Climate Change: The Science of Global Warming and our Energy Future

THE GROWING AMOUNT OF carbon dioxide (CO₂) in the atmosphere is the primary cause for global warming (chapters 7 and 9), so it is important to understand the factors that naturally control its abundance. Atmospheric CO₂ levels have been measured continuously only since 1958. In that year, Charles Keeling (1928–2005), who would go on to become a professor at Scripps Institution of Oceanography, began measuring CO₂ at a newly established observatory on the top of Hawaii’s Mauna Loa volcano.¹ The change in CO₂ content with time, widely referred to as the Keeling curve, shows two interesting features (figure 4.1). First, every year CO₂ concentration reaches a maximum in May and then decreases until October, when it begins to rise again. The reason for this change is that in the Northern Hemisphere early summer plants begin to grow rapidly, drawing down CO₂, but in the fall they become dormant or die and decay, returning CO₂ to the atmosphere. This process is like the natural rhythm of breathing, but on a much grander scale.

The second feature of the Keeling curve is equally obvious. Atmospheric CO₂ content has been rising at a rapid clip since Keeling started measuring it. In 1958, the average CO₂ content of the atmosphere was 315 parts per million (ppm) by volume; in 2006, it was 380 ppm and increasing by about 2 ppm per year. This rate of increase is far greater than any common natural phenomenon can produce and is unprecedented in the climate record. Emissions of CO₂ and other greenhouse gases as a consequence of human activity—burning of fossil fuel, cement production, and other activities, together commonly referred to as anthropogenic emissions—account for this rise (chapter 10).

But there is one curious statistic: only about one-third of the carbon emitted from all human activities and one-half of that emitted from fossil-fuel burning and cement production since 1800 are in the atmosphere. Where has the rest gone? The answer is that it has been removed by the ocean and the biosphere and is now stored there. This brings us to the subject of this chapter, the carbon cycle, which refers to the flow of carbon among the various global “reservoirs” that store carbon. The chapter explores how the ocean and biosphere control the CO₂ content of the atmosphere and how the ocean is particularly affected by the increase in atmospheric CO₂.

The CO₂ content of the atmosphere has been rising since the measurement at Mauna Loa volcano began in 1958, when it averaged 315 ppm. In 2006, CO₂ content was 380 ppm and rising by about 2 ppm per year. The annual fluctuation of CO₂ level reflects the natural cycle of plant growth and death in the Northern Hemisphere. (After Scripps Institution of Oceanography CO₂ Program, http://scrippsco2.ucsd.edu, with permission)

To understand the carbon cycle, we must first characterize the carbon reservoirs. One might suppose that most carbon on Earth is present in living organisms or perhaps in the form of coal or oil. That, however, is emphatically not the case. More than 99.9 percent of the carbon on or near the surface of Earth is sequestered in rocks, mostly limestones, carbon-rich shales, and coal and oil deposits. The rock “reservoir” amounts to some 50 million billion metric tons (50 million gigatons, or 10¹⁵ metric tons) (table 4.1).

In contrast, all the surface reservoirs combined contain only about 42,000 gigatons of carbon. By surface reservoirs, we mean the atmosphere, where carbon exists mainly as CO₂; soil, where carbon occurs in a variety of organic compounds; the ocean, where it is present mainly as the dissolved bicarbonate ion (HCO₃^–); and, of course, living biomass. The largest by far of these surface reservoirs is the ocean, which holds 39,000 gigatons, a fact that, as we shall see, has an important bearing on climate. The other reservoirs are minuscule in comparison; for example, all the biomass contains a piddling 600 gigatons of carbon, even less than the atmosphere!

TABLE 4.1 SIZES OF CARBON RESERVOIRS
Rock reservoir	50 × 10⁶ metric gigatons
Limestone	40 × 10⁶
Organic carbon in sedimentary rocks	10 × 10⁶
Fossil fuels	4.7 × 10³
Marine carbonate sediments	2.5 × 10³
World ocean	39 × 10³
Bicarbonate ion	37 × 10³
Carbonate ion	1.3 × 10³
Dissolved CO₂	0.74 × 10³
Organic carbon in soils and terrestrial sediments	1.6 × 10³
Organic carbon in permafrost	0.9 × 10³
Atmospheric CO₂	0.76 × 10³
Living biomass	0.6 × 10³
Sources: L. R. Kump, J. F. Kasting, and R.G. Crane, The Earth System, 2d ed. (Upper Saddle River, N.J.: Prentice Hall, 2004); S. A. Zimov, E. A. G. Schuur, and F. S. Chapin III, “Permafrost and the Global Carbon Budget,” Science 312 (2006): 1612–1613.

It is convenient to think of the carbon cycle as two cycles rather than one. The long-term carbon cycle operates over hundreds of thousands to millions of years, but carbon also cycles through surface reservoirs in timescales measured only in months to years to centuries. The short-term carbon cycle bears directly on the state of the present climate and the way it changes over human timescales, but we also have to understand the long-term carbon cycle to appreciate our present predicament.

The long-term carbon cycle involves the movement of carbon between the solid Earth and the ocean and atmosphere. The weathering of rock, which is especially rapid in mountainous regions where erosion rates are high, removes CO₂ from the atmosphere. This happens because CO₂ reacts with surface water and with the silicate and carbonate minerals of rocks. The result is the formation of calcium (Ca), magnesium (Mg), bicarbonate (HCO₃^–), and silica (SiO₃, SiO₄) in solution in the water (H₂O), which may be represented by a simplified reaction:

4CO₂ + 6H₂O + CaSiO₃ + MgSiO₃ → Ca⁺⁺ + Mg⁺⁺ + 4HCO₃ + 2H₄ SiO₄

atmospheric carbon dioxide + water + calcium- and magnesium-bearing silicate minerals → dissolved calcium and magnesium cations + bicarbonate ions + silicic acid

Coquina is a form of limestone composed of fossilized shell debris. The shells and the cement that holds them together are made of the calcium carbonate mineral calcite. This sample is on display in the Gottesman Hall of Planet Earth at the American Museum of Natural History, New York. (Photograph by D. Finnin, American Museum of Natural History)

Chalk, a variety of limestone, makes up the White Cliffs of Dover, in southeastern England. The chalk in the cliffs formed about 70 million years ago by the accumulation of the calcium carbonate skeletal remains, known as coccoliths (see figure 4.9), of single-celled, microscopic algae. (Photograph by Maki Itoh, http://www.makikoitoh.com, with permission)

As shelled organisms die, their carbonate shells accumulate on the ocean floor to form limestone (figures 4.2 and 4.3). New layers of sediment cover older layers, and with burial the sediments slowly transform to rock. The rocks are sometimes buried to depths of thousands of meters, a process that may take millions of years, in which case the heat and pressure of the deep Earth may break down the carbonate minerals to give off CO₂. The CO₂ slowly percolates out of the crust and back into the atmosphere, completing the cycle. Alternatively, the sedimentary rocks may be dragged down into Earth’s mantle in subduction zones, where one of Earth’s lithospheric plates is thrust beneath another. Here the carbonates similarly break down, and the CO₂ finds its way back to the surface as part of the gas emanating from erupting volcanoes. In either case, a reaction for the return of carbon from the deep Earth to the surface reservoirs is

Coal is the remains of decayed plants buried in swamps and marshes and then subjected to high temperature and pressure. This metamorphism from dead plants to coal, which typically occurs over 1 million years or more, drives off water and most hydrocarbons, leaving coal as the carbon-rich residue. This sample is on display in the Gottesman Hall of Planet Earth at the American Museum of Natural History, New York. (Photograph by D. Finnin, American Museum of Natural History)

A variant of the cycle is the removal of CO₂ from the atmosphere by plants as they engage in photosynthesis and then the direct burial of dead plant matter. This process commonly occurs in swamps, where burial of the plant matter may eventually produce coal beds (figure 4.4), or in river deltas, where plant matter may be buried in mud to form carbonaceous shale. The reaction for photosynthesis, which is driven by sunlight, may be represented by

Now we have to return to the notion of carbon reservoirs. The enormous mismatch in the amounts of carbon held in the rock and in the surface reservoirs has a very important implication—that over millions of years, there must have existed a close balance in the exchange of carbon between them. The amount of carbon that is being put into the rock reservoir through the formation of carbonate-bearing and organic-carbon-bearing rocks must over the long term be balanced by the amount returned by degassing as the rocks get buried or subducted and thermally decompose. Indeed, this long-term balance may be why Earth has not experienced a runaway greenhouse, in which so much CO₂ accumulates in the atmosphere that the planet heats up to the extent that it cannot support life. In other words, the long-term carbon cycle appears to have acted as a natural thermostat, maintaining conditions on Earth’s surface conducive to the evolution and survival of life since nearly the beginning.

However weighty this thought might be, the important point in the context of climate change is that fossil-fuel consumption and cement production represent a rapid transfer of carbon from the rock reservoir to the surface reservoirs. This street goes only one way as far as the climate system is concerned because carbon returns to the rock reservoir only by natural processes that operate much more slowly than the rate at which we are removing carbon from this reservoir. This, broadly, is the fundamental problem: by burning fossil fuel, we are removing carbon from the long-term reservoir and putting it into the short-term, surface reservoirs, but natural processes cannot return it to the long-term reservoir quickly enough.

By examining the geological record, we can gain a sense of how long it takes for a large excess of carbon to be drawn down into the rock reservoir. About 55 million years ago, a sudden influx of CO₂ into the atmosphere–ocean system sent temperatures soaring at a time known as the Paleocene–Eocene Thermal Maximum (PETM) (chapter 6). The recovery of the climate system and return to cooler conditions took 40,000 to 60,000 years, which may be the shortest period of time that the long-term carbon cycle requires to correct sudden perturbations in the climate system.

The short-term carbon cycle refers to the circulation of carbon among the surface reservoirs—the ocean, atmosphere, soil (land), and biosphere (figure 4.5). It does not involve rocks at all, and the cycling can be rapid, taking only from months to decades to a millennium. Because the circulation is rapid, the short-term cycle is immediately relevant to climate over the next several decades to millennia. In the land-based part of the short-term carbon cycle, photosynthesis removes carbon from the atmosphere (as described earlier). About half the terrestrial organic material from photosynthesis is returned to the atmosphere by respiration of plants and animals, including ourselves:

The burning of fossil fuels represents a net transfer of carbon from the long-term rock reservoir to the short-term surface reservoirs. The transfer is irreversible on timescales relevant to humanity unless we find ways of putting carbon back into the rock reservoir. Gt C = billion metric tons of carbon. (Data from table 4.1 and NASA)

whereas the remainder ends up in soil by accumulation of dead debris. Microbes in the soil engage in aerobic respiration to produce CO₂ as well. Or, if they live in deeper, oxygen-poor levels of the soil, microbes may produce methane (CH₄):

The resulting methane enters the atmosphere but eventually combines with oxygen (oxidizes) to form more CO₂. The methane content of the atmosphere is a measure of the terrestrial methanogenesis rate and, indirectly, of the extent of wetlands in mainly tropical and temperate climates.

The average carbon atom resides in the atmosphere for about a decade,³ which is about how long it takes for the atmosphere to respond to changes in the rates of carbon flowing in and out of the biosphere. The natural flows of carbon in and out of the surface reservoirs balance themselves, but adding to the natural flow is a steady trickle of carbon in the form of anthropogenic CO₂ (see figure 4.5).

Although the atmosphere’s interactions with the land-based biosphere have an immediate (that is, months to decades) effect, they are not really the main story. On the longer timescale of several decades to a millennium and more, the ocean is the primary regulator of CO₂ in the atmosphere. Indeed, the ocean has taken up about one-half of all the carbon originating from fossil-fuel emission and cement production for the period 1800 to 1994 and about one-third of it for the two decades from 1980 to 2000 (table 4.2);⁴ climate models suggest that 90 percent of anthropogenic CO₂ will eventually find its way into the ocean.⁵

Several mechanisms are involved in the exchange of carbon between the atmosphere and the ocean. First, photosynthetic organisms known as phytoplankton live in the photic zone (the region penetrated by light, extending to about 100 meters [330 feet] below the ocean’s surface). Phytoplankton are the lunch of various zooplankton, such as foraminifera, radiolarians, and copepods. The zooplankton produce fecal pellets, which rain down into the ocean depths, where the organic material is consumed by other microorganisms, releasing CO₂ (and nutrients) to deep water. The combined processes of photosynthesis in the shallow ocean and the CO₂ transfer to and enrichment of the deep waters is known as the biological pump. The biological pump is particularly effective at high latitudes where surface waters are relatively rich in nutrients⁶ and thus contain abundant phytoplankton (figure 4.6).

TABLE 4.2 WHERE ANTHROPOGENIC CARBON HAS COME FROM AND WHERE IT GOES

Source: C. L. Sabine, R. A. Feely, N. Gruber, R. M. Key, K. Lee, J. L. Bullister, R. Wanninkhof, C. S. Wong, D. W. R. Wallace, B. Tilbrook, F. J. Millero, T.-H. Peng, A. Kozyr, T. Ono, and A. F. Rios, “The Oceanic Sink for Anthropogenic CO₂,” Science 305 (2004): 367–371.

Surface ocean waters at high latitudes (light green) are relatively rich in nutrients and thus contain abundant phytoplankton. The chlorophyll concentration scale refers to the ocean; the normalized difference vegetation index is a measure of the density of land vegetation. (Satellite imagery courtesy of GeoEye/NASA. Copyright 2008. All rights reserved, http://oceancolor.gsfc.nasa.gov/)

Another biological pump, sometimes referred to as the carbonate pump, involves the removal of carbon from the ocean when foraminifera, mollusks, and other organisms build calcium carbonate (CaCO₃) shells, as described earlier. When these marine “calcifiers” die and their shells settle to the deeper ocean, one of two things happens. The shells may reach the bottom, accumulate as sediment, and eventually become part of the rock reservoir, with the carbon removed from the surface reservoirs for millions of years. Alternatively, because carbonate is not stable in the deep ocean, shells that settle below the carbonate compensation depth (CCD) (about 4,000 meters [13,000 feet]) dissolve and release CO₂ faster than they accumulate on the bottom as sediment.

It may seem counterintuitive, but the reaction in the ocean that produces carbonate for shells also produces CO₂, as illustrated by the second reaction in the long-term carbon cycle. So, despite the fact that the carbonate pump may transfer carbon to the deep ocean or even to the rock reservoir, in those parts of the ocean where calcifying organisms are active, the ocean actually pumps CO₂ into the atmosphere rather than removing it.

A third pump for carbon is called the solubility pump. It turns out that the solubility of CO₂ (and most other gases) in water increases as temperature decreases and pressure increases.⁷ For this reason, the deep ocean can hold substantially more CO₂ than can the shallow ocean. The solubility pump refers to the cooling and sinking of surface water at high latitudes, thus removing CO₂ from the atmosphere and transporting it to the deep ocean.

Like the biological pump, the solubility pump operates mainly at high latitudes, where CO₂-rich surface waters tend to sink as they cool, making these regions particularly important in the regulation of atmospheric CO₂. The deep, CO₂-rich waters that form at high latitudes may take many decades to centuries and longer to well up to the surface, depending on location. Upwellings of deep water occur along certain coastlines in equatorial regions (figure 4.7). As the deep water rises and warms, it gives up CO₂ to the atmosphere.

The ocean therefore both removes CO₂ from the atmosphere and adds it back, and over the course of years the transfer back and forth establishes a CO₂ balance between the two.⁸ But because the ocean contains so much more carbon than the atmosphere, it is the 800-pound gorilla in this story and exerts the main control in the transfer process.

Because the atmosphere and the ocean are in chemical equilibrium, as the CO₂ content of the atmosphere increases, so does that of the ocean. This equilibrium has what may turn out to be an important consequence: the CO₂ taken up by the ocean goes through a series of reactions that reduce the pH of water (pH is a measure of acidity; the lower the pH the higher the acidity):⁹

The ocean removes CO₂ from the atmosphere in the North Atlantic Ocean, in western parts of the South Atlantic Ocean and the Pacific Ocean between about 40 and 50°S latitudes, and throughout the Southern Ocean around Antarctica. It gives up CO₂ to the atmosphere mainly in the equatorial Pacific Ocean. (After Takahashi et al. 1997)

CO₂(aq) + H₂ O → H₂CO₃

dissolved CO₂ + water → carbonic acid (1)

H₂CO₃ → HCO₃ + H⁺

carbonic acid → bicarbonate + hydrogen ions (2)

HCO₃^– → CO₃^2–+ H⁺

bicarbonate → carbonate + hydrogen ions (3)

In preindustrial times, ocean pH was about 8.2. Now it is 8.05, and if the CO₂ content of the atmosphere doubles, it will decrease to about 7.9.¹⁰ (It is important to understand that the ocean is not actually acidic, which refers to compositions for which pH is less than 7; rather, it is becoming less basic because its pH, although decreasing, remains well higher than 7.) Although the pH change may seem small, the acidification is likely to have a profound influence on ocean biogeochemistry, particularly on calcifying organisms. As pH decreases (H⁺ concentration increases), the amount of carbonate (CO₃^2–) in seawater decreases (reaction 3), which has the effect of decreasing the stability of calcite and aragonite, the carbonate minerals that constitute the skeletons and shells of calcifying organisms (but this decrease appears to depend on the organism).¹¹

Because CO₂ solubility in ocean water increases with decreasing temperature and increasing pressure, there are depths in the ocean above which calcite and aragonite (both of which are composed of calcium carbonate but differ in how the atoms are put together) are stable and below which they are not.¹² Furthermore, these depths will migrate to shallower water depths as ocean acidification proceeds. The specific depths at which carbonates become unstable (saturation depths) differ from one location to another (figure 4.8). In the North Atlantic, for example, the depths exceed 2,500 meters (8,200 feet), but in the North Pacific they rise to within a couple of hundred meters of the surface. One of the consequences of CO₂ buildup and acidification is that the carbonate-saturation depths will continue to rise closer to the surface. If CO₂ emissions continue to increase, calcite and aragonite may become unstable in the surface waters in the Southern Ocean beginning sometime around the mid-twenty-first century, and no shelled organisms will live there.¹³

The saturation depth for aragonite (a) is shallower than the saturation depth for calcite (b) because aragonite is more soluble in seawater. The depths are sensitive to ocean water acidity. The color scales on the right are in meters of water depth. (After Feely et al. 2004, with permission)

The changes occurring in the ocean lead to two important questions, neither of which has a clear answer at this point. Will ocean acidification influence the ocean’s ability to absorb CO₂ from the atmosphere and thus affect climate? And how severely will acidification affect marine calcifiers and disrupt marine ecosystems?

Regarding the first question, if less carbon is being converted into shells, there should be a concomitant decrease in CO₂ production in ocean waters (the second reaction in the long-term carbon cycle) and thus an increase in the net flow of CO₂ from the atmosphere into the ocean. Calculations suggest, however, that the increased rate of CO₂ removal from the atmosphere will be small compared with anthropogenic CO₂ emissions.¹⁴

Or there may be feedbacks (chapter 5) associated with the ocean’s changing biochemistry that we simply do not know about, some of which may have the opposite effect. For example, it is possible that as CO₂ in the atmosphere increases, the rate of carbon uptake by the ocean (and land) will slow, resulting in a progressive increase in the fraction of anthropogenic CO₂ emissions that remain in the atmosphere, thereby also amplifying warming.¹⁵

Increasing CO₂ content and decreasing pH of the ocean may influence marine organisms in several ways. Most obvious is the negative influence on the ability of calcifying organisms to reproduce and to build and maintain their shells, as noted earlier.¹⁶ Of particular concern is how the ocean’s changing composition will affect the ability of corals to build their skeletons, which is necessary to maintain reef structures, not to mention entire reef ecosystems. Corals are sensitive to ocean temperature, pollution, and other factors. For example, the particularly warm 1997/1998 El Niño year destroyed 16 percent of the world’s coral reefs, and many more are under serious threat from the combined stresses related to climate change. Several studies suggest that corals will be unable to build their carbonate reef structures as the CO₂ content of the atmosphere approaches 480 ppm, a level that will be reached within 50 years at the current rate of emissions.¹⁷

The organism Emiliania huxleyi secretes and is armored with delicate calcium carbonate coverings known as coccoliths. It lives near the ocean surface, where it photosynthesizes. Such shell-building microorganisms are important for removing carbon from the ocean, and how they will fare as ocean pH decreases is uncertain. (Photomicrograph by Jeremy Young, http://www.noc.soton.ac.uk/soes/staff/tt/eh/index.html)

As for phytoplankton, the few existing experiments suggest that increasing ocean CO₂ content and declining pH increase photosynthesis only marginally in most phytoplankton.¹⁸ There is at least one important exception, however—the abundant phytoplankton Emiliania huxleyi (figure 4.9). For this organism, both photosynthesis and calcification rates increase markedly with increasing CO₂.¹⁹ According to reaction 3, increasing hydrogen ion (H⁺) concentration decreases carbonate ion (CO₃^2–) concentration, but at the same time increases bicarbonate ion (HCO₃^–) concentration. The bicarbonate ion is the source of carbon for carbonate formation; at least for Emiliania huxleyi, then, the effect of increasing bicarbonate ion concentration apparently more than offsets that of decreasing carbonate ion concentration.

Finally, a reduced pH should increase the concentrations of dissolved copper, zinc, and other toxic metals as well as nutrients dissolved in seawater.²⁰ There is little research on this topic, so how these competing factors will ultimately affect marine organisms’ health is not known.

The numerous unknowns in how the carbon cycle operates translate into a huge question about how climate will respond to the 36 gigatons (and growing) of CO₂ we are dumping into the atmosphere every year as a consequence of our activities.²¹ First, one needs to appreciate the single, overriding uncertainty that concerns all elements of the carbon cycle. The atmosphere contains more CO₂ now (385 ppm) than at any time in at least the past 800,000 years (chapter 6). In the past, atmospheric CO₂ varied regularly and in step with the glacial and interglacial periods (there were 10 during that time). At the ends of the glacial periods, atmospheric CO₂ content rose along with temperature at a rate of perhaps 10 to 20 ppm per 1,000 years. The current rate of increase is 100 times faster than that, and all the natural biogeochemical processes are now operating beyond the long-term bounds of atmosphere composition. Will global biogeochemical processes operate in the same way in this new world as they did in the past? We do not know.

Another uncertainty is how climate change will affect carbon uptake by terrestrial plant ecosystems. In general, elevated CO₂ should enhance many plants’ rate of growth. However, growth rate is also sensitive to moisture, availability of nutrients, and other factors, any of which may be the ultimate limit on plant growth. To illustrate, for certain perennial grassland species, the CO₂-enhanced growth rate is transient only because growth rate is eventually limited by the availability of nitrogen in the soil.²² Also, although climbing temperature causes plants to increase their rates of CO₂ respiration dramatically, this phenomenon also may be transient because plants tend to acclimate to the warmth and eventually thus to reduce respiration.²³

A related question is what happens to the carbon in soils. Warming and increased CO₂ content should increase the rate at which plant matter is added to soils, but it should also lead to an increase in soil microbial respiration and a consequent reduction of carbon content in the soil.²⁴ Over decades to centuries, such effects are potentially important as far as atmospheric CO₂ is concerned because of the far greater abundance of carbon in soils than in the atmosphere. The carbon content of and the rates of CO₂ and methane generation from soil depend on temperature, atmospheric CO₂ content, local rainfall, soil moisture, soil oxygen content, fertility, the plant species present, the time since a fire, the nature of the rock substrate, the extent of physical and chemical protection, and the enzymes and inhibitors present. Needless to say, soils vary greatly in terms of this long list of characteristics, which is one reason they are poorly understood.

Yet another question is how climate change will affect the rate at which the ocean removes CO₂ from the atmosphere. In addition to the uncertainties associated with ocean acidification, climate models suggest that increased warming should increase ocean thermal stratification. In other words, the ocean should become more stable, and downwelling should slow, which in turn should reduce the ocean’s ability to remove CO₂ by means of the solubility pump. However, the general expectation is that warming will cause the biological pumps to become more active, thus off setting the reduced activity of the solubility pump.

Although uncertainty persists in how the global carbon cycle will operate in a warmer, more CO₂-rich world, the natural processes of that cycle are clearly incapable of absorbing all 36 gigatons of anthropogenic CO₂ now being injected into the atmosphere every year.