2. What makes it go?Copyright 2014, David A. RandallThe Earth’s radiation budget: An “upper boundary condition” on the global circulationRadiation is (almost) the only mechanism by which the Earth can exchange energy with the rest of the Universe. The solar energy flux at the mean radius of the Earth's orbit is about 1365 W m-2. One way to get an intuitive grasp of this number is to imagine fourteen 100-Watt light bulbs per square meter. Another is to consider that 1365 W m-2 is equivalent to 1.365 GW km-2, which is the energy output of a large power plant for each square kilometer of area normal to the solar beam.The globally averaged top-of-the-atmosphere radiation budget is summarized in Table 2.1. These numbers are now known to three significant digits. The Earth’s albedo is near 0.30, independent of season; this number has been known to better than 10% accuracy only since the 1970s. The energy absorbed by the Earth isSabs= Sπa24πa2⎛⎝⎜⎞⎠⎟1−α( )=14S 1−α( )≅ 240 W m−2(annual mean)(1)Incident solar radiationAbsorbed solar radiationPlanetary albedoOutgoing longwave radiationBrightness temperature of the Earth341 W m-2239 W m-20.30239 W m-2255 KTable 2.1 Summary of the annually averaged top-of-the-atmosphere radiation budget, after Trenberth et al. (2009).! Revised Tuesday, January 21, 2014! 1An Introduction to the Global Circulation of the AtmosphereHere Sabsis the average absorbed solar energy per unit area. S is the “solar constant,” ais the radius of the Earth, and αis the planetary albedo. Note that this rate of energy absorption is referred to the total surface area of the Earth, i.e. 4πa2, rather than to the cross-sectional or “projected” area presented to the solar beam, which is four times smaller. The most important upper boundary condition on the global circulation of the atmosphere is the incident solar radiation, also called the insolation. It is determined by the energy output of the sun and the geometry of the Earth’s orbit (see Fig. 2.1), including the obliquity, the eccentricity, and the dates of the equinoxes, as well as the Earth’s rotation rate. These all vary over geologic time (e.g., Crowley and North, 1991).As a matter of common experience, the insolation varies both diurnally and seasonally. At a given moment, the insolation also varies strongly with longitude. Because a year is much longer than a day, the daily-mean insolation is (almost) independent of longitude, but it varies strongly with latitude in a way that depends on the season, as summarized in Fig. 2.2. As we move from the solar Equator to the summer pole, the insolation initially decreases, because at a given local time (e.g., local noon) the sun appears to be lower in the sky. In addition, however, the length of day increases at higher latitudes, and this obviously tends to make the daily-mean insolation increase. Near the poles, the length-of-day effect dominates, so that the insolation begins to increase again. That is why there is a minimum of the insolation in the mid-latitudes of the summer hemisphere.As discussed later, seasonal and, to a lesser extent, diurnal cycles are clearly evident in the circulation patterns. Around the time of the solstices, no insolation at all occurs near the Figure 2.1: March of the seasons. As the tilted Earth revolves around the sun, changes in the distribution of sunlight cause the succession of seasons. (From Imbrie and Imbrie, 1979.)! Revised Tuesday, January 21, 2014! 2An Introduction to the Global Circulation of the Atmospherewinter pole (the polar night1), but at the same time, near the summer pole, the daily mean insolation is very strong despite low sun angles, simply because the sun never sets (the polar day). As is well known, these effects arise from the sun-Earth geometry. In addition, the distance from the sun to the Earth varies with time of year, resulting in a few percent more globally averaged insolation in January than in July. The month of maximum insolation varies over geologic time. According to the astronomical theory of the ice ages, extensive glaciation is favored when the minimum insolation occurs during the Northern Hemisphere summer, because the Northern Hemisphere contains about twice as much land as the Southern Hemisphere (e.g., Crowley and North, 1991).As you probably already know, the infrared radiation emitted by the Earth is very nearly equal to the solar radiation absorbed by the Earth, i.e., it is about 240 W m-2. This near balance between absorbed solar radiation and emitted terrestrial radiation has been directly confirmed by analysis of satellite data. The balance is observed to hold within a few Watts per square meter, which is comparable to the uncertainty of the measurements.Figure 2.2: The seasonal variation of the zonally (or diurnally) averaged insolation at the top of the atmosphere. The units are W m-2.200 300 400 500 500 500 200 200 300 300 400 400 100 100 100 SOLAR DECLINA TION J F M A M J J A S O N D 0 80N 80S 60 40 20 20 40 60 ! Revised Tuesday, January 21, 2014! 3An Introduction to the Global Circulation of the Atmosphere1 Where the sun don’t shine.Fig. 2.3 shows aspects of the Earth's radiation budget as observed in the Earth Radiation Budget Experiment (ERBE; Barkstrom et al., 1989). The zonally averaged incident (i.e. incoming) solar radiation at the top of the atmosphere varies seasonally in response to the Earth's motion around the sun. The zonally averaged albedo, which is the fraction of the zonally averaged incident radiation that is reflected back to space, is highest near the poles, due to snow and ice as well as cloudiness, but it tends to have a weak secondary maximum in the tropics, again associated with high cloudiness there. The zonally averaged terrestrial radiation at the top of the atmosphere, also called the outgoing longwave radiation or OLR, has its maxima in the subtropics. It is relatively small over the cold poles, but it also has a minimum in the warm tropics, due to the trapping of terrestrial radiation by the cold, high tropical clouds, and by water