Unit1 Lecture2

Earth Science 1111: 

Earth Science 1111 Lecture 2: Electromagnetic Radiation and the Global Energy Budget

Means of Energy Transfer: 

Means of Energy Transfer Energy can be transferred from one area to another using three means: Conduction Convection Radiation

Conduction: 

Conduction Conduction is the transfer of energy directly from molecule to molecule This mode of energy transfer is most important in solids An example of this is a metal spoon left in a pot on the stove. Energy will travel slowly up the spoon’s handle. It is important in the atmosphere only in the “contact layer”, or the lowest 2 mm of the air.

Convection: 

Convection Convection is the means of energy transfer through fluid motions Most important in fluids, such as liquids and gases Hot fluid rising and cold fluid sinking in a pot of boiling water is an example Convection is very important in the troposphere (thunderstorms are a form of convection).

Radiation: 

Radiation Radiation is the transfer of energy by electromagnetic waves, which can travel through empty space The earth receives the majority of its energy from the sun, and it is received by radiation All objects having a temperature above absolute zero emit electromagnetic radiation Electromagnetic radiation can be viewed either as waves or particles (photons)

Electromagnetic Waves: 

Electromagnetic Waves Wavelength refers to the distance between one part of a wave (peak) and the same part of the next wave There is a huge range of possible wavelengths from the very short gamma rays (ten-millionth of a millimeter) to the very long radio waves (thousands of meters)

Electromagnetic Spectrum: 

Electromagnetic Spectrum FIGURE 2.2, PAGE 19

Electromagnetic Spectrum: 

Electromagnetic Spectrum The complete set of all possible wavelengths, from radio to gamma rays, is called the electromagnetic spectrum Climatologically important radiation is between 0.1 to 100 μm The human eye is sensitive to a small portion between 0.4 μm (violet) and 0.7 μm, called visible light Wavelengths immediately shorter than violet light are called ultraviolet, and wavelengths immediately longer than red are called infrared

Radiation Laws: 

Radiation Laws The amount of radiation emitted by an object at a particular wavelength depends on the temperature of that object. The basic law of radiation that governs the amount of radiation emitted by an object at a particular wavelength is called Planck’s Law.

Black Bodies: 

Black Bodies If an object were to emit precisely the amount of energy specified by Planck’s Law, that object is referred to as a black body. Black bodies are theoretical. The emissivity is the degree to which that object behaves like a perfect black body. Solids usually have emissivities between 0.9 and 1.0. Gases don’t behave like black bodies. They radiate only at specific wavelengths.

Planck Curves: 

Planck Curves FIGURE 2.3, PAGE 20

Wien’s Law: 

Wien’s Law The wavelength of maximum emission (where the peak in the Planck curve exists for that object) is given by Wien’s Law: Λmax = 2897 T Since temperature is in the denominator, an increase in temperature results in a decrease in the maximum wavelength emitted.

Stefan-Boltzmann Law: 

Stefan-Boltzmann Law The total area underneath the Planck curve represents the total energy emitted by that object. The total energy emitted by an object is given by the Stefan-Boltzmann Law: E = σT4 σ is a constant value Note that the total energy emitted is proportional to the temperature to the 4th power! If temperature were to double, the total energy emitted would be 16 times greater!

Kirchhoff’s Law: 

Kirchhoff’s Law At a particular wavelength, a good emitter is also a good absorber. Ozone is a good absorber of ultraviolet light. Theoretically, it would also emit ultraviolet light as well. However, the temperature of ozone doesn’t allow for the emission of ultraviolet light, so it is simply absorbed and not re-emitted by ozone.

Planck Curves: 

Planck Curves FIGURE 2.3, PAGE 20

Radiation Laws: 

Radiation Laws Planck’s Law, Wien’s Law, and Stefan-Boltzmann’s Law all relate to black bodies. The emissivity must be included in each equation for any object that is not a black body.

Solar Radiation: 

Solar Radiation The Sun has a surface temperature and Planck curve of an object at 5800 Kelvin Kelvin temperature is equal to the temperature in Celsius + 273 The peak color (Wien’s Law) is blue-green for the Sun, but we see yellow because of a variety of factors (eye sensitivity, interactions with the atmosphere, and the shape of the curve is thicker on the yellow end.

Solar Radiation: 

Solar Radiation We receive only 4.5 x 10-10 of the total energy emitted by the Sun, but this energy drives the climate system 99% of the solar radiation emitted has a wavelength between 0.15 and 4 μm. 9% is in the ultraviolet band 45% is visible light 46% is infrared

Solar Constant: 

Solar Constant The solar constant is the amount of energy passing in a unit time through a unit surface perpendicular to the Sun’s rays at the outer edge of the atmosphere at the average Earth-Sun distance. The solar constant is about 1372 Watts per square meter. The Sun is a very stable star, but the solar “constant” is known to vary on a number of time scales due to solar evolution, sunspots, etc.

Earth’s Orbit of the Sun: 

Earth’s Orbit of the Sun The Earth revolves around the Sun in an elliptical orbit in the period of one year. The eccentricity is a measure of departure from a perfect circle. For the Earth, the eccentricity is small. The eccentricity changes very slowly over time, but has a very minor impact on climate over our lifetimes. We are closest to the Sun in January (perihelion), and farthest from the Sun in July (aphelion). Obviously, eccentricity is not what drives the seasons.

The Earth’s Axis of Rotation: 

The Earth’s Axis of Rotation The Earth makes one complete spin about its axis over the course of 24 hours. The Earth’s axis is tilted 23 ½˚ relative to the plane that we revolve around the Sun. Because of the tilt, the duration of daylight and the height of the Sun in the sky changes during the course of a year.

The Earth’s Orbit: 

The Earth’s Orbit FIGURE 2.4, PAGE 22

The Earth’s Tilted Axis: 

The Earth’s Tilted Axis On the summer solstice, the Sun’s rays are striking the Tropic of Cancer (23 ½˚ North Latitude) directly. On the winter solstice, the Sun’s rays are striking the Tropic of Capricorn (23 ½˚ South Latitude) directly. On the vernal and autumnal equinoxes, the Sun’s rays are striking the Equator directly. It is because of the tilt of the Earth’s axis, and the resultant change in sunlight, that we experience seasons.

INSOLATION: 

INSOLATION FIGURE 2.5, PAGE 22

Insolation: 

Insolation Insolation is an acronym that stands for INcoming SOLar radiATION. The North Pole gets more insolation than any other latitude on the summer solstice. The North Pole gets NO direct insolation between the equinoxes The Equator receives a consistently high amount of insolation throughout the course of a year. This spatial and temporal variation is vital for creating our climate system.

Radiation/Atmosphere Interactions: 

Radiation/Atmosphere Interactions Radiation passing through the atmosphere can only have one of three things happen: Scattering: When radiation interacts with small particles, the direction of travel for the photon changes (reflection is a special case of scattering where the particle is BIG and there is a complete change in direction). Transmitted: The photon passes through unimpeded. Absorbed: The photon is captured by an atom or molecule in the atmosphere, causing the energy of that atom or molecule to increase.

Scattering: 

Scattering The direction of scattering can be basically broken down into two directions: Upward: unless it is scattered back down again, the photon is lost to space. Downward: the photon is still in the climate system and able to interact later on with the climate system.

Rayleigh Scattering: 

Rayleigh Scattering When the scattering particles are small compared to the wavelength of radiation. This is the case for air molecules. The amount of scattering is inversely proportional to wavelength (shorter wavelengths are scattered more). Scattering of blue light is 10 times greater than the scattering of red light. Rayleigh scattering is why the daytime sky is colored blue. For an evening red sky, the longer path length through the atmosphere results in most of the wavelengths being scattered out, including red being scattered out slightly.

Mie Scattering: 

Mie Scattering When the size of the particles are comparable to the wavelength of radiation. Cloud water droplets and pollution particles cause Mie scattering. All wavelengths are scattered in the same manner, which results in a light blue or greyish sky

Scattering: 

Scattering FIGURE 2.6, PAGE 24

Multiple Scattering: 

Multiple Scattering A photon may be scattered numerous times. Between or within clouds is a prime area for multiple scattering. Whiteout pertains to the condition where low clouds and snow cover results in multiple scattering which doesn’t allow the horizon to be distinguished.

Albedo: 

Albedo A measure of the amount of scattering and reflection of radiation. The textbook equation (2.10) has an error. Albedo is the amount of shortwave energy sent up divided by the total amount coming down (it represents the percentage reflected). Planetary albedo is the albedo at the top of the atmosphere for the entire globe (about 30%). Surface albedo is the albedo for a local area which is highly dependent upon the nature of the surface (significant for local climates).

Absorption: 

Absorption Absorption of shortwave radiation is a relatively small amount compared to the amount of scattering. The atmosphere is transparent to the visible wavelengths. Strong absorption of ultraviolet radiation happens with ozone and oxygen in the stratosphere (ozone layer). This leads to a strong heat source (why the stratosphere has temperature increasing with height). Only about 18% of the insolation entering the top of the atmosphere is absorbed by the atmosphere (most UV). Clouds absorb a remaining 2% for a 20% total absorption.

Energy at the Surface: 

Energy at the Surface Direct radiation are those photons that were not scattered (direct solar beam). Diffuse radiation are those photons that have been scattered (“skylight”). Both direct and diffuse radiation behave in the same way energetically at the surface. A maximum amount of insolation is received at the surface in the subtropical deserts (high sun angle combined with a lack of cloud cover) A minimum amount of insolation is received at the poles (low sun angle with a high albedo surface). The radiation that reaches the surface must either be reflected or absorbed.

Insolation at the Surface: 

Insolation at the Surface FIGURE 2.7, PAGE 29

Surface Albedo: 

Surface Albedo The percentage of insolation reflected by the Earth’s surface Varies with the type of surface Not wavelength dependent (single value) Natural surfaces have a 10-25% albedo Water albedo depends on sun angle Snow has a high albedo. Clean snow will last longer than dirty snow. Global average is 15% (water dominates the surface of the Earth)

Albedos and Emissivities at Surface: 

Albedos and Emissivities at Surface TABLE 2.1, PAGE 29

Surface Absorption: 

Surface Absorption Because most surfaces have a low albedo, the majority of insolation reaching the surface will be absorbed. Of the entire insolation that reaches the top of our atmosphere, 50% will be absorbed by the surface (recall 20% will be absorbed by the atmosphere). The remaining 30% is reflected back out to space is results in our planetary albedo of 30%.

Short-wave Budget: 

Short-wave Budget FIGURE 2.1(A), PAGE 18

Surface Heating: 

Surface Heating The primary source of heating for the lower atmosphere is the underlying surface of the Earth With the heat source from beneath, Vertical air motions are established (convection) Vertical air motions are critical for cloud formation Different surface types can result in regional weather and climate differences

Terrestrial Radiation: 

Terrestrial Radiation Absorption of solar radiation at the surface leads to heating Almost all surface types have an emissivity greater than 0.9, which means the Earth’s surface is fairly close to a blackbody for long wave (infrared) radiation.

The Atmosphere: 

The Atmosphere The atmosphere is not a blackbody The values of absorptivity and emissivity vary depending on wavelength Each individual gas will absorb only specific wavelengths, and this appears as spectral absorption lines on the radiation curves Lines of spectral absorption can be grouped together to form absorption bands The location of the bands depend on the atomic structure of the gas Major absorbers of infrared are water vapor, carbon dioxide, and to a lesser extent, ozone and methane

The Atmosphere and Infrared: 

The Atmosphere and Infrared At wavelengths greater than 3 μm are almost all absorbed by one gas or another The exception is a narrow band between 8 μm and 11 μm, called the atmospheric window Therefore, most long wave radiation emitted by the surface is absorbed by the atmosphere

Long-Wave Budget: 

Long-Wave Budget FIGURE 2.1(B), PAGE 18

Atmospheric Heating: 

Atmospheric Heating Atmospheric heating is the total short wave radiation absorbed directly by sunlight, plus the absorbed long wave radiation emitted by the Earth’s surface This atmospheric heating causes the atmosphere to also emit long wave radiation Upward: Outgoing long-wave radiation that is eventually lost to space Downward: Incoming long-wave radiation can be eventually received by the surface again

Clouds and Infrared: 

Clouds and Infrared Clouds close the atmospheric window and absorb those extra infrared bands Cloudy nights don’t cool down as rapidly as clear nights because the clouds act like an insulating blanket and prevent outgoing long-wave radiation from leaking out to space

“Greenhouse Effect”: 

“Greenhouse Effect” Solar energy passes mostly unimpeded through the atmosphere and is absorbed at the surface Outgoing long-wave radiation off the surface is absorbed by the atmosphere A portion of the absorbed long-wave radiation is sent back down to the surface The surface has a temperature 30˚ C warmer than it normally would have This process is termed the “greenhouse effect”

“Greenhouse”?!: 

“Greenhouse”?! Greenhouses are certainly warm places on sunny days However, greenhouses are not warm because of the trapping of infrared radiation A greenhouse is warm because convection is inhibited by the glass (the warm air is stuck inside the greenhouse) Therefore, the term “greenhouse effect” is a misnomer and does not apply to the atmosphere

Greenhouse Effect vs. Global Warming: 

Greenhouse Effect vs. Global Warming The “greenhouse effect” is a good thing, otherwise we would be an ice planet The terms “greenhouse effect” and “global warming” do not mean the same thing Global warming is the concern that, by increasing the gases that trap infrared radiation in our atmosphere, we will increase the average surface temperature of the Earth

Global Radiation Budget: 

Global Radiation Budget Over long time scales, and over the entire globe, the amount of incoming short-wave radiation equals the amount of outgoing long-wave radiation This results in a global radiation balance Some minor variations can occur, and this drives climate variations

Smaller-Scale Radiation Budgets: 

Smaller-Scale Radiation Budgets Within the long-term global energy balance, there are local and short-term imbalances These imbalances drive the weather and climate system In calculating radiation budgets, we are concerned with only two levels: Top of the atmosphere Surface of the Earth

Top of the Atmosphere Budget: 

Top of the Atmosphere Budget The radiation budget for the top of the atmosphere looks like this: Q = K – Lup Net Radiation = absorbed - outgoing long-wave short-wave from below below We only have a balance when Q is zero (when K = Lup)

Variations in K: 

Variations in K The absorbed solar radiation beneath the top of the atmosphere depends on Albedo Total amount of solar radiation incoming The poles receive less because of a low sun angle and high albedo There is a considerable difference between the Equator and the poles

Average Short/Long-Wave Budget by Latitude: 

Average Short/Long-Wave Budget by Latitude FIGURE 2.10, PAGE 33

Variations in L: 

Variations in L There is less of a difference between the amount of outgoing long-wave radiation at the Equator and at the poles The atmosphere and oceans are very efficient at moving energy from the tropics to the polar regions, and this makes the difference less

Seasonal Variations in K & L: 

Seasonal Variations in K & L Because of snow and ice, there is a seasonal variation in surface albedo, especially in the Northern Hemisphere Because K is less in the winter due to a higher albedo, L is also lower because of lower surface temperatures Changes in albedo also occur in the mid-latitudes and tropical regions due to changes in vegetation, but this is a smaller variation

Seasonal Variations in K & L: 

Seasonal Variations in K & L FIGURE 2.11, PAGE 34

Surface Energy Budget: 

Surface Energy Budget The surface energy budget also experiences a long-term global balance with local and short-term variations The energy budget for the surface is: Q = Kdown – Kup + Ldown – Lup net radiation = incoming – outgoing + incoming – outgoing short-wave short-wave long-wave long-wave

Surface Energy Budget: 

Surface Energy Budget The values in the equation can change over many time scales, including daily The solar radiation term, which depends on latitude, season, time of day, clouds, etc., is the most variable of all the terms The long-wave term is less variable, but it does change with the surface temperature and humidity values, which can change by the wind At night, Q (net radiation) is most likely to be negative (only L is active), depending on clouds and winds Other factors need to be included later to balance it out

Diurnal Variation in Surface Budget: 

Diurnal Variation in Surface Budget FIGURE 2.12, PAGE 35