16.6: Climate Change, Acidification, and Deoxygenation

The global radiation budget is not currently in equilibrium. The imbalance between heat energy entering and leaving Earth’s atmosphere is difficult to measure because it varies with the sun’s energy output, which varies in short term to decadal and longer cycles time scales. However, it has been estimated that in 2025 Earth radiated to space about 1.2 W/m² less energy than it receives from the Sun, creating an imbalance of 0.35% of the total incoming solar radiation at the top of the atmosphere. This imbalance is estimated to be growing at a rate of about 0.045 W/m² per year. In response to this radiation imbalance, additional energy is accumulating in Earth’s climate system and Earth is warming. Satellite sensors and instrumented buoys and floats at sea show an accelerating increase in heat content of the upper 2 km of the ocean since the 1970s.

Because Earth’s climate system has considerable thermal inertia, which is mostly due to the ocean’s high heat capacity, a considerable length of time is required for Earth’s climate system to achieve equilibrium under new environmental conditions. Unless there are unforeseen changes (such as a major increase in global volcanic activity) global climate models predict that rising concentrations of atmospheric carbon dioxide and other greenhouse gases will cause global warming to persist through this century and beyond. Even if greenhouse gas emissions were to stabilize at present levels, global warming would continue well beyond the 21st century. The magnitude of warming will depend on the rate of continued greenhouse gas emissions and related feedbacks.

Climate change is geographically non-uniform in both magnitude and sign so the rise in global mean annual temperature predicted by global climate models does not represent what might happen everywhere. Climate models predict polar temperature amplification, meaning that warming will be greater at higher latitudes.

The residence time of carbon dioxide in the atmosphere is between about 300 and 1000 years. The slow rate of removal of carbon dioxide from the atmosphere means that, if the rate of carbon dioxide emissions were to stabilize at the current level, the atmospheric concentration of carbon dioxide would still continue to increase. If anthropogenic emissions can be substantially reduced but not eliminated entirely, carbon dioxide would still enter the atmosphere faster than it is cycled out, and the atmospheric carbon dioxide concentration would continue to increase, although at a slower rate. Only the complete elimination of anthropogenic carbon dioxide emissions would stabilize and then, eventually, reduce the concentration of atmospheric carbon dioxide from current levels.

Approximately 40% of the estimated anthropogenic carbon dioxide released into the atmosphere since the beginning of the Industrial Revolution has been absorbed by the oceans. Currently, about 15% to 30% of anthropogenic carbon dioxide gas released to the atmosphere annually is removed from the atmosphere to the oceans. Some of the anthropogenic carbon dioxide absorbed by the oceans is transported to the deep oceans below the permanent thermocline by sinking of cold surface water in the global thermohaline circulation. Some carbon dioxide is also transported to the deep waters by sinking of organic matter produced by photosynthesis in the surface layer. The residence time of ocean deep waters is long, so this carbon dioxide is sequestered for as much as a thousand years before returning to the atmosphere-ocean interface.

Climate change, driven by rising concentrations of atmospheric carbon dioxide and other greenhouse gases, causes ocean surface waters to warm. Warming increases ocean density stratification and slows thermohaline circulation. Slowing the thermohaline circulation causes deep ocean water to spend longer in the deep oceans, where its oxygen concentration declines steadily due to the transport of organic matter from above and the respiration of decomposers. The amount of organic matter transported to the lower layers of the oceans is increasing due to enhanced photosynthetic primary production in the surface waters, which is due to higher concentrations of carbon dioxide and nutrients introduced by humans. The increased primary production contributes to declining oxygen concentration or deoxygenation in deep oceans. Absorption of anthropogenic carbon dioxide in ocean water also causes the water to acidify.

Thus, the oceans are subjected to acidification, deoxygenation, and circulation changes all caused by the same anthropogenic activity - carbon dioxide release to the atmosphere. Each of these can cause adverse effects on ocean ecosystems, and the synergistic effects of their acting together can be severe. Potentially so severe that this combination of rapid change in atmospheric greenhouse gases, ocean acidification, and deoxygenation of the deep oceans has been associated in the historical record with most, if not all, of the five major mass species extinctions that have occurred in Earth’s 4 billion year history. The possibility that our anthropogenic releases of carbon dioxide could be driving the ocean ecosystem toward another mass extinction is both real and serious. The consequences of a 6th mass species extinction to humans and Earth’s ecology would be catastrophic beyond anything that could possibly be caused by any of the other ocean pollution issues discussed in this chapter. Consequently, we discuss these three interrelated issues together in this section.

The Oceans and Climate Change

The ocean is the largest heat sink in Earth’s climate system and has been found to contain more than 90% of the additional heat energy captured by Earth due to anthropogenic greenhouse gases. About one-third of this heat energy is now in the deep ocean (below 700 m). Since 2005, the rate of total anthropogenic heat accumulation in the deep ocean has exceeded the rate in the upper layer (0 to 700 m).

Models predict that the ocean will continue to warm during the 21st century, with the strongest ocean warming projected to occur in the surface layer in tropical and Northern Hemisphere subtropical regions. At greater depth, the warming will be most pronounced in the Southern Ocean.

Temperature and Salinity Changes in the Oceans

Atmospheric global warming has accelerated in recent years, with the top 10 warmest years occurring since 2015. Since 1979, the rate of warming has been about 0.20±0.02°C (0.36±0.04°F), resulting in approximately 0.9°C (1.6°F) of warming over the last 45 years. The atmospheric warming has been accompanied by the warming of ocean waters

A warming atmosphere and ocean surface lead to a redistribution of rainfall and ocean salinity patterns. Salinity increases in regions where there is increased evaporation of ocean water. Conversely, increased inputs of fresh water from rivers, precipitation, and melting ice decrease salinity. Observations show that mid-latitude surface ocean water has become saltier, indicating the evaporation rate has increased in this zone. Surface waters in tropical and polar regions have become less salty. In other words, the moisture surpluses from the mid-latitudes condensed into liquid that precipitated in nearly equal amounts in these tropical and polar regions. This pattern is expected to lead to weakening of the meridional overturning circulation, impacting the climate of northern latitudes, contributing to increasing deoxygenation of the deep water layers of the ocean, as discussed in more detail later in this section.

Polar Regions

Polar temperature amplification means that the current global warming trend is greater at higher latitudes. Amplification of warming at higher latitudes threatens the ice sheets of Antarctica and Greenland. About 90% of Earth’s glacial ice blankets Antarctica, and melting this glacial ice would cause a catastrophic rise in sea level.

Two ice sheets cover most of Antarctica, separated by the Transantarctic Mountains. The larger of the two, the East Antarctic Ice Sheet (EAIS), is situated on a continent about the size of Australia, averages about 2 km in thickness, and accounts for two-thirds of Antarctic ice. Complete melting of the EAIS would raise mean sea level by about 60 m, although such large-scale melting is highly unlikely to occur under the climate predictions that are deemed likely during the next century or so, which depend on assumptions of future global carbon dioxide emissions. Geological evidence suggests that the EAIS has been stable for the past 30 million years and remains fairly stable today. However, it has undergone episodes of rapid disintegration and may have completely melted at least once in the past 600,000 years. The West Antarctic Ice Sheet (WAIS) sits on a series of islands and the floor of the Southern Ocean, with parts of the ice sheet more than 1.7 km below mean sea level. Complete melting of the WAIS would raise mean sea level by an estimated 5.8 m.

Researchers have concluded that the annual temperature of West Antarctica increased by about 0.20±0.18°C (0.35±0.32°F) per decade between 1957 and 2024: one of the fastest-warming regions on the globe. This warming has been accompanied by higher SSTs around Antarctica and the breakup of ice shelves along the Antarctic Peninsula coast. The majority of glaciers of the Antarctic Peninsula are retreating at an accelerating rate. In fact, the melting rate for the WAIS has tripled over the last 10 to 15 years. Scientists find that the section of the WAIS with accelerated melting appears to be in persistent decline, which will lead the glaciers in this area to melt into the ocean and cause global sea level rise. The question is whether that occurs sooner or later (that is, centuries or millennia).

Unlike the Antarctic ice sheets, which are polar (cold), the Greenland Ice Sheet is temperate. While the Antarctic ice sheets are well below the pressure-melting point (temperature at which ice melts at a given pressure), the temperature of much of the ice on Greenland is near the pressure-melting point. The melting point of ice decreases with increasing confining pressure (or depth within the glacier). Temperate ice sheets more readily generate meltwater and, with less frictional resistance, tend to flow downhill faster than polar ice sheets. The Greenland Ice Sheet has exhibited accelerated melting, and indications are that this melting rate might accelerate even further. Complete melting of the Greenland Ice Sheet would raise mean sea level by an estimated 7.3 m. The consequences of sea level rise are one of the most important adverse effects of climate change on human civilization.

Another major concern associated with the current global warming trend is shrinkage of Arctic sea-ice cover. Arctic sea-ice cover is shrinking at an accelerated rate so that the Arctic Ocean may be free of summer ice within a few years or decades. Although melting of floating sea ice does not raise sea level, it can alter climate significantly. Less sea-ice cover on the Arctic Ocean is causing the humidity of the overlying air to increase, leading to more cloudiness. Clouds cause both cooling (by reflecting sunlight to space) and warming (by absorbing outgoing infrared radiation and emitting infrared radiation in all directions, including some directed towards Earth’s surface). During the long, dark polar winter, additional cloud cover is expected to have a net warming effect. In summer, the impact of greater cloud cover depends on the altitude of the clouds. Cooling would be expected with an increase in low-altitude cloud cover, whereas warming would likely accompany an increase in high cloud cover. The mean annual air temperatures in the Arctic region have increased at a rate of 0.67±0.08°C (1.2±0.15°F) per decade, more than triple the global temperature increase rate.

In addition to shrinking sea ice, Northern Hemisphere snow cover has decreased, mountain glaciers are shrinking, permafrost is beginning to thaw, and fresh water runoff into the ocean has increased. Input of more fresh water from rivers and melting glaciers could impact the ocean thermohaline circulation by reducing the salinity and thus density of surface water in the area where dense, cold, salty water sinks at high latitudes of the Atlantic Ocean to become a critically important part of the MOC. Reductions in the strength of MOC could impact Northern Hemisphere weather patterns and would increase ocean deoxygenation.

Global Mean Sea Level

During the last glacial maximum, about 20,000 to 19,000 years ago, mean sea level was estimated to be 113 to 135 m lower than it is today. Sea level is currently close to the highest level it has been during the past 100,000 years, but it has been substantially higher, as much as several hundred meters higher, during warmer periods prior to that. Sea level falls during ice ages when global temperatures cool, and much of the world’s water is stored in glacial ice on land. In contrast, sea level rises when the climate and ocean water warms and glacial ice melts.

The mean rate of globally-averaged sea-level rise between 1901 and 2010 was about 1.7 mm• yr^-1, which increased to 3 mm•yr^-1 between 1993 and 2020. Considering the inertia in the Earth’s climate system, it is a virtual certainty that global mean sea-level rise will continue beyond 2100 and that sea-level rise due to thermal expansion will continue for many centuries. Due to the rapid melt now occurring in Greenland and Antarctica and the instability of parts of the Antarctic ice sheet, sea-level rise by the end of this century is expected to range from 0.3 to 1 m, depending on future anthropogenic release rates of greenhouse gases.

The impacts of sea level rise are extensive. Higher mean sea level is expected to accelerate coastal erosion by wave action, allow ocean water to inundate wetlands, estuaries, and islands, and make low-lying coastal plains more vulnerable to storm surges. Rising sea level is also expected to disrupt coastal ecosystems, ruin agricultural lands, threaten historical, cultural, and recreational resources, and displace coastal populations. Ports, airports, and other coastal infrastructure are likely also to be disrupted by rising sea levels, causing large economic losses and extensive costs to either protect them from flooding or relocate them. In some coastal areas, higher sea level is also likely to exacerbate saltwater intrusion into groundwater.

According to a report by the U.S. Office of Science and Technology, a 0.5 m rise in sea level, close to the lower end of the range expected before the end of the century, would result in a substantial loss of coastal land, especially along the U.S. Gulf and southern Atlantic coasts. Particularly vulnerable is South Florida, where the elevation of one-third of the Everglades is less than 0.3 m above sea level. For people currently living on low-lying islands (such as the Maldives and Tuvalu), abandonment appears to be their only option as sea level rises. Globally, 147 to 216 million people live on land that will be below sea level or regular flood levels by the end of this century.

Marine methane hydrates (Chapter 2) are not stable at surface temperatures and pressures, but form in pore spaces within ocean floor sediments at depths greater than 400 m, where temperatures are sufficiently low and/or pressures are sufficiently high. When methane hydrates are brought to the surface, the reduction in pressure and rise in temperature cause the hydrates to become unstable and decompose, releasing methane. Warming of ocean water may release methane from methane hydrates and, if large quantities of methane are released and not absorbed by marine organisms before reaching the atmosphere, may contribute to global climate change. Atmospheric methane is a greenhouse gas, more potent than carbon dioxide, although with a shorter lifetime in the atmosphere. If the released methane is consumed by marine bacteria that use it as a food source, this will consume dissolved oxygen and contribute to deoxygenation. Methane concentrations in the atmosphere have been monitored since about 1985 and have indeed been rising, with a faster rate since 2007 than in previous decades.

Marine Ecosystems

Marine organisms are vulnerable to a changing climate. Materials and energy flow from one organism to another via food webs within and between ecosystems. Climate change alters the physical and chemical conditions in the ocean, possibly exceeding the tolerance limits of many organisms. If these organisms are unable to avoid or adapt to these changed conditions, they may individually perish or, as a species, become extinct.

One important example of the vulnerability of marine organisms is the phenomenon of coral bleaching, which is not limited to corals. In addition to hard corals (Figs. 13-12) and sea fans (Fig. 14-8e), many mollusks such as giant clams (Fig. 14-35d), some flatworms, many anemones, and some protists all host symbiotic zooxanthellae. In response to elevated sea surface water temperatures, zooxanthellae become toxic, and their coral polyps expel them. The specific temperature susceptibility varies with the particular strain of the zooxanthellae hosted by each species. Without the colorful zooxanthellae, the corals appear nearly transparent against the white exoskeleton, so the condition is known as coral bleaching. The same “bleaching” process also occurs in anemones, clams, and other zooxanthellae hosting species.

For reef-building corals, a rise in SST of 1 to 2°C is sufficient to stress the corals, causing temporary bleaching. While corals can recover from bleaching, persistent or severe episodes cause coral polyps and the reef to die. Furthermore, higher ocean temperatures and bleaching make corals more vulnerable to infection and likely acidification, adding to reef mortality. In many tropical coral reef ecosystems, sea surface temperature has been high enough to cause coral bleaching during some El Niño events, but temperatures have been low enough between events to allow the corals to recover. However, as sea surface temperatures have continued to rise, temperatures between El Niño events are becoming progressively higher, and bleaching has occurred over greater areas every year. As a result, coral bleaching events are becoming more frequent and are now occurring annually. Because the corals no longer have sufficient time to recover between bleaching events, these recurring events often result in the collapse and death of much of the reef. Many reefs have already been observed to collapse and die from coral bleaching.

A key consideration in the potential impacts of climate change on marine organisms is the rate of change. Marine ecosystems can more readily adjust to gradual rather than abrupt changes in the ocean environment brought on by global-scale climate fluctuations. The current rate of climate change due to the enhanced greenhouse effect is approximately 10 times faster than at any time during the past 10,000 years.

Acidification

Recall from Chapter 1, that as a result of anthropogenically released carbon dioxide, the ocean is becoming more acidic. Evidence shows that at times the ancient ocean was more acidic than today, but modern acidification is happening faster than in the past. Currently, the ocean is still weakly alkaline, but the average pH value has dropped from 8.2 to 8.1 during the past 200 years. This corresponds to a 30% increase in acidity. By 2100, the ocean is expected to be 100 to 150% more acidic than today. This rate of acidification is expected to be too fast to allow many marine species to adapt or evolve. As a result, marine organisms and ecosystems may be adversely affected and many marine species may become extinct.

A more acidic ocean makes calcium carbonate, the material of many hard parts (shells and skeletons) of marine organisms, more soluble and therefore more difficult for marine organisms to grow and maintain. The adverse effects of ocean acidification on marine ecosystems are already significant. In those areas of the ocean where acidity levels are highest, bivalves, marine snails, and other mollusks show decreases in the growth rate of their shells, especially in the early larval stages. Echinoderms, most notably sea urchins, have shown a decrease in fertilization success, developmental rates, and larval size. For vertebrates like fish, the direct impacts are less understood, but recent studies indicate that fish, like invertebrates, are most susceptible at the egg and larval life stages. Research has documented changes in fish behavior and issues with the development of their otoliths, a calcareous structure of the inner ear linked to balance, orientation, and sound detection. These effects, especially when taken together, can disturb food webs and the entire marine ecosystem.

The observed adverse effects of acidification on oysters and pteropods in both the Bering Sea and the California Current off the west coast of the U.S. and Canada are described in Chapter 1.

Deoxygenation

Ocean deoxygenation is the loss of dissolved oxygen from the ocean. Multiple sources of observations show that during the past century, oxygen concentration in the ocean has declined. According to research published in 2017, scientists detected a decline of more than 2% in global ocean oxygen content between 1960 and 2010, while some areas experienced a larger drop. For example, the largest volume of oxygen loss occurred in the North Pacific, but the largest percent of decline occurred in the Arctic. A 2% drop may not seem significant, but, similar to the global mean annual temperature and ocean water acidity, small changes have large implications. Decreasing oxygen levels in the ocean pose a huge risk to marine ecosystems and, consequently, to humans.

The solubility of gases in water decreases as temperature increases, so warmer water holds less dissolved gas, including oxygen. Approximately 15% of the observed decline in global ocean oxygen content is attributed to warmer ocean temperatures. In addition, as surface ocean water warms, it becomes less dense and less likely to sink and mix with colder, less dense deep water. The overall result is a reduction in physical mixing, which translates to a longer residence time and a decrease in oxygen content of the deep ocean.

Recall from Chapter 1 (Figure 1-3) that the photic zone is the top layer of the ocean where sunlight penetrates. Here, oxygen is produced much faster by photosynthesis than it is consumed by animal respiration and the concentration of dissolved oxygen in the upper few meters of ocean water is saturated for that temperature. Photosynthesis declines rapidly with depth, but respiration continues throughout the ocean, so below the photic zone, the concentration of oxygen decreases continuously as the ocean water flows through the subsurface oceans. Just below the pycnocline is where the water mass has been out of contact with the atmosphere for the longest period of time and where the rate of respiration is high due to the rapid consumption of detritus falling from the photic zone. At this depth, there is an oxygen concentration minimum coinciding with the nutrient maximum.

Figure 1-3. Oxygen concentration in ocean water is controlled by both physical and biological processes. The oceans consist of two main layers, a warm surface layer and a cold deep layer, with a transition zone, called the pycnocline zone, between them. (a) Surface layer water is well mixed and in contact with the atmosphere, and its oxygen concentration is primarily controlled by gas exchange with the atmosphere. (b) Photosynthesis by marine life in this layer converts carbon dioxide to organic matter and releases oxygen, but this is offset by respiration, which converts oxygen to carbon dioxide. Respiration also occurs in the deep layer, fueled by animals that migrate vertically or live in the deep and decomposition of organic material (detritus), but there is no photosynthesis and no contact with the atmosphere, so respiration continuously consumes oxygen in this deep layer, and, therefore, the longer the water remains in the deep layers, the less oxygen it contains. (c) The deep layer water is continuously replenished by sinking cold water in some near-polar regions and then circulates through the oceans while being slowly mixed upward, eventually into the surface layer again. The pycnocline provides a barrier that inhibits mixing, so the upward mixing process is slow.

Currently, large areas of the tropical and subtropical oceans exhibit a strong oxygen minimum zone below the pycnocline (Figure 16-11). In these areas, the oxygen minimum zone water has oxygen concentrations below 70 μmol•kg^-1, below which marine species (other than those microbial species that are anaerobic) are subject to oxygen deficit stress, so the water is considered to be hypoxic. In smaller areas, where the concentration is essentially zero (anoxic), aerobic marine species can not survive. The volume of both hypoxic and anoxic water has been observed to have expanded during the past five decades. The hypoxic zone has expanded by an area that is more than half the area of the continental U.S., while the area of the anoxic zone has quadrupled. If this trend continues, it will become increasingly more likely that the low oxygen zone will mix upward to affect the upper water layers where there are higher concentrations of marine animal habitats. Recall that for the past few years, hypoxic (concentration of dissolved oxygen low enough to be detrimental to organisms), and at times anoxic (totally depleted of dissolved oxygen), water from this water mass has seasonally moved or been mixed onto the Oregon continental shelf, forming a dead zone (Chapter 13).

Figure 16-11. The concentration of dissolved oxygen at 300 m depth. This depth is generally below the pycnocline and at approximately the depth where an oxygen minimum occurs. Concentrations in the figure are expressed as μmol•kg^-1. The water is considered to be hypoxic at concentrations below 70 μmol•kg^-1.

While the cause of the expansion of oxygen minimum zones is not yet fully known, it has been estimated that about 15% of the total oxygen loss in the oceans is due to the warming-induced reduction in the solubility of oxygen in ocean water. However, warming is responsible for about 50% of the loss in the upper 100 m of the oceans. The remaining 85% of total oxygen loss in the ocean deep waters is likely due to increased stratification, which slows the sinking of oxygen-rich surface water, combined with faster transport of carbon-rich particles to the deep water as primary production increases as a result of nutrient enrichment and warming of the ocean surface layers. It is expected that oxygen in the deeper layers will decline further during the next several hundred years as the deep water is replaced by the meridional overturning circulation, and that this oxygen loss will occur even if there is no continued increase in stratification or nutrient enhancement of primary production. This means that the areas of oxygen minimum zones, in which the oxygen concentration is below the concentration needed to support aerobic marine species, will continue to grow for centuries, even if humans were to reduce anthropogenic carbon dioxide and nutrient emissions to zero or near zero immediately.

Recent data and modeling suggest that the post-industrial loss of oxygen concentration now observed in deep ocean waters represents only about 25% of the loss that will occur over the next several hundred years, even if anthropogenic emissions were to be completely stopped immediately. Unless anthropogenic emissions are reduced sufficiently and quickly enough to stop additional warming and further nutrient enrichment of the oceans, the rate at which the oxygen minimum zones will expand will be higher, likely substantially higher, than the now unavoidable rate of increase that will result from the existing warming and solubility loss in surface waters. We are already committed to much expanded deep ocean hypoxic and anoxic zones in the deep oceans with consequent but largely unknown effects on ocean biology and long-term climate. These already committed changes will not be reversed for as much as a thousand years or more even if we immediately stop all anthropogenic greenhouse gas emissions and nutrient releases.

Hypoxia and anoxia have also grown in coastal and estuarine regions during the past 50 years. During this time period, over 500 locations in coastal waters worldwide have been known to experience hypoxia. Fewer than 10% of these ocean areas are known to have experienced any hypoxia before the mid-20th century.

Hypoxic and anoxic water is currently found in only relatively limited areas of the ocean, mostly in the oxygen minimum zones below the pycnocline in the open ocean, and in an increasing number of coastal regions and estuaries where the hypoxia or anoxia is most often caused by inputs of nutrients from sewage and agricultural runoff. However, geologic evidence shows that periods of widespread anoxia within the ocean have occurred throughout Earth’s history. Uncertainties remain about the causes of these anoxic events, but they were usually associated with warmer climate conditions, rises in sea level, and occasionally with mass extinctions. For example, global ocean anoxia has been proposed as the main cause for the mass extinction event at the close of the Permian (245 ma).

Climate Change Summary

In summary, while climate change will cause sea-level rise and changes in nutrient supply and ocean circulation that may adversely affect some ocean ecosystems, the more serious global threat to marine life from climate change comes from the combined effects of rising ocean temperature, increased acidity, and decreased oxygen concentration. Given the anticipated range of changes to each of these parameters, many marine species will be stressed or killed, and at least some species will become extinct within the next several decades. Indeed, climate-driven changes in each of these parameters have already been observed to cause harm to marine ecosystems in some parts of the ocean. For example, as discussed above, warming ocean water has caused a rise in the frequency of coral bleaching and damaged coral reef ecosystems. Increased acidity has caused mortality of oyster larvae and reduced reproductive success of pteropods in some parts of the northern ocean. Decreased oxygen concentration has caused episodic dead zones in many parts of the coastal and estuarine ocean and in the deep layers of some ocean areas.

As stated above, deoxygenation has been implicated as a cause of at least one of five mass extinctions in Earth’s past and probably at least contributed to the other four mass extinctions. Warming and acidification are thought to have also played a part in these five mass extinctions. Elevated carbon dioxide concentrations in the atmosphere have also been associated with most extinctions. Some scientists have suggested that increased carbon dioxide in the atmosphere drives warming, acidification of the ocean, and deoxygenation of the ocean, and it is the combined effects of these stresses that were responsible for most of Earth’s mass extinctions in the past. Moreover, the rate of species extinction has increased in the past few centuries, leading some scientists to hypothesize that Earth may already be experiencing a sixth mass extinction of species.

Much more detailed reliable scientific information on the likely future impacts of climate change, acidification and deoxygenation can be obtained from reports of the Intergovernmental Panel on Climate Change (IPCC).