermal Radiation 3-12
3.14 Atmospheric Absorption Bands in the IR Spectrum 3-13
3.15 Atmospheric Absorption Bands in the Microwave Spectrum 3-14
3.16 Remote Sensing Regions 3-14
CHAPTER 5 - THE RADIATIVE TRANSFER EQUATION (RTE)
5.1 Derivation of RTE 5-1
5.10 Microwave Form of RTE 5-28
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Re-emission of Infrared Radiation
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Molecular absorption of IR by
vibrational and rotational excitation
CO2, H2O, and O3
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Radiative Transfer Equation
The radiance leaving the earth-atmosphere system sensed by a
satellite borne radiometer is the sum of radiation emissions
from the earth-surface and each atmospheric level that are
transmitted to the top of the atmosphere. Considering the
earth's surface to be a blackbody emitter (emissivity equal to
unity), the upwelling radiance intensity, I, for a cloudless
atmosphere is given by the expression
I = sfc B( Tsfc) (sfc - top) +
layer B( Tlayer) (layer - top)
layers
where the first term is the surface contribution and the second
term is the atmospheric contribution to the radiance to space. 46
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Rsfc R1 R2 R3
τ4 = 1 _______________________________
τ3 = transmittance of upper layer of atm
τ2 = transmittance of middle layer of atm
τ1= transmittance of lower layer of atm sfc for earth surface recalling that i = 1- τi for each layer, then
Robs = sfc Bsfc τ1 τ2 τ3 + (1-τ1) B1 τ2 τ3 + (1- τ2) B2 τ3 + (1- τ3) B3
Satellite observation comes from the sfc and the layers in the atm
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Radiative Transfer through the Atmosphere
λ
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The emissivity of an infinitesimal layer of the atmosphere at
pressure p is equal to the absorptance (one minus the transmittance
of the layer). Consequently,
(layer) (layer to top) = [1 - (layer)] (layer to top)
Since transmittance is multiplicative
(layer to top) - (layer) (layer to top) = -Δ(layer to top)
So we can write
I = sfc B(T(ps)) (ps) - B(T(p)) (p) .
p
which when written in integral form reads
ps
I = sfc B(T(ps)) (ps) - B(T(p)) [ d(p) / dp ] dp .
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Weighting Functions
1
z1 z2
zN
d/dz
z1 z2
zN
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In standard notation, I =
sfc B(T(ps)) (ps) + (p) B(T(p)) (p)
p
The emissivity of an infinitesimal layer of the atmosphere at pressure p is equal
to the absorptance (one minus the transmittance of the layer). Consequently, (p) (p) = [1 - (p)] (p)
Since transmittance is an exponential function of depth of absorbing constituent, p+p p
(p) (p) = exp [ - k q g-1 dp] * exp [ - k q g-1 dp] = (p + p)
p o
Therefore (p) (p) = (p) - (p + p) = - (p) .
So we can write I =
sfc B(T(ps)) (ps) - B(T(p)) (p) .
p
which when written in integral form reads ps
I = sfc B(T(ps)) (ps) - B(T(p)) [ d(p) / dp ] dp .
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To investigate the RTE further consider the atmospheric contribution to the radiance to space of
an infinitesimal layer of the atmosphere at height z, dIλ(z) = Bλ(T(z)) dλ(z) . Assume a well-mixed isothermal atmosphere where the density drops off exponentially with
height ρ = ρo exp ( - z), and assume kλ is independent of height, so that the optical depth can
be written for normal incidence
σλ = kλ ρ dz = -1 kλ ρo exp( - z)
z and the derivative with respect to height dσλ = - kλ ρo exp( - z) = - σλ .
dz Therefore, we may obtain an expression for the detected radiance per unit thickness of the layer
as a function of optical depth, dIλ(z) dλ(z)
= Bλ(Tconst) = Bλ(Tconst) σλ exp (-σλ) .
dz dz The level which is emitting the most detected radiance is given by d dIλ(z) { } = 0 , or where σλ = 1.
dz dz Most of monochromatic radiance detected is emitted by layers near level of unit optical depth.
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When reflection from the earth surface is also considered, the Radiative Transfer
Equation for infrared radiation can be written o
I = sfc B(Ts) (ps) + B(T(p)) F(p) [d(p)/ dp] dp
ps
where F(p) = { 1 + (1 - ) [(ps) / (p)]2 }
The first term is the spectral radiance emitted by the surface and attenuated by
the atmosphere, often called the boundary term and the second term is the
spectral radiance emitted to space by the atmosphere directly or by reflection
from the earth surface. The atmospheric contribution is the weighted sum of the Planck radiance
contribution from each layer, where the weighting function is [ d(p) / dp ].
This weighting function is an indication of where in the atmosphere the majority
of the radiation for a given spectral band comes from.
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Including surface emissivity atm
ps 'λ(p) ref atm sfc
Isfc = ελ Bλ(Ts) λ(ps) + (1-ελ) λ(ps) Bλ(T(p)) d ln p
λ o ln p
ps 'λ(p)
Iλ = ελ Bλ(Ts) λ(ps) + (1-ελ) λ(ps) Bλ(T(p)) d ln p
o ln p
o λ(p) __________
+ Bλ(T(p)) d ln p sfc
ps ln p
using 'λ(p) = λ(ps) / λ(p) then.
'λ(p) λ(ps) λ(p)
_____ = - _______ _____
ln p (λ(p))2 ln p
Thus
o λ(p)
Iλ = ελ Bλ(T(ps)) λ(ps) + B(T(p)) Fλ(p) d ln p
ps ln p
where
λ(ps)
Fλ(p) = { 1 + (1 - ελ) [ ]2 } .
λ(p) 54
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The transmittance to the surface can be expressed in terms of transmittance to the
top of the atmosphere by remembering
1 ps
'λ(p) = exp [ - kλ(p) g(p) dp ]
g p
ps p
= exp [ - + ]
o o
= λ(ps) / λ(p) .
So
'λ(p) λ(ps) λ(p)
= - .
ln p (λ(p))2 ln p
[ remember that λ(ps, p) λ(p, 0) = λ(ps, 0) and λ(ps, p) = λ(p, ps) ]
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Earth emitted spectra overlaid on Planck function envelopes
CO2
H20
O3
CO2
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Spectral Characteristics of
Atmospheric Transmission and Sensing Systems
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Earth emitted spectra overlaid on Planck function envelopes
CO2
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Longwave CO2
14.7 1 680 CO2, strat temp
14.4 2 696 CO2, strat temp
14.1 3 711 CO2, upper trop temp
13.9 4 733 CO2, mid trop temp
13.4 5 748 CO2, lower trop temp
12.7 6 790 H2O, lower trop moisture
12.0 7 832 H2O, dirty window
Midwave H2O & O3
11.0 8 907 window
9.7 9 1030 O3, strat ozone
7.4 10 1345 H2O, lower mid trop moisture
7.0 11 1425 H2O, mid trop moisture
6.5 12 1535 H2O, upper trop moisture
Weighting Functions
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High
Mid
Low
ABC
A
B
C
line broadening with pressure helps to explain weighting functions MODIS IR Spectral Bands
MODIS
ABC
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CO2 channels see to different levels in the atmosphere
14.2 um 13.9 um 13.6 um 13.3 um
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H
e
i
g
h
t
A
b
s
o
r
p
t
i
o
n
Wavenumber Energy Contribution
line broadening with pressure helps to explain weighting functions
- k u (z)
(z ) = e
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Tau
H
e
i
g
h
t
100% 0 =dTau/dHt
Wet Atm.
Dry Atm.
Dry
Wet
Moderate Moderate
For a given water vapor spectral channel the weighting function depends on the
amount of water vapor in the atmospheric column
CO2 is about the same everywhere, the weighting function for a given CO2
spectral channel is the same everywhere
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Improvements with Hyperspectral IR Data
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GOES-12 Sounder – Brightness Temperature (Radiances) – 12 bands
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Characteristics of RTE
* Radiance arises from deep and overlapping layers
* The radiance observations are not independent
* There is no unique relation between the spectrum of the outgoing radiance
and T(p) or Q(p)
* T(p) is buried in an exponent in the denominator in the integral
* Q(p) is implicit in the transmittance
* Boundary conditions are necessary for a solution; the better the first guess
the better the final solution
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Profile Retrieval from Sounder Radiances
ps
I = sfc B(T(ps)) (ps) - B(T(p)) F(p) [ d(p) / dp ] dp .
o
I1, I2, I3, .... , In are measured with the sounder
P(sfc) and T(sfc) come from ground based conventional observations
(p) are calculated with physics models (using for CO2 and O3)
sfc is estimated from a priori information (or regression guess)
First guess solution is inferred from (1) in situ radiosonde reports,
(2) model prediction, or (3) blending of (1) and (2)
Profile retrieval from perturbing guess to match measured sounder radiances
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Example GOES Sounding
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GOES Sounders –Total Precipitable Water
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GOES Sounders –Lifted Index Stability
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Sounder Retrieval Products ps I = (sfc) B(T(ps)) (ps) - B(T(p)) F(p) [ d(p) / dp ] dp .
o
Direct
brightness temperatures Derived in Clear Sky
20 retrieved temperatures (at mandatory levels)
20 geo-potential heights (at mandatory levels)
11 dewpoint temperatures (at 300 hPa and below)
3 thermal gradient winds (at 700, 500, 400 hPa)
1 total precipitable water vapor
1 surface skin temperature
2 stability index (lifted index, CAPE) Derived in Cloudy conditions
3 cloud parameters (amount, cloud top pressure, and cloud top temperature) Mandatory Levels (in hPa)
sfc 780 300 70
1000 700 250 50
950 670 200 30
920 500 150 20
850 400 100 10
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Microwave RTE
Lectures in Brienza
19 Sep 2011
Paul Menzel
UW/CIMSS/AOS
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Radiation is governed by Planck’s Law
c2 /T
B(,T) = c1 /{ 5 [e -1] }
In microwave region c2 /λT << 1 so that
c2 /T
e = 1 + c2 /λT + second order
And classical Rayleigh Jeans radiation equation emerges
Bλ(T) [c1 / c2 ] [T / λ4]
Radiance is linear function of brightness temperature.
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19H Ghz 83
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10 – 11 um 84
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Microwave Form of RTE atm
ps 'λ(p) ref atm sfc
Isfc = ελ Bλ(Ts) λ(ps) + (1-ελ) λ(ps) Bλ(T(p)) d ln p
λ o ln p
ps 'λ(p)
Iλ = ελ Bλ(Ts) λ(ps) + (1-ελ) λ(ps) Bλ(T(p)) d ln p
o ln p
o λ(p) __________
+ Bλ(T(p)) d ln p sfc
ps ln p
In the microwave region c2 /λT << 1, so the Planck radiance is linearly proportional to the
brightness temperature
Bλ(T) [c1 / c2 ] [T / λ4]
So
o λ(p)
Tbλ = ελ Ts(ps) λ(ps) + T(p) Fλ(p) d ln p
ps ln p
where
λ(ps)
Fλ(p) = { 1 + (1 - ελ) [ ]2 } .
λ(p) 85
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Transmittance
(a,b) = (b,a)
(a,c) = (a,b) * (b,c)
Thus downwelling in terms of upwelling can be written
’(p,ps) = (ps,p) = (ps,0) / (p,0)
and
d’(p,ps) = - d(p,0) * (ps,0) / [(p,0)]2
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WAVELENGTH
FREQUENCY
WAVENUMBER
cm
m
Å
Hz
GHz
cm-1
10-5 0.1 1,000 Near Ultraviolet (UV)
3x1015
4x10-5 0.4 4,000 Visible
7.5x1014
7.5x10-5 0.75 7,500 Near Infrared (IR)
4x1014
13,333
2x10-3 20 2x105 Far Infrared (IR)
1.5x1013
500
0.1 103 Microwave (MW)
3x1011
300
10
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89
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Microwave spectral bands
23.8 GHz dirty window H2O absorption
31.4 GHz window
60 GHz O2 sounding
120 GHz O2 sounding
183 GHz H2O sounding
90
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23.8, 31.4, 50.3, 52.8, 53.6, 54.4, 54.9, 55.5, 57.3 (6 chs), 89.0 GHz 91
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II II I |I I ATMS Spectral Regions
Microwave Spectral Features
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AMSU
23.8
31.4
50.3 GHz
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AMSU
23.8 dirty
window
31.4 window
50.3 GHz
atm Q
warms
BT
94
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Tb = s Ts m + m Tm + m rs m Tm
Tb = s Ts (1-m) + m Tm + m (1-s) (1-m) Tm
So temperature difference of low moist over ocean from clear sky over ocean is given by
ΔTb = - s m Ts + m Tm + m (1-s) (1-m) Tm
For s ~ 0.5 and Ts ~ Tm this is always positive for 0 < m < 1
Low mist over ocean
rs , s
m , m
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AMSU
52.8
53.6
54.4 GHz
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AMSU
54.4
54.9
55.5 GHz
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Spectral regions used for remote sensing of the earth atmosphere and surface from
satellites. indicates emissivity, q denotes water vapour, and T represents temperature. 98