Chapter 6
Greenhouses such as the one in Figure 6.1 are covered with glass and the temperature inside is higher than outside. W hy? The short explanation is that the glass is transparent to
Figure 6.1. Greenhouse in the Brooklyn Botanical Garden
79
the incoming visible light from the sun but partially opaque to the (invisible) infrared radiation that the ground emits back to the atmosphere. This energy imbalance causes the inside to warm up. The wavelength and intensity of the outgoing radiation adjusts to reestablish the balance between the incoming and the outgoing radiation at a higher temperature.
W hy does the ground emit radiation at all? We discussed in the previous chapter that if net energy is absorbed by an object, and nothing else (such as evaporation, melting, etc.) takes place, then its temperature will rise. For the temperature to be stable, a balance in energy input and output must be maintained. The way for any object to establish such a balance is to reradiate the radiation. W hat kind of radiation? To try to answer this question I will use another example— we take an iron poker and insert it into a furnace. Initially we will see no change, but if we touch the end of the poker, we get burned and will not try to touch it again. If we keep the poker in the furnace, and if the furnace is powerful enough, the end of the poker will turn red. We describe this as “red hot,” and the object emits red light. If we continue to keep the poker in the furnace, the color will gradually change to blue— even hotter than “red hot.” Th is example demonstrates that we can generate light by heating an object, and the kind of light emitt ed will depend on the temperature of the object. Can we arrive at a quantitative description that correlates the temperature of the object with the emitt ed light?
The idealized model that enables us to calculate the relationship between the emitt ed radiation and the temperature assumes that the hot object is an ideal blackbody, which is an object that is an ideal absorber of radiation, and therefore there is no reflection at any wavelength (color). As we discussed in Chapter 5, in order to maintain a steady temperature, one requires that if the object is an ideal absorber it also has to be an ideal emitter of radiation. Th e radiation emitted by such an object is called blackbody radiation. The model assumes that the radiation from such an object depends only on the temperature of the object— not on the nature of the incoming radiation and not on the material that the blackbody is made of. The closest one can get in constructing such an object is to construct a cavity, paint it black on the inside, and make a very small hole in the cavity for the incoming and outgoing radiation. Once we assume such an object, it is fairly easy to calculate the spectrum of the radiation that comes out of such a cavity as a function of the temperature of the cavity. The calculation is not much diff erent from calculating the possible tones that emerge out of a musical instrument (however, the radiation is diff erent). The spectrum of the radiation may be seen in the plots of the intensity of the radiation as a function of wavelength. Two such plots, for two diff erent temperatures, are shown in Figures 6.2 and 6.3.
Figure 6.3. Spectral distribution of blackbody radiation at a temperature of 290 K
There are a few things to notice in these fi gures: First, the horizontal axis gives the wavelength of the emitted radiation. I present the wavelengths not in meters, as I did previously, but in units of nanometers (nm; billionths of a meter). We do this by multiplying the wavelength by a billion, thus avoiding carrying the negative exponents that will always be present as long as the wavelength is considerably smaller than 1 m. All the wavelengths of interest here will be considerably smaller than 1 m. Second, the scales are diff erent. The scale on the horizontal axis in Figure 6.2 is between 0 and 4000 nm. The scale on the horizontal axis in Figure 6.3 is between 0 and 60,000 nm. The maximum of the peak in Figure 6.2 is around 500 nm, whereas the maximum in Figure 6.3 is around 10,000 nm. It is clear that the spectrum in Figure 6.3, which corresponds to a much lower temperature, extends toward much longer wavelengths as compared with the spectrum in Figure 6.2. The scales of the vertical axes in both figures are normalized to the height of the peak in Figure 6.2. This means that we arbitrarily assign the value of 1 to that peak and the rest of the amplitudes are expressed as a fraction of this value. We can see that the amplitude at the maximum of the peak in Figure 6.3 is considerably smaller than the corresponding amplitude in Figure 6.2— in fact more than 5000 times smaller. Aside from these quantitative differences, the shapes of the two figures are remarkably similar. This shape is the fingerprint that indicates that the objects behave very much like the ideal blackbody for which we can make the calculations.
Let us now summarize the main conclusions from our definition and calculations of blackbodies:
One can express the last two conclusions in terms of two important formulas that will allow us to make quantitative connections between light sources such as the sun and the temperaThe T in both equations refers to the absolute temperature (in Kelvin) and the A in the Stefan- Boltzmann equation refers to the area of the surface of the blackbody in units of m2 .
| ture of objects with which their light interacts: | |
|---|---|
| Wien’s displacement law: position of the maximum of | |
| the peak of the spectrum (in nm) = 2.9 ×106 T . | [6.1] |
| Stefan-Boltzmann law: energy flux (in Watts) = 5.67 × 10- 8 × A × T4 . | [6.2] |
Let us now try to predict some consequences from these relationships. Th e temperatures in Figures 6.2 and 6.3 were not chosen arbitrarily. One (Fig. 6.2) represents the surface of the sun and the other one (Fig. 6.3) represents the surface of Earth. W hen we check the spectral distribution of the radiation from the sun and Earth, they roughly resemble the distribution that we expect from radiation from an ideal blackbody. The maximum intensity of the solar radiation corresponds to the color green, which approximately corresponds to wavelength of 500 nm. We see the sun as white-yellow because of the mixture with other wavelengths. If we insert the value of 500 nm in Wien’s formula, then we get a temperature of 5800 K . Figure
6.2 corresponds to the equivalent blackbody radiation of the surface of the sun. Th e average temperature on Earth is around 15°C, which corresponds to 288 K (Fig. 6.3). We see that the much shorter wavelengths of the solar radiation correspond to a much higher temperature than Earth’s radiation. In the previous chapter I discussed the concept of entropy, its relation to disorder, and the second law of thermodynamics that tells us that if left on its own, a system tends to maximize its entropy. I defined changes in entropy as the ratio between the amount of heat added or removed from a system and its absolute temperature. Here we discover that short-wavelength light has a much higher temperature than longer-wavelength light. So for a given amount of energy, the light with a short wavelength will have much lower entropy than light with a longer wavelength (equation 5.7). In the last chapter we saw that a given amount of energy with a lower entropy content (high temperature) can be converted more effi ciently to work than the same amount of energy with higher entropy content (equation 5.5). Th at is why we refer to the short-wavelength solar radiation as “high-quality ” radiation and the long-wavelength thermal radiation emitted by Earth as “low-quality ” radiation. In the energy interplay between Earth and the sun, there should be a balance between the short-wavelength solar energy absorbed by the Earth and the long- wavelength thermal energy emitted by Earth into outer space. Otherwise Earth’s temperature will change over time. We see that the incoming energy is of a higher quality than the outgoing energy. The physical property that describes this quality is the entropy. One can translate these statements to say that the only net input from outer space is the “order” with which everything on Earth (including us) is being energetically maintained.
Box 6.1 uses the Stefan-Boltzmann relationship, together with very simple geometric considerations, to calculate the energy balance between Earth and the sun, from which we calculate the average temperature on Earth. We use the Stefan-Boltzmann equation to derive the energy fl ux from the sun and then estimate the fraction of this energy intercepted by Earth. For Earth’s temperature to be in equilibrium, this energy should be equal to the energy flux that Earth emits into space.
Box 6.1
Figure 6.4. Schematic diagram of Earth’s orbit around the sun
This box will demonstrate a simple use (although with big numbers) of the Stefan- Boltzmann law with simple geometry to estimate the global average temperature. This provides us with the baseline on which all the Earth– sun interplay that determines Earth’s climate is based.
Let us first use the Stefan- Boltzmann equation to calculate the energy flux from the sun.We have already calculated that the temperature equivalent of the green light from the sun is 5800 K.The radius of the sun is 700 million meters (7 × 108 m). So the surface area of the sun is 4πR2 = 4 × π × (7 × 108)2 = 6.1 × 1018 m2. Thus
solar energy flux = 5.7 × 10– 8 × A × T4 = 5.7 × 10– 8 × 6.1 × 1018 × (5800)4 =
3.9 × 1026 W.
This energy is spreading in space in all directions. The average distance between Earth and the sun is often known in astronomical calculations as 1 astronomical unit (1 AU) and is equal to 1.5 × 1011 m. At this distance, the light spans the surface of a big sphere with this radius. Most of this light goes unobstructed into outer space. A small fraction is intercepted by Earth.This fraction is equal to the ratio of the cross section of Earth to the area of this big sphere. Let us calculate this ratio:
area of the sphere of 1 AU radius = 4 × π × (1.5 × 1011)2 = 2.8 × 1023 m2 ,
area of Earth’s cross section = π × (radius of Earth)2 = π × (6.4 × 106)2=1.3 × 1014 m2 .
The ratio between these two numbers times the solar energy flux will give us the solar energy flux intercepted by Earth:
energy flux intercepted by the Earth: = 3.9 × 1026 × 1.3 × 1014/2.8 × 1023 = 1.8 × 1017 W.
Not all of this energy is absorbed. About 30% of this energy (given the name albedo) is reflected back into outer space by the ground, oceans, and clouds— a process that will be discussed more fully in the next chapter. So the net absorbed energy flux is 1.8 × 1017 × (1 – 0.3) = 1.3 × 1017 W.
In order to maintain a constant temperature, all the absorbed energy is radiated back into space at a wavelength characteristic of Earth’s temperature. Applying again the Stefan- Boltzmann equation,
Earth energy flux = 1.3 × 1017= 5.7 × 10- 8 × 4 × π × (Earth radius)2 × T4 = 5.7 × 10- 8 × 4 × π × (6.4 × 106)2 × T4
will give us T = 258 K = – 15°C. This is cold. The recorded global average temperature is 15°C, which in absolute degrees is 288 K. The “error” in our very simple calculation is on the order of 10%.
“Outside the box” we will see that there are good physical reasons why our calculated temperature is on the low side.
Again I use the Stefan-Boltzmann equation— this time to calculate the blackbody temperature that emits the energy. The value that we arrive at is an average temperature of – 15°C. This is cold. The measured average temperature is 15°C. It looks like a big difference— but if we translate these temperatures to the absolute scale with which we did the calculations, the temperature that we calculated is 258 K while the measured average temperature is 288 K— a diff erence of about 10%, small but very important because it swings around the freezing point of water. At
–15°C almost all water would be frozen and life would have certainly been different from what we experience. W hat did we forget in our calculations? We “forgot” that we have an atmosphere.
Let us go back to the greenhouse shown in Figure 6.1. Without the glass, the temperature inside the structure would be the same as the temperature outside because the interior of the greenhouse is part of the outside. We know that the glass is responsible for the hott er temperature because we have relevant experiences within different contexts that we actually can control to some extent. We leave our car locked outside on a sunny day—we return and the car is steaming. If we want to equilibrate the temperature with the outside, we open the window. If we want to minimize the effect, we leave the windows of the car shaded with a cardboard sun shade. W hat the glass is doing is letting the visible light from the sun get in but blocking the thermal radiation that the interior is trying to radiate back. Glass can do that because it is almost completely transparent to the incoming short-wavelength radiation from the sun but is partially opaque at the wavelengths that the interior is trying to radiate back. In the case of the greenhouse, the sun shines and the light just passes through and gets absorbed in the interior by the plants and the ground.
To equilibrate the temperature, the interior radiates back the thermal radiation characteristic of its temperature. But the glass is not transparent to this radiation. The radiation in a sense will be trapped inside. This will create an imbalance between the incoming and the outgoing radiation (which can be quantified and is known as radiative forcing). As a result the temperature of the interior will rise. This temperature rise will result in some shift in the outgoing radiation toward shorter wavelengths (Wien’s displacement law) and an increase in the energy flux of the outgoing radiation (Stefan-Boltzmann law) until a new equilibrium between the incoming and the outgoing radiation is established that leaves the interior of the greenhouse at a higher temperature than the outside.
Let us now go back to our global temperature calculations in Box 6.1— what does the glass have to do with this calculation? In a sense, the atmosphere acts as our global “glass.” Th e approximate composition of the atmosphere is given in Table 6.1.
The two main elements of the atmosphere, nitrogen and oxygen, are transparent to both the visible and infrared parts of the spectrum and have only indirect roles in the energy balance of Earth. They are active in transporting heat by convection (movement of air). Argon, neon,
Approximate composition of the atmosphere
| Gas | Concentration (percent volume) |
| Nitrogen (N2) | 78.1 |
| Oxygen (O2) | 20.9 |
| Water (H2O) | 0.05– 2 (variable) |
| Argon (Ar) | 0.9 |
| Carbon dioxide (CO2) | 0.036 |
| Neon (Ne) | 0.0018 |
| Helium (He) | 0.0005 |
| Methane (CH4) | 0.0002 |
| Krypton (Kr) | 0.0001 |
| Hydrogen (H2) | 0.00005 |
| Nitrous oxide (N2O) | 0.00005 |
| Ozone | Variable |
helium, and krypton are inert gases that also do not have direct roles in the energy balance. Methane, nitrous oxide, and ozone are “greenhouse gases” (“GHGs”), and their role will be discussed separately. Their concentration in the atmosphere is relatively small. The two constituents that play the most important roles are water and carbon dioxide.
The full energy budget of Earth is presented in Figure 6.5. The outer atmosphere intercepts
341.3 W/m2 solar radiation. For the temperature to remain constant on the average, we need the same amount of energy to be emitted into outer space. About 30% of the incoming radiation is refl ected and scattered back into outer space. This percentage is known as the albedo. As we will see shortly, it plays a crucial role in the balance. The remaining 70% is returned to outer space as infrared radiation. Figure 6.5 shows that not only the overall energy flow but also the regional energy flow is balanced in order to have approximately constant temperature on the surface of Earth.
Figure 6.5. Global energy balance (in W/m2). Solar radiation is on the left and thermal radiation is on the right. Source: Adapted from Kiehl and Trenberth (1997).1
Greenhouse Gases
Atmospheric water is the most important element responsible for the rise of the average temperature on Earth to values above freezing (0°C) and has nothing to do with the human presence on Earth. The intimate connection between the energy balance and the water cycle will be discussed in the next chapter. Following our discussion in Chapter 2, humans are probably responsible for about 30% of the carbon dioxide in the atmosphere, and as long as the main energy source for human development is fossil fuels, this fraction will continue to rise. Th is issue is the common thread throughout the book. Before we continue with our focus on our dependence on fossil fuels, we need to examine the effects of the other gases in the atmosphere on global warming. These gases include methane, nitrogen oxides, ozone, chlorofluorocarbons, and various particulate, atmospheric components known collectively as aerosols. Before we proceed we need a common scale that will quantif y the contributions of the diff erent atmospheric constituents to global warming.
Radiative forcing is defined as the change in the net radiation at the top of the troposphere that occurs because of a change in concentration of an atmospheric component or some other change in the overall climatic system, such as solar insolation. Sudden doubling of the pre– Industrial Revolution atmospheric concentration of carbon dioxide while holding everything else unchanged will result in a radiation imbalance at the top of the troposphere to the extent of 4 W/m2. More solar radiation will reach the top of the troposphere than the infrared radiation that will be radiated back by Earth because of the trapping of the outgoing radiation by carbon dioxide. This imbalance will, in time, readjust itself to restore a new equilibrium at a higher average global temperature. Positive forcing, on average, will cause readjustment by warming Earth and negative forcing results in cooling Earth. A summary of the radiative forcing of various atmospheric components is given in Table 6.2.
Concentrations in Table 6.2 are given in units of parts per billion by volume (ppbv). If we take the pre–Industrial Revolution concentration of methane to be 700 ppbv, then we mean that there are 700 methane molecules per billion gas molecules.
Table 6.2 shows that the concentrations of non– carbon dioxide GHGs are small compared to carbon dioxide. However, their relative radiative forcing is much larger than their relative concentration. This is because they are much better absorbers of thermal radiation than carbon dioxide is. The result is that the total radiative forcing of methane, nitrous oxide, and the most prominent chlorofluorocarbons amount to over 50% of that of carbon dioxide. The contributions of these gases, and a few not included in Table 6.2 that will be discussed shortly, to global warming are so important that one of the most respected experts in this field,
Radiative forcing of some of the atmospheric components
| GHG |
|---|
| Carbon dioxide |
| Methane |
| Nitrous oxide |
| CFC- 12 |
| Chemical | Pre– Industrial |
| formula | Revolution |
| concentration | |
| (ppbv) |
Atmospheric concentration in 1994 (ppbv) R adiative forcing due to the concentration increase (W/m2) Approximate radiative forcing per molecule relative to CO2
CO2
278,000
358,000
1.46
1
CH4
700
1721
0.48
20
N2O
275
311
0.15
200
CCl2F2
0
0.5
0.17
18,000
Dr. James Hansen, head of the NASA Goddard Institute for Space Studies, in his testimony to the Senate Commerce Committee, suggested that in the short term it might be benefi cial to focus policy attention on reducing their concentrations at the expense of less emphasis on carbon dioxide.2
An alternative measure of contribution to global warming is the global warming potential (GWP). It measures the contribution to global warming of a given mass of gas relative to the same mass of carbon dioxide over a given period of time. For short- lived materials such as methane, the GWP will be considerably greater for a short interval than a long interval.
I will now proceed to discuss the origin, including the anthropogenic origin, of these gases. I will include in this discussion gases such as ozone and aerosols that are not included in table 6.2.
Methane
Methane is the main component of natural gas. About 70% of the current methane emissions to the atmosphere are anthropogenic in origin. Table 6.2 shows that the atmospheric increase of methane since the Industrial Revolution far outpaces (as a percentage of the pre–Industrial Revolution level) that of carbon dioxide. Methane is produced when biological materials decompose in environments deficient in oxygen. Such conditions prevail when the decomposition takes place under waterlogged conditions such as in swamps. That is the reason that methane is sometimes referred to as “swamp gas.” Ruminant animals, such as cattle and sheep, that can digest cellulose produce large amounts of methane as a “ by- product” of their digestive process. Decomposition of organic garbage in landfills that takes place under oxygen-starved conditions also produces large amounts of methane.
The presence of large quantities (probably larger than the total known land fossil- fuel reserves) of “methane hydrates” trapped at the bottom of the oceans will be discussed in Chapter 10. The lifetime of methane in the atmosphere is a relatively short 12 years. Methane is destroyed in the atmosphere through oxidation to carbon dioxide by very active intermediate products of the solar-induced decomposition of water and oxygen molecules.
Nitrous Oxide
The common name for nitrous oxide is “laughing gas.” Table 6.2 shows that per molecule, this gas is about 200 times more effective in blocking outgoing thermal radiation than carbon dioxide. About 40% of the current nitrous oxide emissions results from anthropogenic sources. Most of the natural supply of nitrous oxide originates through oceanic release. The gas is a by- product of biological reactions that add or subtract nitrogen from biological molecules. The anthropogenic contributions are due to the increased use of fertilizers and the industrial synthesis of products such as nylon. The lifetime of nitrous oxide in the atmosphere is about 100 years. It slowly rises to the stratosphere where it absorbs energetic solar radiation that decomposes it to nitrogen gas and atomic oxygen.
Air Suspensions
Routine monitoring of air quality, such as that conducted by the US Environmental Protection Agency (EPA), includes tests for chemicals such as carbon monoxide, nitrogen dioxide, ozone, sulfur dioxide, and so forth. These chemicals originate from some form of human activity, and all can cause a variety of environmental damage. They are well-specified chemicals with known properties. An additional, widely monitored pollutant is referred to as particulate matt er, often with a designation such as PM- 10. This pollutant is not a chemical compound but a mixture of solid particles and liquid droplets suspended in the air. Scientists oft en refer to these particles as aerosols. Some of these particles are large or dark enough to be seen as soot or smoke, whereas others are so small they can be detected only with powerful electron microscopes. These particles have the ability to scatter light and to reduce visibility. Th ese abilities make them participants in Earth’s energy balance, but their importance as significant participants in global warming was only recently recognized. These particles come in a wide variety of sizes. The size designation appears as the number associated with the letters PM: PM- 10 are course particles with size less than 10 μm (one thousandth of a centimeter), whereas PM- 2.5 are fi ne particles with sizes less than 2.5 μm. These particles originate from many anthropogenic and natural sources. They can be emitted from fuel combustion in motor vehicles, power generation stations, industrial emissions, and smokestacks. They are the products of atmospheric reactions of other air pollutants such as sulfur dioxide (SO2), nitrogen oxides, and volatile organic carbon. They form in powerful volcanic eruptions, from which they can reach the upper atmosphere or stay in the lower troposphere. The ability of these suspensions to scatter light means that they can reflect incoming solar radiation and absorb or reflect outgoing thermal radiation. In the next chapter I will discuss the role that clouds play in the energy balance, as clouds are local aggregates of particulate water droplets.
A vivid demonstration of the contributions of these suspensions to the global climate came with the eruption of Mount Pinatubo in the Philippines in 1991. A massive injection of volcanic ash into the stratosphere took place due to the eruption. This injection of largely black ash, composed of relatively large particles, immediately resulted in the absorption of incoming sunlight and outgoing thermal radiation. The result was a significant heating of the lower atmosphere. This local heating lasted only few months. Within a few months, the heavy ash particles had been subjected to the gravitational force, and they precipitated by falling back to Earth. W hat remained in the atmosphere from the eruption was about 30 million tons of sulfur dioxide. In time, the sulfur dioxide got oxidized and recombined with water to form droplets of sulfuric acid (which will be discussed in the next section on Venus). Th ese droplets scattered the incoming solar radiation to a degree that the average global temperature over the next 2 years dropped by about 0.2°C.
It is now recognized that the net effect of the air suspensions is negative radiation forcing in the radiation balance. On average they cause a reduction in the global temperature. In scientific publication and in his testimony before the US Senate, James Hansen has estimated that the sum of the forcing of sulfate aerosols and suspensions caused by organic carbon, biomass burning, and soil dust contribute about – 1.2 W/m2. Volcanic aerosols contribute an additional – 0.2W/m2. The sum of these contributions is approximately equal to the contribution of carbon dioxide. If one adds indirect effects such as cloud changes, then the negative radiation forcing can be even larger. Because of the diversity of mechanisms, the uncertainty in these estimates is very large. One might be tempted to look at these numbers and suggest countering the carbon dioxide contribution with a deliberate increase in the release of material that contributes to air suspensions. As Hansen pointed out, this will be a Faustian bargain because the potential improvement in the energy balance will be more than compensated for by environmental deterioration in other areas, such as acid rain and the adverse health eff ects of breathing these particles.
Ozone
One chemical in Table 6.2 was not yet mentioned. It is designated there as CFC-12. It is a member of the chlorofluorocarbon family, which consists of compounds that have carbon chains in which the hydrogen atoms are replaced with chlorine and fluorine atoms. CFC-12 is one of the simplest and has a the chemical formula CCl2F2. This family of compounds became very well-known because they are major contributors to the recent anthropogenic creation of a hole in the stratospheric ozone layer. The issues of global warming and the destruction of the stratospheric ozone layer are often lumped together. These are two separate issues, but as with most other environmental issues, there are connections.
W hat is ozone? A chemical reaction between molecular oxygen and atomic oxygen that can be writt en as
O2 + O → O3 [6.3]
produces a new molecule— ozone. If we are unlucky enough to be in an urban area that suffers from severe smog, then we can smell it. We can also smell it if we want to look good and get an artificial tan in a tanning bed, which contains mercury lamps. Ozone is a very reactive molecule. At low levels of the troposphere, it is an indirect environmental pollutant, excessive amounts of which can invoke government warnings to limit our activities and cause severe health damage. On the other hand, stratospheric ozone acts as a shield against the ultraviolet parts of the solar radiation, which would cause irreversible chemical changes in the principal organic molecules that constitute the backbone of all living organisms, such as DNA, RNA, and proteins, making them unsuitable for life support.
Ozone in the Stratosphere
Box 2.6 in Chapter 2 shows a graph that describes changes in the atmospheric temperature with height. As we rise in the atmosphere, the temperature decreases until we reach heights of around 12 km, where the temperature starts to rise again. The height where this temperature change takes place constitutes the boundary between the troposphere and the stratosphere. The main reason the temperature starts to rise at this point is that the stratospheric ozone absorbs the ultraviolet part of the solar radiation. This absorption is the important shield discussed in the previous section.
The ozone in the stratosphere is formed by the reaction of molecular oxygen with the most energetic part of the solar radiation. This radiation is energetic enough to break the bond in the oxygen molecule and create oxygen atoms that can react with other oxygen molecules to create ozone according to the previous reaction. This simple molecule provides the shield for Earth- bound biological systems for the billion years or so that oxygen was present in the atmosphere. Suddenly, very recently, the ever- innovative human mind came up with a wonderful product: chlorofluorocarbons, commonly known as Freon. These were mainly used as refrigerant fluids. Other uses included solvents, aerosol propellants, and so forth. They were ideal: very stable, not corrosive, and nontoxic. But in these applications, we were releasing them to the environment, and because of their stability, they stayed in the troposphere for a very long time and eventually entered the stratosphere. Once they entered the stratosphere, the intense ultraviolet radiation is energetic enough to break the carbon– chlorine bond according to the following reaction:
CF2Cl2 → Cl· + CClF2 · . [6.4]
(The dots in these compounds indicate that these are “free radicals”—very active chemicals.) The chlorine atom is very active and reacts with ozone according to the following reaction:
Cl· + O3 → O2 + ClO· . [6.5]
ClO· is a very active intermediate that, with the help of oxygen atoms, can decompose to form chlorine atoms that can find other ozone molecules to decompose. So the chlorofluorocarbons here act as catalysts in a sense because very few of them are needed to decompose large quantities of ozone. From Table 6.2 we can see that the chlorofluorocarbons are powerful GHGs with a very large radiative forcing per molecule. This positive radiative forcing will be partially compensated for by the reduction in stratospheric ozone.
Ozone in the Troposphere
How do we get ozone into the troposphere? In order to get ozone, we need oxygen atoms. We do not get enough energetic solar radiation in the troposphere to break the oxygen–oxygen bond of molecular oxygen. The secret to the formation of ozone in the troposphere is in the activity of nitrogen. W hat happens with nitrogen? In the case of the internal combustion engine, fossil fuel is burned by oxidizing the fuel with oxygen from air. Close to 80% of air is molecular nitrogen. Under ordinary conditions nitrogen is a very inert gas. However, at the high temperatures inside the engine, nitrogen can interact with oxygen according to the following reaction:
N2 + O2 → 2NO. [6.6]
This reaction produces nitric oxide, which reacts relatively easily with oxygen to produce nitrogen dioxide—NO2. Together NO and NO2 are often called NO x . NO2 is a pungent, red- brown gas that can effectively absorb solar radiation. W hen the gas absorbs solar radiation, it will decompose back to nitrogen oxide and atomic oxygen according to the following reaction:
NO2 + solar radiation → NO + O. [6.7]
Now we have again our atomic oxygen that can react with molecular oxygen to produce ozone. Ozone is a direct health hazard, is a GHG, and also reacts chemically with many other chemicals in the atmosphere to have a significant effect on almost every environmental issue the atmosphere is involved in.
I shall end this chapter by discussing the situation on Venus, a neighboring planet. Table 6.3 shows a comparison between Earth and Venus.
The properties of the two planets are very similar—all the way down to the last entry. Th e calculation of the temperature in the absence of atmosphere follows exactly the same model we used in Box 6.1— these calculated temperatures are also very similar. However, the last row
Table 6.3.
Some physical characteristics of Earth and Venus
| Property | Earth | Venus |
|---|---|---|
| Radius (m) | 6.4 × 106 | 6.3 × 106 |
| Mass (relative to Earth) | 1 | 0.814 |
| Density (g/cm3) | 5.52 | 5.24 |
| Average distance to sun (m) | 1.5 × 1011 | 1.1 × 1011 |
| Albedo | 0.39 | 0.59 |
| Temperature calculated in the absence of atmosphere (°C) | – 15 | – 11 |
| Average surface temperature (°C) | 15 | 464 |
shows a big diff erence. On Earth, the measured average temperature is a comfortable 15°C, whereas on Venus the average surface temperature is 464°C, which is probably unsuitable for any life- form. W hat causes such a diff erence? It turns out the atmosphere of Venus is made of 96.5% carbon dioxide, as compared with 0.036% on Earth. In addition the atmospheric pressure on Venus is about 90 times that of Earth. Venus has a much more eff ective thermal blanket than Earth. In addition, clouds and rain on Venus are made of sulfuric acid. One might be inclined to speculate that Venus was “created” to demonstrate what will happen to a planet run amok. As far as we know, no humans or any other advanced life- form ever drove cars on Venus. The reasons behind the atmospheric differences on the two planets are still subjects of active research focused on the premise that Earth and Venus have similar amounts of carbon dioxide but on Earth most of the carbon dioxide is dissolved in the oceans, whereas on Venus oceans never formed and the carbon dioxide stayed in the atmosphere. The original carbon dioxide on both planets came from volcanic eruptions that took place very early in the life of both planets. On Earth, water was a major part of the eruptions, and the surface temperature was cold enough for water to condense, creating oceans able to dissolve the carbon dioxide. On Venus, either the eruption was dry (no water) or, because of the closer proximity of Venus to the sun, the temperature was high enough (the albedo probably also changed with time) to prevent condensation. The atmospheric water was decomposed by the ultraviolet radiation of the sun. The resulting hydrogen escaped into outer space, and the oxygen probably was used in other chemical reactions. The end result in both scenarios is the same: no water, no oceans, and nothing to dissolve the carbon dioxide. The carbon dioxide remains in the atmosphere to form a dense thermal blanket that absorbs the thermal radiation— and the temperature rises until energy balance is reestablished.
