WHAT IS THE NATURE OF THE LIGHT THAT REACHES EARTH'S SURFACE?

Incident Radiation at The Top of The Atmosphere

Variations in Insolation Due to Earth-Sun Distance

Variations in Solar Insolation Due to Solar Zenith Angle

Key Points about Radiation that Reaches Earth's Atmosphere

Interactions Between Light and Matter in the Atmosphere

The Energy of an Atom

Absorption and Emission of Photons by Atoms

The Energy of a Molecule

Key Points about The Interaction of Light and Matter

Absorption in the Atmosphere

Basic Ozone Photochemistry

Atmospheric Scattering - Some Basics

Rayleigh Scattering and Mie Scattering

Single and Multiple Scattering

Key Points about Atmospheric Scattering

Incident Radiation at The Top of The Atmosphere

The Sun emits electromagnetic radiation in all directions, but we are only concerned with the amount of energy that reaches Earth's atmosphere. The amount of energy transferred to the atmosphere at a given position depends on where the Sun is relative to the local zenith, which is the direction directly overhead. The angle between the Sun and zenith is known as the solar zenith angle ( ), as illustrated in Figure 11. The heating effectiveness of solar radiation, called insolation, varies with latitude and time of day since these factors affect the solar zenith angle. We can see from Figure 11 how different solar zenith angles affect the amount of solar radiation that is transferred into the atmosphere. Clearly, during the day we get the most solar radiation at noon, when the Sun is highest in the sky, and the least at sunset, when the solar zenith angle is near 900.

Figure 11. Solar Zenith Angle???

Variations in Solar Insolation Due to Solar Zenith Angle

The other factor affecting the amount of incoming solar radiation at a given position is the changing solar zenith angle. As we saw in Figure 11, the lower the solar zenith angle (the higher the Sun is in the sky) the greater the incoming radiation. We know that the solar zenith angle changes daily as Earth rotates about the Sun. However, the induced diurnal changes in the ozone occur very high in the stratosphere, and at most levels the ozone does not respond to these very short-term changes. As a result, in ozone studies we are more interested in the longer-term changes.

The solar zenith angle also changes as a function of season. These variations drive the seasonal cycle in temperature and atmospheric circulation. Figure 13 shows the average distribution of solar insolation as a function of latitude and season. To understand this plot, we must first remember how Earth's tilt changes as it revolves around the Sun. During northern hemisphere summer, the northern hemisphere tilts towards the Sun, causing the Sun to be higher in the sky with a smaller zenith angle and longer amounts of daylight. Summer solstice, around June 22, marks the highest point of the Sun in the northern hemisphere, and the longest amount of daylight of the year. On this day the Sun is directly overhead at Tropic of Cancer ( 23.5o N latitude) at noon, and the Sun never sets anywhere north of the Arctic Circle at 66.5o N latitude. At the same time, the southern hemisphere is tilted away from the Sun, and is undergoing the shortest daylight period of the year with the highest solar zenith angle. Just as the region north of the Arctic circle remains in light all day, all regions south of the Antarctic circle ( 66.5o S latitude) remain in darkness. This is known as polar night. The situation is reversed at the winter solstice, which occurs around December 22. In Figure13 , polar night is indicated by the shaded regions, where there is no solar insolation. Notice the maximum solar insolation occurs at the North (South) Pole at the summer (winter) solstice, due to the exceptionally long days. At the equinoxes (September 22 and March 22) the Sun is directly over the equator. The Vernal, or spring equinox, marks the end of polar night at the North Pole, and the beginning of polar night at the South Pole, with the largest amount of solar insolation occurring at the equator. The opposite scenario occurs on the Autumnal, or fall equinox. Note that the length of daylight at all latitudes is exactly 12 hours on the equinoxes. If the amount of insolation were to be added over the year, we would find more insolation in the tropics when compared to high latitudes, leading to the temperature differential between the tropics and the poles.

Figure 13. Solar Insolation as a Function of Latitude and Season

Key Points about the Solar Radiation that Reaches Earth's Atmosphere:

a.	The amount of energy transferred to Earth's atmosphere from the Sun depends on the solar zenith angle (position of the Sun in the sky) with most energy transferred when the solar rays are most direct (smallest solar zenith angle).

b.	Radiation transfer depends on the Sun-Earth distance which varies about 3.2% during the year as the Earth moves on its elliptical orbit. Radiation intensity decreases as the square of the distance between the Sun and Earth.

c.	The solar zenith angle changes seasonally depending on the Earth's tilt toward or away from the Sun.

Interactions Between Light and Matter in the Atmosphere

Now that we know about the solar radiation coming into the atmosphere, we want to look at what happens to that radiation. Solar radiation that penetrates Earth's atmosphere may either be: 1) reflected back into space; 2) scattered to other parts of the atmosphere by atoms, molecules, cloud droplets, and aerosols; 3) or absorbed by atmospheric constituents (most notably ozone), increasing the energy of the absorbing materials. Absorption entails the disappearance of a photon, with all of the photon's energy being transferred to an atom or molecule. Scattering occurs when light interacts with an atom, molecule, or aerosol such that the light's direction of travel is changed, but its energy, and the energy of the scattering particle, remains the same. Radiation that reaches the surface is either reflected back into the atmosphere or absorbed by the surface. All of these processes involve radiative transfer, which is defined as the exchange of energy between photons and matter. In the following sections we are going to take a closer look at radiative transfer through absorption and scattering processes in the atmosphere. However, to understand these processes, we must first understand what the energy of an atom or molecule is, and how it can be changed.

Key Points about the Interaction of Matter and Light:

1. An electron in an atom at any instant of time is characterized by its energy.

2. There are only certain energies that an electron is allowed to have.

3.	If a photon comes by that just happens to have an energy that corresponds to the difference between two of the allowed energies, it may be absorbed. The electron winds up in a higher energy level and the photon ceases to be.

4.	Since electron energy levels are well defined, so are their differences. If you shine white light (with all visible wavelengths) through a sample containing these atoms, most of the photons that go in will come out. However, photons having certain wavelengths (energies) will be absorbed by some of the atoms and they won't come out.

5.	We call the pattern of light absorption as a function of wavelength the absorption spectrum of an atom.

6.	If an atom encounters a photon with enough energy to tear its electron free from the nucleus, the atom becomes ionized, and any excess energy is transferred from the photon to the free electron in the form of kinetic energy.

7.	Molecules have nuclear vibrational and rotational levels that add to the energy states of a molecule. Transitions between these states take less energetic photons than electronic transitions.

Absorption in the Atmosphere

Absorption of solar radiation in the atmosphere serves two purposes: 1) it is a heat source [click to radiative heating] for Earth-atmosphere system; and, 2) it prohibits high-energy radiation from reaching Earth's surface. Solar radiation at wavelengths shorter than about 0.31 m , which covers a portion of the ultraviolet region (UV-B) through the X-ray region, can be extremely dangerous to the biosphere if it is subjected to exposure for extended periods of time. Figure 16 shows how this potentially dangerous solar radiation is absorbed in the atmosphere. Specifically, it shows the altitude at which the radiation has been reduced by a factor of e (about 2.7) from its value at the top of the atmosphere and the species responsible for the absorption. For these wavelengths, nearly all of the radiation is absorbed before it reaches the surface. In fact, only a small portion of radiation of the least energetic of these wavelengths ever reaches the ground (Gordon- % of UV-B reaching the ground?).

Figure 16. Solar Radiation Absorption as a Function of Altitude

Solar radiation at wavelengths shorter than 0.1 m contains enough energy to photoionize

the atomic constituents of the far upper atmosphere (altitudes above about 90 km), including oxygen and nitrogen (O and N), and photodissociate molecular oxygen and nitrogen ( O2 and N2 ). We can recall that in the process of ionization ( photodissociation ), the excess energy is transferred from the photon to the free electron (atom) in the form of kinetic energy. This ultimately results in increased heat in the layer. More absorption takes place at the top of the layer so the temperature of the layer increases with altitude. This atmospheric layer is the thermosphere [click to temperature structure of the atmosphere], also known as the ionosphere because of the high percentage of ionized particles. The radiation responsible for maintaining the ionosphere is emitted primarily from the Sun's chromosphere, which is the second hottest portion of the Sun's atmosphere. The temperature of the thermosphere varies significantly depending on the amount of solar activity. The temperature in this altitude region varies from 500 K during periods of low solar activity to 2000 K during periods of high solar activity.

Solar radiation with wavelengths from 0.1 m to 0.2 m is strongly absorbed by molecular oxygen at altitudes of 50-110 km, which is in the mesosphere and lower thermosphere. In the case of O2 , photodissociation is the absorption process in operation. The atomic oxygen produced in this process is a primary constituent of the upper atmosphere. Between 0.2 and 0.31 m , the photodissociation of ozone in the stratosphere is the important process. These reactions maintain the ozone layer, and are discussed in more detail in the next section, and in the Ozone Photochemistry Chapter.

Of the radiation reaching the top of the atmosphere, more than 98% is at wavelengths longer than 0.31 m which includes the near ultraviolet (0.32-0.40 m , known as UV-A), visible, and infrared regions of the spectrum. At these wavelengths, there is very little absorption in the atmosphere. Ozone has a weak absorption band in the visible region due to vibrational- electronic transitions of the molecule. Molecular oxygen has weak absorption at red light wavelengths, CO2 has a number of weak absorption bands in the near infrared, and ozone also displays weak absorption in the infrared. However, these IR bands are more important to the radiative transfer of longwave radiation. The most significant absorber of longer-wavelength solar radiation is tropospheric water vapor. Tropospheric water vapor concentrations are highly variable, and the amount of light that penetrates to the surface is highly dependent on the local amount of water vapor in the troposphere. It also depends on the amount of cloud and aerosol present in the troposphere, which can scatter the radiation back towards space.

Basic Ozone Photochemistry

Stratospheric ozone is created and destroyed primarily by ultraviolet radiation. The chemical reactions that control the gas-phase photochemical production and destruction of ozone are summarized in Figure 17. When high energy ultraviolet photons (about 0.1 - 0.2 m ) strike molecular oxygen ( O2 ) they split the molecule into two single oxygen atoms by photodissociation ( Eqn . A). The free oxygen atoms can then combine with molecular oxygen in a three-body collision (M being the third body) to form ozone ( O3 ) molecules ( Eqn . B). Note that in this reaction M can be any atom or molecule. Its presence is necessary to conserve energy and angular momentum.

Figure 17. Ozone Photochemical Reactions

Ozone is a very important absorber of UV-B radiation (0.29-0.32 m ). When an ozone molecule is exposed to a UV photon with a wavelength in the 0.2-0.31 m range, it may photodissociate back into O2 and O ( Eqn . C). During dissociation, the atomic and molecular oxygen gain kinetic energy which produces heat and causes an increase in atmospheric temperature. The free oxygen atom may then combine with an oxygen molecule, creating another ozone molecule, or it may collide with an existing ozone molecule and break it apart, creating two oxygen molecules. Whenever O and O2 recombine, there is no change in the amount of either constituent, but only an increase in heat due to the absorption of a photon. This is why ozone is capable of absorbing most of the incident UV radiation, despite the low concentrations of ozone in the stratosphere (average of 3 ozone molecules per 1 million molecules air). While most of the UV photons are absorbed high in the stratosphere by some form of oxygen (O, O2 , or O3 ), longer wavelength photons can penetrate deeper into the atmosphere because they are not energetic enough to react with O or O2 and are only weakly absorbed by O3 . Although weak, ozone does absorb radiation at visible wavelengths and this absorption can lead to ozone photodissociation , creating regions of ozone production and destruction lower in the stratosphere. The processes of ozone creation and destruction through interactions with UV and visible photons are known as Chapman Reactions. [ sentence on why ozone layer is at particular altitude]

Most ozone is destroyed through catalytic chemical processes rather than Chapman Reactions. Ozone is a highly unstable molecule that readily gives up its extra oxygen atom to free species such as nitrogen, hydrogen, bromine, and chlorine ( Eqns . E, F on Figure 17). The net reaction converts an ozone molecule to two oxygen molecules. Note that the catalyst in the reaction (denoted by X) is recycled and can go on to destroy more ozone. These catalytic compounds occur naturally in the stratosphere, released from sources such as soil, water vapor, biological processes, and the oceans. Several of these species also have man-made contributions from fertilizer, pesticides, fire extinguishers, refrigerants, Styrofoam, and solvents. These have consequences for the rate of ozone destruction in the stratosphere. Less stratospheric ozone may lead to increases in the amount of UV-B radiation reaching the Earth's surface. For more details on current results concerning ozone decreases and UV-B increases see the information in Ozone Variability and Trends [CLICK TO CHAPTER 8]. If you would like more details on reactions involving ozone in the stratosphere, see Ozone Photochemistry [CLICK TO CHAPTER 5].

Atmospheric Scattering - Some Basics

When light encounters some material, whether as a gas in the atmosphere, a liquid or solid, the light can interact with molecules of that material in such a way that the light's direction of travel is changed but its energy remains the same (or is modified very slightly). We call such an interaction scattering. Unless we are looking directly at a light source, almost all of the light that enters our eyes has been scattered by the atmosphere or solid objects around us.

While absorption is best explained in terms of the particle nature of light (photons), scattering is a wave phenomena of light. In the last section we saw how a molecule can only absorb photons of certain energies, depending on the electronic and vibrational levels of the molecule. However, molecules can interact with electromagnetic radiation of any energy in the scattering process. Even when absorption does not occur, electrons are still somewhat free to move around in a molecule and, in particular, to respond to electric oscillating field of the light. To illustrate the wave nature of light and its scattering interaction, think of a seagull sitting on the water when a boat passes nearby. The wake of the boat (a group of waves in the water) moves toward the bird and eventually passes under it. As it does so, the bird moves up and down following the surface of the wave under it. Now think of a molecule with an electromagnetic wave (light) approaching it from some direction. Since the wave is an oscillation of an electrical field it will be a force felt by a charged particle in the vicinity of another charged particle. So a negatively charged electron will feel a force applied to it by the passing wave, just as the seagull felt the vertical force applied by the passing boat wake. Of course, electrons in a molecule are also under the influence of their electrostatic attraction for the positively charged nuclei. They are going to be pulled only so far before their attraction for the nucleus takes over.

So, as the sinusoidal oscillation of the electric field pass the molecule, the electrons are forced back and forth in a direction that is perpendicular to the direction the light is traveling. This process entails increasing the kinetic and potential energies of the electrons, which means that some of the energy that the electromagnetic wave carries must be transferred to the oscillating electrons.

Now here's the tricky part. Newton's First Law, F = ma, tells us that since a force F is being applied to the electron which has a small mass m, the electron accelerates. It is also a fact that electromagnetic radiation is generated by accelerating charged particles. So, an oscillating force (from the incident light) is applied to the electron, which therefore accelerates, emitting electromagnetic radiation. Since the electron is oscillating back and forth along a line, the electromagnetic radiation it emits is a regular, sinusoidal oscillation at a frequency the electron is oscillating with, which is just the same frequency that the incident radiation had!

Let's summarize the process. An incident electromagnetic wave interacts with a molecule forcing the electrons to oscillate back and forth and removing some energy from the incident wave in the process. Just by oscillating back and forth, the electron emit electromagnetic radiation of the same frequency (and energy) it interacted with in all directions. If the incident wave is only present for a finite amount of time, a certain amount of energy will be lost from the field but by the time the electron is finished oscillating, it will have emitted the same amount of energy as had been lost. The energy of the molecule itself will not have changed, only the direction of the radiation is altered which gives rise to the term scattering. Scattering is called elastic scattering if the energy of the re-emitted radiation is exactly identical to the incident radiation. If the emitted radiation is slightly greater or less, we refer to it as inelastic scattering.

The radiation may have come in one direction but the re-emitted radiation leaves in all directions away from the scattering particle. For a more detailed discussion of scatter direction CLICK HERE [material in brackets below.]

[Radiation from a scattering particle leaves in all directions but it is greatest in the plane perpendicular to the line along which the particle is accelerating. It falls off as the square of the cosine of the angle away from this plane. We can illustrate this in Figure 1?. The coordinate system is selected with the positive z axis in the direction of the propagation of the incident wave, the origin at the center of the molecule, and the x axis lined up with the direction of the electric field vector of the incident wave. In general, the light that is scattered by a particle is scattered more in some directions than others. The phase function P( , ) tells us what fraction of light is scattered in different directions. Here, and are angles that express the direction of scattered light relative to the incident light, as seen in the figure. The angle that the out-going light's direction makes with the z-axis is . If the light leaves going in the same direction then =0 and if it is scattered back toward the same direction as the incident light came from, = . The angle expresses the azimuth of the outgoing light measured from some direction. If the particle is spherical, there will be no -dependence in the phase function. If the particle is not spherical, can be measured from some special geometrical axis of the scattering particle].

As a general rule, the larger the scattering particle is relative to the wavelength of the incident radiation, the more energy is scattered in a forward direction. Although light is scattered in all directions, it is useful to characterize what fraction is scattered in more or less forward direction compared to what is scattered in a more or less backward direction. This is because????????? [Need a statement here about why direction is important]. One measure of the directional scattering is the forward asymmetry parameter (g). A particle that scatters equally as much energy in the forward and backward directions has a value g=0. A particle that scatters light only in the forward direction has g=1 while a particle that scatters only in backward direction has g=- 1. Among the constituents of our atmosphere, g typically ranges from 0 for molecules to about 0.85 for cloud water droplets. The value of g is generally larger for larger particles.

Because of the physics of light scattering by molecules, nearly nine times as much blue light is dispersed by Rayleigh scattering as red light. This accounts for two everyday occurrences. First, if you look straight up at a clear sky it looks blue. The reason the sky is blue is that the light from the Sun travels through the atmosphere, and some of it passes the atmosphere directly overhead. The molecules up there scatter far more blue light than red, so you see the blue. Second, sunsets are red because at the time the Sun sets, light coming from the Sun travels through a more atmosphere to reach the observer, and as a result has much of the blue light scattered out of the direct beam from the Sun to you. What is left is perceived as reddish-orange by your eye.

Larger atmospheric particles, such as aerosols, are always condensed (liquid or solid) aggregates of atoms and molecules. When an electromagnetic wave strikes an aerosol particle, electrons behave as we have described before. However, in these condensed states, the distances between molecules, and therefore the oscillating electrons, are small compared to the wavelength of light. This gives rise to interference (some constructive, some destructive) between out-going electromagnetic waves. This greatly complicates scattering calculations for these particles. Gustav Mie performed these calculations for special cases and so scattering of light by particles larger than the wavelength of incoming radiation is called Mie scattering. Figure 18 summarizes the type of scattering that can occur depending on the particle size and incoming radiation.

Figure 18. Scattering as a Function of Particle Size and Wavelength

Single and Multiple Scattering

NEED MATERIAL HERE

Key Points about Atmospheric Scattering

1.	During scattering, particles absorb light from one direction and re-emit the radiation in all directions.

2.	The amount and angle of scattering depends on the size of the scattering particle and the wavelength of radiation that interacts with the particle.

3.	Generally, no energy is lost or gained by a particle in the scattering process.