Understanding the Greenhouse Effect

Transcription

1 EESC V2100 The Climate System spring 200 Understanding the Greenhouse Effect Yochanan Kushnir Lamont Doherty Earth Observatory of Columbia University Palisades, NY 1096, USA

2 Equilibrium (Effective/Emission) Temperature of Earth In equilibrium conservation of energy implies that Solar radiation absorbed = planetary radiation emitted T e = S 0 (1 α p ) σ If Earth had no IR absorbing gases in its atmosphere shortwaves longwaves πr p 2 (1 α p )S 0 = πr p 2 σt e On Earth (S 0 =1367, α p =0.3) we have: T e =255 K or -18 C

3 Composition of the Atmosphere percent concentration by volume in dry air

4 Absorption of Photons by Gas When photons hit a volume of gas part of the energy is absorbed (outside of scattering and transmission). Absorption implies that the photons entire energy is transformed to one of forms of the gas internal energy: E absorbed = E rotational + E vibrational + E electronic + E translational E rotational, E vibrational, and E electronic can increase only in quantized increments according the the properties of the gas. Different gases absorb in different wavelength bands: a photon will not be absorbed if its energy (i.e. frequency) does not correspond to the steps between the gas energy levels. E translational corresponds to the gross movement of the gas molecules and its temperature and is the only energy that can change in non-quantized increments. It is responsible for collisions between molecules and doppler effects, both of which contribute to the broadening of spectral absorption or emission bands. The structure of water, CO 2 and other polyatomic molecules allows for the absorption and emission of IR radiation.

5 Absorption of Radiation in the Earth s Atmosphere Various atmospheric constituents absorb electromagnetic radiation. The important absorbing gases and their absorption efficiency as a function of wavelength is shown on the right. Note the strong oxygen absorption in the UV part of the spectrum and that of water vapor, which absorb effectively in large sections of the IR wavelength range. CO 2 absorption band overlaps a gap in the water vapor band (referred to as window ) hence its importance in the climate system. Note also that non of the gases are truly blackbodies.

6 Atmospheric Normalized Spectra Absorption Ground level absorption Fraction of energy emerging at the top of the atmosphere

7

8 Atmospheric Absorption - Greenhouse effect The greenhouse effect is caused by IR absorbing gases (mainly water vapor) in the Earth s atmosphere, which reduce the efficiency of the climate system to lose heat to space. A radiative equilibrium must be achieved, but because of the absorption, the ground has to warm to a temperature higher than the emission temperature for the outer atmosphere to achieve a balance with the incoming shortwave radiation. The following slide illustrates the greenhouse effect. It shows the Earth climate system receiving solar (shortwave) radiation in the amount S 0 / (averaged around the globe). Aside from the albedo effect due mainly to clouds we assumed that the atmosphere is entirely transparent to the incoming solar flux. An absorbing layer in the troposphere traps all the longwave radiation emitted by the planet. In radiative equilibrium this layer emits energy towards the ground and outer space (the same radiative flux is emitted both ways) warming the surface and maintaining radiative balance with space.

9 Effect of Atmospheric Absorption 1 w/o absorption w. absorption S 0 α S 0 στ g S 0 α S 0 H=σΤ a atmosphere atmosphere IR absorbing layer στ g (1-α) absorption of solar ground S 0 emission of IR (1-α) S 0 ground H στ g = (1-α) S 0 στ a στ g = = (1-α) S 0 S (1-α) 0 + στ a

10 Effect of Atmospheric Absorption 2 w/o absorption w. absorption στ g = (1-α) S 0 = 239 Wm -2 στ a = (1-α) S 0 The Greenhouse Effect στ g = στ g = S (1-α) 0 + στ a (1-α) S 0 2 Τ g = (1-α) S 0 σ Τ g = (1-α) S 0 2σ W/O absorption: T g =255 K or -18 C W. absorption T g =303.5 K or C

11 Radiative Transfer in the Atmosphere In reality the situation is much more complex. Solar energy is absorbed in the atmosphere, and IR absorption happens continually through the troposphere. The surface looses heat not only through radiation, but also through convection (the transfer of heat through turbulent fluid motion), which moves warm and moist air masses from the surface up causing surface sensible and latent heat loss (more later). The globally averaged picture of the Earth radiation balance is depicted on the left. Based on this figure: greenhouse effect= =155 Wm -2 T g =(390/σ) 1/ =288 K=15 C

12 Earth Radiation Budget from Space: the Spatial Pattern

13 Incoming Solar (Shortwave) at TOA December March June September

14 Incoming Solar (Shortwave) at TOA Global average 32 Wm-2 [S 0 /] Longitudinal (zonal) symmetry with uniform variation with latitude. Hemispheric symmetry during equinox seasons with radiation decreasing uniformly from equator to pole. Polar regions receive large amounts of radiation during summer (more than the equator, even during equinox) because of the length of the day. Present climate perihelion during winter is causing larger overall radiation levels during December than during June.

15 Reflected Solar at TOA December March June September

16 Reflected Solar (Shortwave) at TOA Global average 235 Wm-2 [(1-α)S 0 /]. Longitudinal variations with non-uniform variation in latitude. Weak hemispheric symmetry during equinox seasons. Polar regions reflect largest amount of radiation during summer, a result of high insolation but also of high reflectivity. Variations in longitude and non-uniform latitudinal behavior reflect changes in planetary albedo.

17 Planetary Albedo December March June September

18 Planetary Albedo Global average ~0.3. The albedo field reflects the properties of the surface and the cloud distribution and brightness (see following two slides). Over the oceans albedo is a function of the Sun s zenith angle (deviation from the perpendicular). The larger the angle the larger is the albedo. Albedo values are large around the poles - a result of the reflective lowclouds/ice/snow cover and the low angle of the Sun over the oceans. Over land albedo is high in dry (desert) areas and low in forrest regions - however, some of these forrest regions are also regions of deep and highly reflective clouds, which mask the surface properties to create high albedo. Over the tropical oceans there are narrow regions of high albedo flaked by large regions of low albedo - evidence for narrowly confined cloud regions flaked by vast ocean areas with little cloudiness.

19 Earth s Surface Properties as seen from Space

20 Global Rainfall - a Proxy for Clouds

21 Net Shortwave (Solar) Radiation December March June September

22 Net Shortwave Radiation at TOA Global average 235 Wm-2 [(1-α)S 0 /]. Results from the combination of length-of-day, solar zenith angle, and local planetary albedo. The poles are now regions of minima (high albedo) and the tropics, regions of maxima, particularly over the vast subtropical areas, which are relatively cloud free and have low albedo. The distribution of net shortwave is close to being hemispherically symmetric with small variation in maximum with the season.

23 Outgoing Longwave (IR) at TOA December March June September

24 Outgoing IR (longwave) at TOA also referred to as OLR (outgoing longwave radiation) Global average 235 Wm-2 [(1-α)S 0 /] - globally approximately balancing net shortwave over the course of the year (imbalances may occur over limited periods due to the ability of the oceans to store heat away from the surface). Outgoing longwave is more uniformly distributed with latitude than net shortwave, reflecting the planet s relatively uniform temperature - a direct result of the dynamical climate system, which works to distributes the heat received from the sun around the globe. Regions with relatively little clouds and dry air upper atmosphere (subtropics) emit larger amount of longwave to space. Regions of deep clouds (Southeast Asia, South America, Africa) display minimum emission.

25 Net Incoming Radiation December March June September

26 Net Incoming Radiation Global average - zero (or close to it) with positive overall in the summer hemisphere and negative in the winter hemisphere. The polar regions are always negative where in the winter OLR is uncompensated by shortwave radiation. Note that both the maximum and minimum values are somewhat larger than 100 Wm -2. Note that over the deserts (such as the Sahara) the net is ~zero: In dry land region the only mechanism that can balance the incoming shortwave (after reflection, see albedo slide) is longwave cooling. The annual mean picture looks more like the equinox states with excess of incoming radiation in the tropical and subtropical regions and a deficit thereof in the middle and high latitudes. The latitudinal distribution results from the relatively strong decrease of net shortwave with latitude on both sides of the equator compared to the weak gradient in OLR. Radiation Wm -2

27 Surface vs. TOA Longwave From surface temperature data we can calculate the surface outgoing longwave radiation (using the Stefan-Boltzman law) assuming emissivity* of 0.95 Annual mean surface outgoing IR Compare this with the outgoing logwave radiation at the top of the atmosphere... Annual mean TOA outgoing IR * emissivity: Natural surfaces are not perfect black bodies. The absorb and emit a fraction of the amount predicted by the Stefan-Boltzman Law. The ratio between actual and predicted emission is the emissivity.

28 Greenhouse Effect The difference between the longwave radiation from the Earth s surface and OLR is the greenhouse effect. Note the strong GH effect in areas which are dominated by deep tropical clouds that precipitate a lot (above). These clouds reach high into the atmosphere (more than 10 Km) where the temperature is low. These clouds tops radiate longwave into space at these low temperatures, while the surface underneath is warm and its emitted longwave radiation is trapped in the cloudy atmosphere.

29 Earth Radiation Budget: the Role of Clouds

30 Cloud Forcing: Longwave By comparing total outgoing IR to outgoing IR in locally cloud free (or clear-sky) conditions we can evaluate the trapping effect (forcing) of clouds. Annual mean total outgoing IR In the longwave band clouds warm the climate system Annual mean clear-sky outgoing IR

31 Cloud Forcing: Shortwave Annual mean total reflected solar Shortwave cloud forcing is determined in a similar way, by comparing total and clear-sky reflected solar. In the shortwave band clouds cool the climate system Annual mean clear-sky reflected solar

32 Cloud Forcing: The Net Effect In the net, clouds slightly cool the climate system. The also provide for reduced zonal symmetry over the oceans. Annual mean total net radiation Annual mean clear-sky reflected solar

33 Cloud Forcing: Global Summary The Table below presents the globally averaged effect of clouds on the radiation balance of the planet as measured by NASA ERBE satellites.

34 Summary The greenhouse effect results from atmospheric absorption of IR (longwave) radiation emitted from the Earth s surface. IR (and shortwave) absorption happens in broad spectral bands related to molecular motion and energy levels. The major constituent of the atmosphere, Nitrogen (78.01% by volume in dry air) is transparent to both short- and longwave radiation. Oxygen (20.95% by volume) absorbs mainly shortwave (UV) radiation. It is the minor atmospheric constituents, mainly water vapor (about 0.33% of total atmospheric mass and 0.8% by volume) and CO 2 (0.035% by volume) that efficiently absorb the longwave radiation from the Earth s surface. Longwave absorption by the atmosphere strongly reduces the efficiency of Earth s cooling to space, forcing its surface to overheat in order to balance the incoming solar radiation. The end result is a higher surface temperature than the emission temperature of 255 K). At present conditions, the earth has to emit 350 Wm -2 to allow 235 Wm -2 to escape to space and balance the 235 Wm -2 received from the sun. This results in an average surface temperature of 288 K or 15 C