(1 of 14) Further Reading: Chapter 04 of the text book Outline - global radiative energy balance - insolation and climatic regimes - composition of the atmosphere
(2 of 14) Introduction Last time we discussed radiation and distinguished between solar (or shortwave) and thermal (or longwave) radiation In addition, we discussed how temperature affects the intensity and wavelength of radiation Today we want to look at how radiation can affect the temperature of a body In particular, we want to: Discuss how radiation from the sun determines the temperature of the earth s surface based upon a very simple model for the global earth system Look at how solar radiation varies with latitude and time of year We are also going to use this lecture to introduce some basic concepts used in atmospheric sciences to describe the composition and state of the atmosphere
(3 of 14) Thermal Equilibrium Amount of energy absorbed equals the amount of energy released (or radiated away) If the input of energy exceeds the output, energy is added to the system and it will heat up If the input of energy is less than the output, energy is removed from the system and it will cool down Hence, thermal equilibrium implies that there is no net heating of the system
(4 of 14) Global Radiation Energy Balance The flux arriving at the top of the atmosphere is constant This constant input of energy goes to heating the earth Heating of the earth raises the temperature As the temperature of the earth increases, this increases amount of longwave radiation it gives off Eventually, the temperature is just right so that the amount of longwave radiation given off exactly balances the incoming solar radiation
(5 of 14) Solar Constant Energy flux arriving at the top of the atmosphere measured perpendicular to Sun s rays at mean Earth-Sun distance Equal to 1370 W/m^2 (Watts per square meter)
(6 of 14) Total Incident Solar Energy Energy flux arriving at the top of the atmosphere Es = 1400 W/m^2 (approx. to the solar constant) Interception area of Earth = pi*r^2 = 129,000,000,000 m^2 Ein = 1.8 x 10^17 Watts
(7 of 14) Total Solar Energy Absorbed Not all energy reaching the earth s atmosphere is absorbed About 33% is reflected back to space Clouds Ice and Snow Plants 33% reflected away Esurf = 1.2 x 10^17 Watts
(8 of 14) Radiation Energy Balance Esurf(incoming) = Esurf(outgoing) Esurf(outgoing) = σt^4 times Area of Earth T^4 = (1.2 x 10^17/Area)/σ T(earth) = 253 K (degrees Kelvin) Water freezes at 273 K; The actual T(earth) is 290 K That is because of the natural greenhouse effect
(9 of 14) Insolation So far we have been talking about global balances; this ignores regional and seasonal effects For a particular point, as opposed to the globe, we refer to insolation For a given point, insolation depends on the angle of the incoming sunlight and the duration of sunlight Angle of sun (solar zenith angle): if this is large, the energy of the sun is spread over a larger area and insolation goes down Duration: the longer the point is exposed to the sun, the more radiation it receives, hence its insolation goes up For daily insolation, maximum occurs at the poles during the solstice because of duration effects (i.e. the day is 24 hours long) However, it also has the lowest insolation during winter, hence it has a large annual temperature range
(10 of 14) Annual Insolation Over the year, however, equatorial regions have the highest insolation because of the consistently low solar zenith angles Because of this, the tropics have the highest average annual temperatures The earth s tilt results in Substantial increase in mean insolation at high latitudes Slight decrease in insolation at low latitudes
(11 of 14) Climatic Regimes Areas of the globe defined based upon the annual insolation they receive (13 regimes) Explain first order control on climate Higher insolation and warmer climate at low latitudes (and vice versa) Climates at higher latitudes characterized by strong seasonality
(12 of 14) Radiation and Temperature Up until now we have looked at the general relationship between solar energy, the earth s orbit and global and latitudinal temperatures For instance, we know that in the tropics its warm and has a constant temperature through the seasons; for the polar regions, it is cool and the temperature changes throughout the seasons NOTE: this doesn t always hold. For instance London at 51N has cooler summers and warmer winters than Boston at 42N Now we want to discuss heterogeneity in temperatures. This deals with how energy interacts with the cryosphere, oceans, biosphere, lithosphere, and atmosphere at different points on the globe
Composition of the Atmosphere Starting with the atmosphere First, we are interested in how energy is transferred through the atmosphere This depends on its composition The earth is 128,000km across The atmosphere is gaseous envelope held close to earth by gravity 97% of atmosphere is within 30km of the surface (13 of 14) Originally the gases from the atmosphere seeped out from volcanoes 10% was CO2, 85% was H2O; No O2; it was also much warmer (100C) As the earth cooled, the H2O condensed into clouds, rain, oceans Photosynthetic organisms evolved, converting CO2 into O2 Evolution of land plants led to rapid conversion of CO2 into O2 Today we find most is N with O2 second The most highly variable constituent is H2O: from 0.25-2%; depends on temperature, location, dynamics
(14 of 14) State Variables Temperature: Average kinetic energy of air molecules Highly variable (-60 -> 50 C) Study is called thermodynamics Pressure: Average mass of air molecules above a given point Related to potential energy Study is called dynamics Humidity: Average number of water molecules in a give volume of air Water is one of the most important constituents of the atmosphere - influences energy balance, water balance, and dynamics Study is called hydrodynamics