Climate Change

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This Wikibook deals with climate change and global warming. If you care to share your knowledge of this topic, please contribute.

Although this book is still in a very preliminary state, we seem to be moving toward a cohesive structure. The book is really two parts: first, a primer on the science of climate change, and second, a study of societal implications of climate change. The first part deals with much more concrete, physically based principles. As such, it can be developed more easily. The second part is potentially more controversial, but not necessarily so; for the early stages of development, the second part should probably be limited to generalities and external links.

[edit] Introduction

[edit] What is Climate?

Climate is a broad term, but it always describes a long-term average of a system. Often 'climate' is used to mean the long-term mean state of the atmosphere, including temperature, humidity, and wind. In other contexts, 'climate' can include the oceanic state, the cryosphere (snow and sea-ice), the biosphere, and sometimes even the lithosphere (Earth's crust). Meteorologists and atmospheric scientists often say that climate is what you expect; weather is what you get.

[edit] What is Climate Change?

Although dynamical systems, like Earth's climate system, are full of complicated processes that lead to chaotic variations, changes to external forcing of the system can lead to significant changes. For the Earth's climate, we usually think of trends in global average quantities (especially surface temperature) as indicative of climate change. When the trend leads to a change larger than the natural variability, a statistically significant change most certainly has occurred. DEFINITIONS

[edit] Anthropogenic Climate Change

Anthropogenic means "human caused," form "anthro-", meaning "human," and "genic," meaning "produced by, origin, cause". The term anthropogenic climate change is used to attribute changes in Earth's climate to activities of humans. In recent times, this has been taken as implying mainly the emission of "greenhouse gases" into the atmosphere, usually by burning fossil fuels.

How can humans change Earth's climate? Even as far back as Arrhenius^[1] people have been aware that the composition of the atmosphere affects the climate. Some gases, like carbon dioxide, have molecular structure that allows the absorption of certain wavelengths of light. In the case of "greenhouse gases," that means absorbing infrared radiation. The distinguishing characteristic of a greenhouse gas is that it absorbs infrared radiation better than it does visible radiation; this allows sunlight to penetrate through the gas (the atmosphere) and warm Earth's surface. The Earth then radiates as a blackbody, emitting infrared radiation that is then trapped in the atmosphere. This is the greenhouse effect.

If humans change the composition of the atmosphere, say by burning fossil fuels which release carbon dioxide, then more energy goes into the atmosphere than would have otherwise. More energy leads directly to higher temperature, hence climate change.

[edit] Historical Footnote

Climatology is a young science, with modern climate science only emerging from meteorology, oceanography, and geology in the late 20th Century. Of course, people have been interested in the natural world, including movements of air and water, for a very long time. One early example of a theory for anthropogenic climate change is George P. Marsh's^[2] "The Earth as Modified by Human Action," published in 1874. The science in this early effort is far from the level of climate science today, but Marsh does link land use change, including deforestation and irrigation, to changes in the local climate.

[edit] A Wikibook digression

One contributing author to this wiki book stated that decreased nocturnal cooling may never have been "considered in any debate about global warming." The argument was stated as

All planets with rotational days unequal to their orbital years absorb their sun's heat during their day and release it at night. In the case of planet Earth, however, not only are we adding to the sun's heat in the daytime; The ever-increasing tendency away from regular day-night cycles of work, play and sleep means that at night, the time when our Earth should be shedding its excess heat, we are still adding to it.

This is a fair argument at first blush, but it does not hold up under scrutiny. While Earth cools much more efficiently at night at the surface, the better cooling does not continue into the upper troposphere very well. That means that most of the energy from the cooling will still end up where it would during the day: either absorbed in the troposphere or emitted to space. Also, the argument seems to imply that increased nocturnal activity by humans makes the cooling less efficient, but it is an extremely small effect. The more efficient cooling at night is due almost entirely to the absence of sunlight. Think of the evolution of the surface temperature as dT = S - F to zeroth order, where S is the solar energy absorbed at the surface and F is the cooling by infrared emission. At nighttime, S = 0, so dT is all due to cooling by emission. During the day, the warming offsets the cooling. We are not adding to the sun's heat, as the contributor states, but just trapping it in the troposphere. That trapping has no diurnal cycle, since there is a negligible diurnal cycle in the concentration of atmospheric constituent gases. Let this be a lesson for the reader: critical thinking should always accompany learning about new topics.

[edit] The Science of Climate Change

Climate change has become a "hot button" issue over the past few years, and this has only become more extreme as United States policy on climate change has diverged from much of the rest of the world and the scientific community. However, climate change is first and foremost a scientific topic.

The study of climate -- sometimes called climatology or climate science -- is actually a relatively young field, but has roots in all the major branches of science. It is most easily associated with atmospheric science (and its older name, Meteorology) and oceanography. There are also clear connections with the cryosphere (glaciology) and the biosphere (biology, ecology). Physical and computer models built to predict climate change are based on evidence gathered from glacial geologists and quaternary geologists, that infer, with varying degrees of precision, the Earth's climate history. Many of the fundamental concepts of climate science are straight out of elementary physics. The equations of motion are the same fundamental equations that govern all classical fluid dynamics, much of energy transfer is based on well-known principles of radiative transfer and nuclear physics and spectrometry, and a lot of observations are based on geological, chemical and biological processes and methods. This is all to say that climate science is a multi-disciplinary field, with diverse (even disparate at times) interests and applications. It is unified only by the end goal: to understand the physical processes governing our natural world.

When focused on issues of climate change, the same physical principles are at work. Instead of describing the average climate, or even the natural variability of climate, climate change studies try to quantify differences or trends. Modeling studies might compare simulations with different levels of carbon dioxide, or an observational study might describe a slowly changing quantity, like sea-ice concentration, over a long period of time. Much of the interest in studying climate change is motivated by the idea that human activity has changed and will continue to change the climate. Most research points to rising levels of carbon dioxide in the atmosphere as the primary factor leading to anthropogenic global climate change.

When the composition of the atmosphere changes, for example by changing the carbon dioxide concentration, the radiative properties of the atmosphere might also change. In the absence of an atmosphere, Earth would look a lot like a black body radiator; that is to say, the sun would shine on Earth, which would warm to an equilibrium temperature, and then a balance would be struck. That balance (radiative equilibrium) would have Earth radiating as much energy to space as the sun delivers to the surface. Mostly due to the fact that Earth is so small and intercepts so little of the total energy emitted from the sun, that radiative equilibrium temperature is much lower than the sun's temperature. Using Wien's law, we can calculate that temperature and establish that Earth is an infrared emitter.

When there is an atmosphere, like the one on Earth, some of the gases that make up the atmosphere can absorb infrared radiation. That interaction between photons and molecules increases the temperature of the atmosphere, which then emits at a slightly different wavelength. The emission from the atmosphere goes both out to space and downward, back to the ground where it is absorbed by the surface. This process, whereby energy that is emitted from the surface is absorbed by the atmosphere which then emits energy back toward the surface, is called the greenhouse effect, and it is one of the basic feedback processes in the climate system. It increases the surface temperature on Earth from the radiative equilibrium temperature to a much more life-friendly temperature. Global warming, or anthropogenic global warming, is the difference in the global mean temperature in a world with artificially elevated carbon dioxide compared to a reference state (which is usually taken as a time before the Industrial Revolution). In the rest of this part of the book, we investigate the processes involved with climate and climate change, from the sun's influence, to the natural greenhouse effect, to observed changes in the composition of the atmosphere. We will focus on feedbacks and processes that are thought important in both stabilizing and amplifying changes to the global climate.

[edit] The Sun's Influence on Earth

Nearly all the energy impacting Earth's climate comes from the sun, even if it is sometimes indirectly as we shall see.

The sun works as a thermonuclear engine, emitting energy that is released by fusion of hydrogen atoms in the sun's core. A complete description of the inner workings of the sun, including sun spots and the solar wind, are beyond the scope of this book. It is a topic worth reading about though, since the sun supplies the energy needed for life on Earth.

The photons (electromagnetic energy) emitted from the sun reach Earth's orbit in about 8 minutes. The temperature of the "surface" of the sun (the photosphere) is about 5700 K, and has been determined in many ways, including a simple calculation that is included below. From conservation of energy, we can guess what happens. As the light travels away from the sun, it spreads out homogeneously over an expanding spherical shell, so a lot of energy gets spread over a huge area, so the concentration at any particular point decreases with distance from the sun. By the time the energy reaches Earth's orbit, the energy flux (energy per time per area) is only about 1367 W/m², the so-called solar constant. This single number is the beginning point for most of climatology, especially historically.

Using the solar constant, we can calculate Earth's temperature if there were no atmosphere. To do so, we treat the Earth as a blackbody, which means that it is in radiative equilibrium, all the energy that it absorbs (that is, all the energy incident at its surface) is emitted. The Steffan-Boltzmann law law governs blackbody radiation, and can be stated as

where the left-hand side is the flux, which is proportional to the temperature of the blackbody to the fourth power. The σ is the Steffan-Boltzmann constant, which has a value approximately . The energy flux incident at the Earth is the solar constant, but notice that the solar flux would not fall evenly on the surface of the planet. To account for this, not that the flux is spread uniformly over the projection of hemisphere facing the sun, that is the cross section, πr², but the energy (on average) is spread over the entire spherical surface, 4πr², giving a factor of 1/4. The temperature of this hypothetical Earth is then

T = (1367 / 4σ) ^{− 1 / 4} = 278.63K

In Celsius, that is equal to 278.63 - 273.15 = 5.48 C. That's a very cold globally averaged surface temperature! Luckily, the presence of the atmosphere allows the Earth system to store somewhat more energy; the observed global average temperature is approximately 287 K (14 C), allowing a more comfortable existence for humans.

The light that reaches the sun arrives over a broad range of frequencies, but the peak frequencies are in the visible portion of the electromagnetic spectrum. This can be shown by considering what we know already about the sun. For this exercise, two lengths are crucial, but can be measured quite well: the distance between the Earth and sun and the sun's radius. The average distance from the center of the sun to the center of the earth, which historically defined one astronomical unit, is about . The sun's equatorial radius is about m^[3]. Combining these distances with the solar constant (Φ_E = 1367Wm ^{− 2}) can provide us with an estimate of the sun's effective surface temperature. To derive the temperature, consider the amount of energy reaching the distance from the sun to Earth, i.e. the solar constant. Since the sun is spherical and radiates in all directions, that amount of energy reaches the same distance on the whole surface area of a sphere, with radius of 1 AU, call this radius R. The surface area of this sphere is given by 4πR². All this energy has to pass through the sphere defined by the radius of the sun as well, which means a shell with surface area 4πr² where r is the sun's radius. Because energy is conserved (it doesn't get created or destroyed between the sun and earth), the energy can be equated by

Φ_E4πR² = Φ_S4πr²,

where Φ_S is the energy flux at the sun's surface. With that flux, the sun's effective surface temperature can be deduced by T_Sun = (Φ_Sσ ^{− 1}) ^{− 1 / 4}. Plug in the numbers! We find that T_Sun = 5762K, which is very close to other estimates. Using Wien's Displacement Law, we can determine the maximum emission wavelength at this temperature as λ_max = b / T, where mK and is called Wien's displacement constant. This wavelength, which comes out to about 503 nm, is the wavelength at which most of the sun's energy is emitted. This value is close to the middle of the visible part of the spectrum, which is why we see the sun as a yellow-orange color when we look at it.

From this simple exercise, we deduce that the sun is close to a blackbody, which is confirmed observationally. What is the peak emission frequency from the Earth? What part of the spectrum is it in? To answer these questions, repeat the use of Wien's law above. The Earth emits at a much lower temperature, so a much longer wavelength, one that is thankfully invisible to human eyes (imagine how difficult it would be to see if the ground were like a lightbulb!).

[edit] The Sun-Earth System

In the above description, a simplified view of the Sun-Earth system has been used. The results are correct, and show how simple physical principles directly inform us about some of the most important aspects of Earth's global climate. However, in some situations this view is inadequate; this will become especially clear in the discussion of paleoclimate (ancient climate) and ice ages. Here let us review a few important aspects of the system to set the stage for later discussions.

Perhaps the most profound influence on the Earth-Sun system is the geometry involved. The most basic part of this geometry is the orbit of Earth around the sun, which is governed by the gravitational attraction between the two bodies. Kepler^[4] showed that orbits are ellipses, rather than perfect circles. Earth's orbit is slightly elliptical, with an eccentricity of just 0.01671 ^[5]; even though this value is small, it has important consequences. The perihelion, or smallest distance from the sun, occurs during the northern hemisphere's winter and is about , and aphelion, the most distant point from the sun, is during the northern hemisphere's summer and is about This difference does not cause Earth's seasons, but can influence the severity of seasons (discussed in the paleoclimate section) and does introduce small variations to the annual incoming solar radiation ("insolation") as there are very slow variations in the eccentricity.

A second important effect to consider is the tilt of Earth's spin axis with respect to the ecliptic plane, which is basically the average plane of Earth's orbit around the sun. The angle between the spin axis and the perpendicular to the ecliptic plane is called Earth's obliquity, and is currently about 23.4 degrees. This angle is the primary reason for seasons on Earth, for as the planet traverses its orbit, the amount of insolation at points on the surface slowly change, with winter occurring when the pole faces away from the sun and summer when the pole faces toward the sun. Seasons are more extreme with larger obliquity, and high latitudes (e.g. Antarctica) experience more extreme changes in insolation than the tropics, leading to more pronounced season. Earth's obliquity slowly changes in time, which has important consequences for very long-term climate change.

The ecliptic and its relation to rotation axis, plane of orbit and axial tilt.

A third important part of the Earth-sun geometry is called precession, and is actually the combination of parameters. Precession is the slow variation of direction of the spin axis, and is affected by both a turning of the spin axis and a slow change in the shape of Earth's orbit. For the contemporary climate, the precession only matters because it determines the relative position of the poles to the sun during Earth's orbit. There are important consequences for long-term climate change, though, which will be discussed later.

The geometry of the Earth-sun system is a large part of the astronomical basis for Earth's climate. Other astronomical factors that are important include the evolution of the solar system and the sun itself, as well as electromagnetic phenomena (e.g. the solar wind). These topics are well worth studying, even in the context of climate, but they are beyond the scope of this book, as they bear little relevance on contemporary climate change.

[edit] Distribution of Insolation

The previous section should make it clear that the geometry of the Sun-Earth system plays a key role in how much sunlight reaches Earth and where it arrives. This section briefly describes the pattern of incoming solar radiation, or insolation, over the course of the year. This pattern is the direct result of the geometrical factors described above, and since the insolation is the energy source for the entire climate system, where the energy enters the system is of fundamental importance to the subsequent distribution of energy. That is to say, where the sun shines has a direct impact on weather and climate.

In an idealized system, where the orbit is circular and the planet's spin axis is perpendicular to the ecliptic (i.e. obliquity equal to zero), the problem becomes nearly trivial because every day is identical. However, the spherical shape of the planet requires some consideration. A first guess at why the poles of this hypothetical planet are cold compared to the equator might be because the pole is farther from the sun than the equator (by a distance equal to the radius of the planet); in fact, this is a common mistake people make about the real Earth as well. To convince us that this can not be the case, consider the change in incoming energy in say 6800 km versus the Earth-sun distance: it is minuscule. However, if we allow the solar constant to be constant at the equator and the poles, there is still an important effect of geometry, namely the angle between the incoming photons (which we can think of as parallel rays of light) and the direction normal to the surface (which can be thought of as the local vertical direction, looking straight up, the sun is not always overhead). The insolation is reduced by the cosine of this angle, which is known as Lambert's cosine law. This is not an atmospheric effect, but simply an optical one, so consider this the insolation at the top of the atmosphere instead of the surface. The principle is simply that the curvature of the Earth means that the same radiance (or photon flux or sunshine) gets spread over a larger area as the angle between the photons and the normal to the surface increases from zero to 90 degrees at the poles. So even though the sun is just as bright everywhere in the world, the power per area incident at the top of the atmosphere changes as the sun appears lower in the sky. The result is a modification to the local insolation, I,

I(φ) = Scos(φ),

where φ is latitude.

Thus far, we have neglected Earth's spin. Of course, only half the planet faces the sun at any instant in time. Consider a snapshot in time. Half the planet faces the sun, half is in darkness. From everyday experience, we know that the sun appears at different distances above the horizon over the course of the day. At dawn, the sun comes over the eastern horizon, traveling in an arc across the sky, reaching its maximum height at local noon, and then descending toward the western horizon. Let us define the zenith as the point directly overhead; the angle from this zenith point to the sun's current position is called the zenith angle. In the hypothetical planet described above, where seasons do not exist, the sun only appears directly overhead along the equator at noon; moving away from the equator, the sun sinks lower and lower toward the equatorward horizon. At the south pole, the sun is just at the northward horizon at local noon, providing essentially no insolation. Define a new angle, that between Earth's equator and the highest local position of the sun (the position at local noon), and call this the declination angle; it is essentially a measure of the height above the horizon that the sun will reach each day, and is equal to the latitude at which the sun is directly overhead at noon. In the hypothetical world above, the conditions are perpetually equinox, so the declination is zero because the sun is overhead directly on the equator every day. Earth's obliquity is about (0.409 radians), which combined with the Earth's revolution of 360 degrees per 365 days provides an expression for the declination angle^[6][7],

where N is the Julian day of the year. This expression is an approximation, but will serve our purposes.

The distribution of incoming solar radiation with current orbital parameters, averaged over each day. Notice polar night is shown by the dark blue regions, reaching to the Arctic and Antarctic circles at the solstices. This distribution is calculated from the equations in the section, and clipped at zero since negative insolation does not exist.

With the latitude and declination, an estimate of insolation can be made given one additional piece of information related to the geometry of the system: the time of day. The time of day is needed since it affects the overall angle between the local vertical and the suns rays, as is clear by noting the difference between night, dawn/sunset, and noon. To describe the time of day, an additional angle in introduced, conveniently denoted the "hour angle,"

where, is the rotation rate (i.e., ) and LT is the local time. Note that at local noon, H = 0. These calculations are normally done in radians, but since most people more intuitively understand angles, we use them here.

Putting all these together, the actual solar zenith angle can be described in terms of the above angles

Z = cos ^{− 1}(sinφsinδ + cosφcosδcosH)

This relation can be derived without explicitly expressing the declination and hour angles, purely from the geometry. These angles make sense physically, so are included here.

Returning to Lambert's cosine law, we can write a simple expression for the insolation: I = ScosZ. This describes the flux of energy into the climate system at any particular moment, given the location and time of day.

[edit] At the Top of the Atmosphere

One may wonder where exactly is the top of the atmosphere, and with good reason. We know that the atmosphere consists primarily of the gaseous envelop around Earth, and that pressure decreases with height, according to they hydrostatic approximation. Does the atmosphere end only when the pressure reaches a vanishingly small value? No, but there is not a good definition of the top of the atmosphere, and it changes with sub-discipline. For our purposes, we can usually take the top of the atmosphere, often abbreviated TOA, as somewhere in the low to mid-stratosphere, or even simply the tropopause. In this chapter, we can imagine it is the level at which the downward shortwave radiative flux is negligibly different from the solar constant and where there is negligible downward longwave flux (that due to the sun, which is small).

Now that we have an idea of what the TOA is, we can ask why it might be useful.

First consider conservation of energy in an equilibrium system. This could be a tank of water with a heating lamp above it all enclosed in a box. It could be a simple blackbody system, or any isolated system. Conservation of energy means that the total amount of energy does not change, which is equivalent to saying that any energy that is input to the system must be balanced by an outward flux of energy. In the case of Earth (as the "system"), this means that the energy coming in (the sunlight) must be balanced by outgoing radiation. We know the solar constant (1367Wm ^{− 2}), so if we integrate over the Earth's surface, we know how much incoming energy there is. This incoming energy, sometimes called solar insolation or downward shortwave radiation, needs to be balanced. Why? Well, if it is not balanced by an equal loss of energy, then the temperature of the system must change (this is the 1st law of thermodynamics). Wien's law tells us that the wavelength of peak emission from a blackbody is inversely related to the temperature, and for normal Earth-like temperatures that puts the emission in the infrared part of the electromagnetic spectrum. Thankfully, this light is invisible to humans, and because the wavelength is longer than visible light (solar or shortwave), the terrestrial infrared radiation is often referred to as longwave radiation. The amount that is radiated to space (which differs from that emitted by the surface because of the greenhouse effect) is often called outgoing longwave radiation (OLR). The OLR (which is equal to the net longwave at TOA) balances the net shortwave at the top of the atmosphere when the system is in equilibrium.

Is the net shortwave radiation at TOA equal to the incoming shortwave? The answer is no. The net shortwave, which when averaged over suitable time and over the global, is the source of energy to the climate system, but not all the solar insolation is absorbed by the earth. Let's not beat around the bush. What happens to incoming solar radiation when it arrives in the atmsphere? There are really just three paths a photon (a "particle" of light) can take. First it can be absorbed, either in the atmosphere or at the surface. Absorbtion means that the energy associated with the photon is imparted to some atom or molecule, resulting in a higher energy level in that particle. Second, the photon can be reflected, which means that the path of the photon is reversed. More generally, we should say that the photon can be scattered, with some probability of being scattered back in the direction it came from, but we do not need to deal with scattering right here. Third, the photon can continue unimpeded, ultimately reaching the surface and being absorbed or reflected; while the photon is traveling through a medium without interacting, it is said to be transmitted. To study climate, one need not (usually) worry about individual photons, but the effects of the light in aggregate. Since we now know what can happen to each photon individually, we can sum over all the photons that make up the solar insolation such that , where F is the total downward shortwave flux, A is the fraction of light absorbed by the atmosphere, R is the fraction reflected back to space before reaching the surface, and T is the light transmitted to the surface.

From our understanding of the downward shortwave flux, we can continue the analysis by considering the surface. The amount of light absorbed by the surface is not exactly equal to the transmitted light, T. Why? Well, the surface can be highly reflective. For example, snow and ice reflect up to 80% of incident light, while open ocean surfaces reflect almost none. The reflectivity of the surface is usually called the albedo, denoted α, and is simply the fraction of incident light that gets reflected. Knowing that the surface has a given albedo, we can now say that the amount of light absorbed at the surface must be equal to . This says the absorbed light at the surface is equal to the transmitted insolation that is not reflected by the surface. Note that the albedo is constrained by definition to always be between 0 and 1, with typical global average of about 0.3.

It should also be noted that the shortwave light reflected by the surface does have a chance of being reflected (by clouds or particulate matter) or absorbed by atmospheric constituents. However, for most discussions of climate, and for the purposes here, we will neglect this process. Furthermore, we can (to a reasonable approximation) assume that the atmosphere is transparent to shortwave radiation, meaning there will be no absorption. This simplifies our previous expressions by eliminating the term A. To further simplify our notation, we can say that the total "planetary albedo" is the sum of the atmosphere albedo (later will will call this the cloud albedo) and the surface albedo, α_p = α_s + α_c. These simplifications allow us to write the "net shortwave at TOA" as

As mentioned above, the radiative flux from the surface acts approximately according to the Steffan-Boltzmann and Wien's laws of blackbody radiation. Even if we take the total energy from the solar constant spread over the full surface area of Earth, the emission must be in the infrared. The calculation is left as an exercise.

When we assume that the surface absorbs some fraction of the incident shortwave, (1 − α_s)F_sfc, and that the temperature reaches equilibrium, the emitted flux is then . Of course, in reality there is some emissivity associated with different surface types, but we neglect that here. Once emitted, these photons face similar consequences as the downwelling shortwave radiation. The difference in the longwave is primarily that the atmosphere is much more opaque in the infrared than visible, so the absorption can not be neglected. Various atmospheric constituents absorb infrared energy, then emit at a wavelength commensurate with the temperature of that part of the atmosphere. This is the natural greenhouse effect, and the active gases are often referred to as greenhouse gases; primary among these are water vapor, carbon dioxide, and methane.

The consequences of the natural greenhouse effect are crucial for life on Earth. In the absence of an atmosphere, the longwave radiation emitted to space would be exactly equal to the shortwave absorbed, and the surface temperature would be a chilly 255 K. Because greenhouse gases absorb infrared radiation, they act to warm the planet. How? We can think of the effects in two ways. First, the gases are heated by the absorbed radiation, and then radiate isotropically (equally up and down), sending energy back toward the surface to act as an extra energy source. Second, the absorption and subsequent emission by greenhouse gases changes the effective emission temperature of Earth (as seen from space). This second effect is a useful way to understand the greenhouse effect, and can be easily applied to changing climates. As a thought experiment, consider all the absorption by greenhouse gases happening in a thin layer of the atmosphere, which can effectively be thought of as a thin shell around the Earth. From space, the emission from the planet will be coming from an elevated level, with a much colder temperature than the surface. Of course, that means that the flux from that surface will be less than the incoming flux of solar insolation. The only way the climate system can achieve equilibrium, which is required by conservation of energy, is for the lower levels to warm, emit more energy as longwave radiation, which in turns warms the atmosphere, and changes the effective emission height and temperature. This adjustment continues until the shortwave and longwave budgets are balanced at the top of the atmosphere.

The figure below shows a rough cartoon of the radiative balances in the atmosphere. So far, we have focused only on the clear-sky scenario (on the left of the cartoon). Later we will consider the modifications that arise in the presence of clouds, and we will also explore the implications of changing the atmospheric composition.

The text in its current form is incomplete.

[edit] Balance at the surface

Climate Change/Surface Balance

[edit] Factors that influence climate

We have seen above that the vast majority of energy in the climate system comes from the sun in the form of electromagnetic radiation, and most of that visible light. From the discussion of blackbody radiation, we saw that the emission temperature of Earth was much colder than the actually global mean temperature. It is the presence of the atmosphere -- because of its ability to absorb the emitted infrared radiation -- that the surface temperature is a much more comfortable temperature. This process is called the greenhouse effect (sometimes modified as the "natural greenhouse effect"). What is it that allows the atmosphere to absorb infrared radiation, and for that matter, why is it so transparent to visible radiation?

To answer these questions completely requires a full treatment of radiative transfer and atmospheric chemistry, which are beyond the scope of this Wikibook. However, we can begin to understand these processes with basic physical principles.

[edit] Some topics to add

• albedo - surface, cloud, ice, and snow
• infrared absorption - water vapor, carbon dioxide, other greenhouse gases
• feedbacks - positive/negative, albedo, cloud, water vapor, vegetative, snow/ice

[edit] Present-day climate

Here we illustrate some important aspects of the current climate. (This should cover not just the late 20th century, but pre-industrial to present climate. )

[edit] Ancient Climate

This section deals with climate and its variability on very long, geologic, time scales. These time scales are greater than centuries, and there are important millennial and multi-millennial signals in the so-called paleo record. Some research focuses exclusively on ancient climates; such work is often referred to as paleoclimatology (also paleoceanography or paleogeography).

While a review of all relevant topics is beyond the scope of this Wikibook, some discussion of paleoclimate can shed light on the natural variability of the climate system. This includes not-so-distant climates, like that of the "little ice age," and also the global-scale oscillations known as ice ages. At time scales even longer than ice ages (10⁵ years), there are large variations of climate associated more with geological processes than atmospheric and oceanic. While intensely interesting, these climates will be ignored here.

[edit] Ice Ages

[edit] Decadal to Centennial Climate Variability

[edit] Climate Modeling

Climate models come in many forms, from very simple energy-balance models to fully coupled, three dimensional atmosphere-ocean-land numerical models. As computers have become faster, climate science has advanced commensurately. The equations that govern how fluids move in time and space (Navier-Stokes Equations) are complicated to solve, and when all the scales of motion and physical processes (radiative transfer, precipitation, etc) are incorporated, the resulting problem is impossible to carry out analytically. Instead, climate scientists turn these systems of equations into a series of computer programs. The resulting set of programs is, in some cases, a "climate model."

When the model is used to approximate the equations of motion on a sphere, it can be called a general circulation model (GCM). These are generic models, which can be specialized to simulate the ocean, atmosphere, or other fluid problems. Mostly because of limitations on computer power, these models do not resolve all scales of motion; instead a grid of points is established as an array of points where the equations are solved. Most modern atmospheric GCMs are run with horizontal grid spacing (distance between adjacent grid points) around 100 km, and with a number of vertical levels (usually around 30). The exact resolution depends on details of the model and the application. Because of this coarse grid spacing, small-scale (or "sub-grid-scale") phenomena (like individual clouds or even hurricanes) are not explicitly resolved. For detailed calculations of smaller scales, more specialized numerical models are often employed, though there are some very high resolution GCMs (e.g. Japan's NICAM^[8]). To incorporate the effects of sub-grid-scale phenomena, conventional GCMs rely on statistical rules, parameterizations, that describe how the processes work on average given the conditions within the grid cell. Parameterizations can be very simple or very complicated, depending on the complexity of the process and the level of understanding of the statistical behavior of the process. Much of the improvement in GCMs today is directly related to improving parameterizations, either by incorporating more elaborate rules to match measured quantities better or by using more sophisticated theoretical arguments for how the physics should work.

[edit] Kinds of Models

There are many classes of models, and within each class there are many implementations and variations. It is impossible to enumerate and describe every climate model that has ever been developed; even doing so for the published literature would be prohibitively difficult. In fact, there are entire volumes devoted to the history of numerical modeling of just the atmosphere; the American Institute of Physics has a brief description of AGCMs available online [9]. Here we discuss several classes of models, with an emphasis on atmospheric models. The discussion closely follows that of Henderson-Sellers and McGuffie (1987), which is an excellent resource on the subject (and has an updated edition).

First of all, we restrict ourselves to numerical models, specifically those designed to be solved with computers. More generally, any equation or set of equations that represents the climate system is a climate model. Some of these can be solved analytically, but those are highly simplified models, which are sometimes incorporated in numerical models.

The ultimate goal of climate models is to represent all physical processes that are important for the evolution of the climate system. This is a lofty goal, and will never truly be realized. The climate system contains important contributions and interactions among the lithosphere (the solid Earth), the biosphere (e.g., marine phytoplankton, tropical rainforests), atmospheric and oceanic chemistry (e.g., stratospheric ozone), and even molecular dynamics (e.g. radiative transfer). In fluid dynamics, some systems are now modeled using "direct numerical simulation" (DNS), in which (nearly) all the active scales are explicitly resolved. This is will never be feasible for the climate system. We cannot possible represent every atom in the climate system, it would essentially take the same number of electrons in the computer. Instead, climate modeling is limited to truly modeling the system; simplifying assumptions and empirical laws are used, the resolved motions are chosen to match the problem and/or the computing resources, and other processes are parameterized. However, these comprehensive climate models are not the only way to model the climate system.

Simpler models have been developed over the years for many reasons. One common reason historically was the computational cost of running large computers; simpler models have fewer processes to represent, and often have fewer space and time points (lower resolution). Two extremely simple classes of climate models are one-dimensional energy balance models and one-dimensional radiative-convective models. The single dimension in each is typically latitude (north-south direction) and altitude (vertical column), respectively.

A typical energy balance model (EBM) solves a small set of equations for the average temperature, T, as a function of latitude, T(φ). These models were introduced in 1969 by Budyko and Sellers independently. They are solved for the equilibrium temperature at each latitude based on the incoming and outgoing radiative fluxes and the horizontal transport of energy. The radiative fluxes are simple schemes (usually) for the radiation reaching the surface, and often include some temperature dependent albedo (reflectivity) to represent ice-albedo feedback. The horizontal transport is typically given by an eddy diffusion term, which is just a coefficient multiplied by the meridional (north-south) temperature gradient. One of the most interesting aspects of these simple models is that they already produce multiple equilibria, having solutions for ice-free and ice-covered Earths as well as a more temperate solution (like the current climate). This result spurred much research in the sensitivity of the climate system.

Radiative-convective models (RCM) are essentially models of an atmospheric column. They can be used to represent the global average atmosphere, a particular latitude (zone), or a particular location. The resolved dimension is vertical, so all the horizontal fluxes (like winds and advected scalars like temperature and moisture) must be passed to the column somehow. The early RCMs (due largely to S. Manabe and colleagues) have a background temperature structure (lapse rate) and a treatment of radiative fluxes throught he column. When the radiative heating of the column brings the lapse rate beyong a critical or threshold lapse rate, a "convective adjustment" is used to reduce the instability. Given constant boundary conditions, the model will equilibrate such that the energy budget is balanced, giving a model of the vertical (especially temperature) structure of the atmosphere. The early RCMs were used to explore the effects of increasing carbon dioxide in the atmosphere.

There are also combinations of EBMs with RCMs that give a simple two-dimensional representation of radiative-convective equilibrium.

Another class of two-dimensional model is the axially symmetric model used, for example, by Held & Hou (1980) and Lindzen & Hou (1988). This is a dynamical model only, and has been used to study the meridional circulation in the absence of baroclinic eddies (midlatitude storm systems). While not truly climate models, these simple dynamical models have provided important theoretical understanding of the atmospheric circulation.

In the ocean, there are simple box models that are somewhat analogous to the axially symmetric models of the meridional circulation of the atmosphere. These box models are traced back at least to Stommel, who used one to show the multiple equilibria of the thermohaline circulation in the Atlantic Ocean.

Other two-dimensional models also exist. For example there are simple equivalent barotropic models of the atmosphere. However these have mostly been used in numerical weather prediction and theoretical atmospheric dynamics.

Occupying a higher region of the modeling hierarchy are three-dimensional numerical models. In terms of dynamics, these are usually fully turbulent fluids, and can be applied to spherical geometry or some simplied geometry like the beta-plane. This class of models should probably be divided into several subclasses. Some are coupled models (atmosphere + ocean, for example) while others only contain a single component of the climate system. Some are described as climate models of intermediate complexity, which covers a large range of models.

At and around the top of the climate model hierarchy are general circulation models (GCM), sometimes called global climate models. These are fully three-dimensional representations of the atmosphere and/or ocean solved in spherical geometry. They are designed to conserve energy (1st law of thermodynamics), momentum (Newton's 2nd law of motion), mass (continuity equation), and (usually) moisture. We will discuss GCMs in much greater detail later, including the primary assumptions that they include, and the uncertainty associated with the results. GCMs are the best available tools for studying climate change.

[edit] What Models Tell Us

[edit] What Uncertainty in Simulations Means

Why can't climate models predict climate change perfectly? There are many answers to this question, and most of them are at least partly true! Here we briefly describe what is meant by "uncertainty" in climate modeling.

Before starting to describe the uncertainty associated with climate models, it is important to emphasize that climate models are the best tools currently available for studying the climate of Earth and other planets. Although they are far from perfect, sophisticated climate models embody the physical processes thought to be important in the real climate. That there is some uncertainty most decidedly does not mean that we can't trust climate model results, nor does it mean there is built in "wiggle room" in the models.

Different climate models, and here we want to imply sophisticated numerical models (usually of the whole globle), get different results for the same experiment. These differences are due largely to different ways of representing physical processes that happen on scales smaller than the distance between model points. These processes are usually called sub-gridscale processes, and the representations for them in numerical models are known as parameterizations. The main idea of a parameterization is that uses information from the large scale, and infers (based on some rules) what is likely happening on smaller scales. A good example of this is wind near mountains. GCMs might have grid points only every 100 km, but mountain ranges can have very drastic elevation changes over much shorter distances. Rather than try to represent the scale of the mountains, which would be very hard with current computers, GCMs have a sub-gridscale topography parameterization. Depending on the details, it may affect the "roughness" of the surface or gravity waves induced by terrain, but the idea is the same, given that mountains change height on small scales, the GCM tries to model that behavior, at least to capture how mountains can affect the large-scale circulations the GCM does resolve. Since parameterizations are different, and given the large number of parameterized processes, it is no wonder GCMs get different results. The fact that the results are not more different is a testament to our current level of understanding of climate processes.

Imagine taking a large number of GCMs and running them all the same way. For example, the IPCC had modeling centers around the world run the same experiments (basically increasing CO2 concentration) to compare each model's climate response. Because the models are built differently, in some cases the are very fundamental differences between models, the results vary from model to model. If we just take the climate response, for example the change in surface air temperature for a given change in radiative forcing, from each model and find the average and standard deviation, that gives us an estimate of the "uncertainty" in the climate response. This is done in lieu of a real experiment because we only have one Earth, and definitely not enough time to run so many experiments!

The above method gives a measure of expected climate response based on very different models. Another method is to use the same GCM, but slightly change parameters or even parameterizations to determine the strength of different processes. As a simple example, imagine that some GCM uses a single value for the albedo (reflectivity) of stratocumulus clouds. If the GCM is run ten times, each with a different value for that parameter, the results of a climate change experiment will change. How much the results differ will determine that GCM's sensitivity to stratocumulus albedo. This gives another measure of "uncertainty," since that model assumes there is only one value of the albedo, which may not be true in the real atmosphere. The distributed computing project ClimatePrediction.net uses such a methodology to study processes important to climate sensitivity.

Another answer to our original question is that the system is not perfectly predictable. The climate system is chaotic, or at least "sensitive to initial data." This just means that we know the equations that govern fluid motion, and we have a pretty good idea of the physical processes that need to be included, but the system of equations has many solutions, and even if the system is perfectly deterministic (no random fluctuations), unless we also perfectly know the initial conditions, we may not get the right answer. In fact, in chaotic systems it has been shown that arbitrarily small errors in the initial conditions can give wildly different results after some amount of time. While the case of Earth's climate is unlikely to be that sensitive, it does mean that we shouldn't expect a perfect long-term (greater than 2 weeks) weather forecast to be on the local television station any time soon (or ever). Note that the science of climate prediction differs fundamentally from that of weather prediction. In weather prediction, the sensitivity to initial conditions is a basic limitation, as perfect knowledge of initial conditions is impossible. Climate models are not sensitive to initial condition; the problem changes from an initial-value problem to a boundary-value problem.

[edit] What geology, hydrology, and biology tell us about climate

[edit] Geology

• Sedimentary rocks: The presence of sedimentary rocks is direct evidence of ancient climatic change. The atmosphere slowly developed in its current composition from very different states related to geology and biology. Geological processes, especially weathering due to wind and rain, helped transform primary igneous rocks to sediments. Those sediments are deposited downstream, and over time they are compacted and form sedimentary rocks. Tectonic activity can also contribute to the process, acting to accelerate compaction of sediment.
• lake sediments
• ocean floor sediments

Tectonic activity has been an abundant source of carbon dioxide in the atmosphere for our earth's history. A single volcanic event can transfer as much carbon dioxide to our atmosphere as our whole nation's output for a year. This tectonic phenomenae is credited for the creation of our original atmosphere. A case can be made that the earth's carbon is simply recycled from original volcanic origin to life forms and then later to limestone in sedimentary rock accumulations. Vast deposits of limestone are known from the Paleozoic era. Increased carbon dioxide in our atmosphere may result in an increase of limestone deposition in our oceans.

[edit] Biology

• ecosystems
• dendrochronology
• coral reefs
• rainforests
• desertification

[edit] Evidence that the climate is changing

By using various methods of chemistry, geology, and even astronomy, past climate variations are well known. These include the relatively periodic ice ages of the past 2-million years as well as more exotic climates from the Cretaceous and other eras. On such long timescales the main reasons for climate changes must be linked to changes in the sunlight reaching Earth (insolation), major shifts in ocean heat transport, or "external" forcing like volcanism or meteor impacts. More recently, human-induced changes (anthropogenic) are likely a strong climate forcing. Much research has focused on quantifying Earth's climate and its variability over the past 30 years or more. It has been shown that the globally averaged surface temperature is now warmer than it has been for at least 150 years. The trend in surface temperature is remarkably well correlated with a trend in atmospheric carbon dioxide. The current trends are increasing, and the warmest years on record are in the last decade. IPCC

Indeed, over the past century or so the global (land and sea) temperature has increased by approximately 0.6 ± 0.2°C

Earth is now absorbing 0.85±0.15 W m^-2 more energy from the Sun than it is emitting to space. (Hansen et al, 2005, Science vol 308)

Some of the observed evidence that Earth's climate is changing:

• Global temperatures are rising
• Melting glaciers
• Sea levels rising

[edit] Evidence for Anthropogenic Climate Change

Wikipedia has related information at Attribution of recent climate change

Wikipedia has related information at Greenhouse gas

Many believe that CO₂ and other greenhouse gases (chlorofluorocarbons, methane, sulfur hexafluoride) cause global warming.

• Observed trend in global mean surface temperature
• Observed radiative imbalance at top-of-atmosphere
• Rising atmospheric concentration of CO₂
• Rising sea level due to thermal expansion of sea-water.

[edit] Criticism of Anthropogenic Attribution

Some people, for a variety of reasons, claim to have found faults with the hypothesis that humans are affecting Earth's climate. While we strive to present any legitimate criticism of the scientific principles where they are presented, this section includes some specific issues that are commonly cited as reasons that humans are not or could not change the climate.

• Lack of scientific consensus

One of the most common arguments against human induced climate change is a supposed lack of scientific consensus. While there were many skeptical scientists in the past, as the evidence has mounted (especially using satellite-based data), even the most ardent skeptics have come to the determination that humans are changing the climate. One recent study found no instances in the peer-reviewed literature of a study on climate change stating that global warming is either fictitious or purely natural Oreskes, N., Science, 2004, 306(5702), DOI: 10.1126/science.1103618. See also Scientific Consensus.

• The Earth's surface obtains energy from four primary sources: space (predominantly solar radiation), the molten core of the Earth, anthropogenic processes that generate excess heat, and radiation from the atmosphere. The second (geothermal heat) is known to be trivially small; the third (direct excess heat) is not as important as increases in the fourth, the "greenhouse effect" due to increases in CO2, etc. While it is nice to think that changing our energy consumption habits will stop global warming, it could very well be that climate change is being driven by processes that we have little control over.
• Please note: Solar radiation varies over time as the orbit of the earth changes due to gravitational inter-action with the other planets and the sun. When the earth's orbit gets more elliptical and most of the land mass is in a seasonal orbital position so it is receiving more direct radiant energy from the sun (or it is summer for most of the land mass when the orbit brings it nearest the sun), then an inter-glacial period usually occurs. Those combining planetary science with geologic evidence have significant findings suggesting that our present inter-glacial period has not peaked. Some pointing to an inter-glacial period about 400,000 years ago that had about 1/3 of the ice on the Antarctic gone when it peaked, as having the most similar pattern of orbits for the planets when compared to the orbits now.
  • Others, like Bill Ruddiman (U VA) think that we are overdue for an ice age, based on orbital parameters.
  • Also note that the current configuration actually has Earth closest to the sun during northern hemisphere winter, and not summer. Seasons are not due to the eccentricity (how "oval" the orbit is), but really much more on the tilt of the spin axis of Earth (obliquity). There is a precession signal as well, which is influenced by the sun-earth distance, but that signal is more directly linked to the Tropics.

[edit] Predicted outcomes

• Global mean surface temperature rises
• Global mean sea surface temperature rises
• Surface temperature over land increases more than over ocean
• Global mean precipitation increases, with a larger increase over the ocean.
• Sea-ice concentration decreases, melt season comes earlier, freezing starts later
• Regional temperature and salinity changes
• Sea level rises, mostly due to thermal expansion of ocean water
• More frequent and severe floods and droughts

[edit] Possible effects on humans

• Increased incidences of respiratory infections, malaria, and other diseases.
• Changing growing seasons lead to bad harvests and widespread food shortages
• Increased droughts cause regional water shortages
• Sea levels rise more than expected (>1m), displacing millions of people
• More frequent flooding leads to coastal destruction in developing nations, increasing the spread of diseases locally.

[edit] Worst-case scenario outcomes

• Release of methane clathrate from ocean bottom releases enormous amounts of methane to atmosphere, leading to runaway greenhouse effect
• Thawing tundra releases methane trapped in permanently frozen organic matter, leading to enhanced warming
• Increases in precipitation over Greenland, combined with other warming effects there, leads to pools of liquid water that melt into the ice sheet as moulins. Liquid water gets deep into the ice sheet, lubricating and destabilizing it, and huge discharges of ice spill into the north Atlantic.
• Huge discharges of ice spill into the north Atlantic, chilling and freshening the surface water, stabilizing the upper ocean. This shuts down deep convection, and we experience a rapid climate change (not quite so fast as The Day After Tomorrow).

[edit] Climate Change and Society: Mitigation strategies

In this section, we assume anthropogenic climate change is well underway and examine what individuals and social structures can do to slow or reverse the trend.

[edit] Community contributions to mitigation

Wikiversity has course materials on Environmental community building.

This section highlights some current thinking about how communities might deal with a changing climate. These include encouraging people to live closer to where they work, building up (not out), increasing use of public transportation, increasing recycling programs, more efficient use of water resources, more renewable energy sources, and even "green" urban planning.

• Cradle to Cradle: Remaking the way we make things
• Pacala S, Socolow R, 2004: Stabilization Wedges: Solving the Climate Problem for the Next 50 Years with Current Technologies. Science, Vol. 305, No. 5686, pp. 968-972.LINK

[edit] Individual contributions to mitigation

Here we discuss how personal conservation can help to limit greenhouse gas emissions.

• Carbon coaching ^[1]
• In the office - computers, printers, windows, etc.
• In the kitchen - appliances, water use, etc.
• Transport - gasoline, diesel, biodiesel, hybrids, and even bikes/walking.
• Lighting - compact fluorescent bulbs, LEDs, OLEDs
• Heating & Cooling
• Alternative energy - Roof-mounted solar panels, personal wind turbines etc
• Alternative propulsion - Plug-in hybrids.
• Carpool

[edit] How societies/countries can reduce emissions

Wikiversity has course materials on Juridical national measures on climate change.

• Carbon sinks - including sequestration
• Public transport - efficient urban planning, trains, buses, and shared resources
• Renewable energy - solar, wind, tidal, geothermal
• Nuclear power - fission now, fusion later?
• Economics of fossil fuel - Hubbert's peak and making it a plateau (maybe)

[edit] Appendices and miscellany

• Press cuttings - the sub-page where we should post press cuttings
• Further Reading - References and recommendations... non-technical, text, and scientific work to supplement this wikibook.

[edit] External links

• National Energy Foundation Home
• CO₂ calculator
• Simple ways to save energy
• National Energy Foundation - Urban Myths or Simple Truths?
• MIT's Global Change website
• WikiPedia: Global Warming
• A distributed computing project to predict climate change
• Pew Center on Global Climate Change - news, explanation, policy, and activism
• The Climate Ark is a Climate Change Portal and Search Engine
• Introduction to climate change... good for younger audience too.
• US Global Change Research Program
• Canada's introduction, including Kyoto Protocol information
• US Climate Change Science Program
• Local effects of climate change
• RealClimate
• Click 4 Carbon Online Community, blog and Carbon Emissions Calculator

[edit] References

1. ↑ http://www.carboncoach.com

1. ^  M. Satoh, T. Matsuno, H. Tomita, H. Miura, T. Nasuno and S. Iga, "Nonhydrostatic icosahedral atmospheric model (NICAM) for global cloud resolving simulations," Journal of Computational Physics Volume 227, Issue 7, 20 March 2008, Pages 3486-3514
2. ^  Sun
3. ^  Svante Arrhenius [NASA]
4. ^  The text of Marsh's manuscript is publicly available [Gutenberg].
5. ^  Johannes_Kepler
6. ^  Eccentricity is defined for all conic sections, and is a relationship between the semimajor (a) and semiminor (b) axes. It can be determined by

For a perfect circle a = b, so the eccentricity is zero, for an ellipse a > b, and the eccentricity is bounded 0 < ε < 1. See also [Wolfram MathWorld].

1. ^  Cooper, P.I., Solar Energy, 12, 3 (1969). "The Absorption of Solar Radiation in Solar Stills."
2. ^  [D. Fletcher]