|
||||||||||||||||
|
||||||||||||||||
The following HTML text is provided to enhance online readability. Many aspects of typography translate only awkwardly to HTML. Please use the page image as the authoritative form to ensure accuracy. 3
|
||||||||||||||||
 |
 |
Page 71 |
 |
 |
 |
|
| ||||||||||||||||||||||||||||||||
|
|
|||||||||||||||||||||||||||||||
Below are the first 10 and last 10 pages of uncorrected machine-read text (when available) of this chapter, followed by the top 30 algorithmically extracted key phrases from the chapter as a whole.
Intended to provide our own search engines and external engines with highly rich, chapter-representative searchable text on the opening pages of each chapter.
Because it is UNCORRECTED material, please consider the following text as a useful but insufficient proxy for the authoritative book pages.
Do not use for reproduction, copying, pasting, or reading; exclusively for search engines.
OCR for page 71
Origin and Evolution of Earth: Research Questions for a Changing Planet
3
A Habitable Planet
Earth’s hospitable climate—with temperatures high enough to keep surface water in the liquid state but low enough to keep too much water vapor from entering the atmosphere—is a special and probably critical feature of the planet. There is growing public awareness that climate can change, and there is abundant evidence in the geological record that climate has changed in the past. The history of Earth’s climate, a peculiar combination of both variability and stability, poses challenging scientific questions. Our current understanding suggests that many factors can change climate, some capable of producing rapid changes and some requiring much more time but also potentially causing much larger changes. However, despite the many ways that natural forces can change Earth’s climate, substantial geological evidence suggests that Earth’s overall climate, although it has oscillated between relatively warm and relatively cold states many times, has somehow been maintained in a reasonable, and quite narrow, range that is conducive to the preservation of life. The equitable climate conditions have been present for perhaps 3.5 billion years, despite the fact that both the Sun and Earth have changed in ways that might be expected to play havoc with Earth’s climate.
This chapter addresses major questions that relate to understanding how Earth’s surface conditions can change and at the same time can be maintained between limits that are conducive to life over extremely long times. Question 7 is concerned with the geological and astronomical factors that affect climate and the geological evidence of climate change. Question 8 considers the relationship between geology, climate, and life. The picture that emerges from Question 7 is that a large number of factors contribute to governing Earth’s climate, but how the interplay of these factors results in a particular climate state is still unknown. The answer to this question is critical for addressing future climate change. Question 8 raises the interesting possibility that life itself helps govern climate and other aspects of Earth’s surface conditions, while at the same time we have conclusive evidence that climate change has at times been seriously detrimental to life, occasionally killing off huge numbers of species and often forcing evolutionary change.
QUESTION 7:
WHAT CAUSES CLIMATE TO CHANGE—AND HOW MUCH CAN IT CHANGE?
Among the systems of planet Earth addressed in this report, climate is the most widely discussed in public forums. We know that human civilizations developed during an unusual period of climate stability over the past 10,000 years or so, but we also know from geological evidence that momentous changes can occur in periods as brief as centuries or even decades in ways that would disrupt human settlement patterns worldwide. Moreover, it is widely recognized that Earth’s mean global surface temperature has risen since the beginning of the industrial age and that emissions of CO2 and other greenhouse gases are at least partly, if not wholly, responsible (IPCC, 2007a). The potentially serious consequences of this global warming, ranging
OCR for page 72
Origin and Evolution of Earth: Research Questions for a Changing Planet
from inundation of densely populated coasts to ocean acidification to the poleward spread of tropical diseases, underscore the need to determine how much of the warming is caused by human activities and what can be done about it. Earth science has an important role in answering both questions.
The immediate grand challenge in climate science is predicting how climate will change over the coming decades. However, the broader challenge is to account for both the long-term consistency of Earth’s climate and its multiple and varied excursions in the context of a constantly evolving global geological and biological framework. Only when we are able to capture past climate changes in models will we have confidence in our predictions of future climate. Reliable models have not been available because the conditions that characterized ancient climates—such as ground surface temperature, sea surface temperature, and mean annual precipitation—vanished thousands or millions of years ago, along with the climate they shaped. Lacking real-time data for ancient events, geologists are assembling toolkits of “proxy” data. The temperature and precipitation of continental regions, for example, can often be inferred from evidence left in the sediments of lake beds or in ancient preserved soils. Earth’s large-scale surface temperature structure, as well as information on ancient ocean currents, is also reflected in fossil and geochemical records of deep-sea sediments and in records of sea-level change. Similarly, atmospheric temperatures for at least the past 100,000 years or so are recorded in glacial ice and retrievable through deep drill cores in the ice. However, the further we journey into Earth’s past, the more different Earth was from our modern planet. To understand Earth’s climate in geologically ancient times, we need to know an enormous amount about the geology and geography of the ancient Earth; this is where geological science and climate science become inseparable.
What Processes Govern Climate Change?
The climate system is regulated by how much energy Earth receives from the Sun and how much is radiated back into space (Figure 3.1). How much energy is absorbed depends on the reflectivity (or albedo) of Earth’s atmosphere and surface. The albedo depends on how much of Earth’s surface is covered by water, land, or ice; how the continents are arranged; the extent of land vegetation; and the amount of reflective material (clouds and particles) in the atmosphere. It is generally believed that the key determinant of Earth’s ability to capture energy from the Sun is the amount of greenhouse gases, predominantly carbon dioxide, present in Earth’s atmosphere. Increasing the CO2 content of the atmosphere stimulates warming, which is then amplified by increasing amounts of water vapor that can evaporate from the oceans at higher temperature. Hence the cornerstone of any broader understanding of Earth’s climate is the question of what controls the amount of CO2 in the atmosphere.
The various processes that contribute to the CO2 content of the atmosphere are referred to collectively as the carbon cycle. The carbon cycle is a key regulator of climate change. The overarching issue is the fraction of Earth’s carbon that is present in the atmosphere in the form of CO2 or other greenhouse gases like CH4. For the modern Earth, most of the carbon is stored in rock, and most of that is stored deep within the mantle and core. Estimates suggest there is 500,000 times more carbon stored in Earth’s mantle than in the atmosphere (McDonough and Sun, 1995; Salters and Stracke, 2004), and there is likely to be more carbon in Earth’s core than in the mantle. Most of the carbon not stored in the mantle and core is found in sedimentary rocks as the mineral calcite or as organic material like kerogen and petroleum. Most of the rest is either dissolved in the oceans, stored in soils, or present as living plant and animal tissue. Only a very tiny fraction (roughly one-millionth) is present in the atmosphere and acting to help warm Earth’s surface. The Venusian atmosphere, which contains about 200,000 times more CO2 than Earth’s preindustrial atmosphere, is clear evidence that the distribution of carbon between a planet’s interior and atmosphere can be very different from that of Earth.
Even though the amount of carbon in Earth’s atmosphere is small, changes in the amount have a major effect on the surface temperature. Although the carbon in Earth’s core is not likely to be transferred to the atmosphere, there are ways that at least some fraction of the enormous amounts of carbon in Earth’s mantle, crust, and oceans could be. Similarly, there are ways to transfer the carbon in the atmosphere to the oceans and to sediments and then to subduct them into the mantle.
OCR for page 73
Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 3.1 Schematic view of the global climate system, showing many of the flows of energy, water, and CO2 that control the climate and the processes that play a role in regulating Earth’s greenhouse and determining what happens to the solar energy. Not shown is the warm circulation near midocean ridges, which moves CO2 from the ocean to the shallow oceanic crust. SOURCE: After <http://www.carleton.edu/departments/geol/DaveSTELLA/climate/climate_modeling_1.htm>. Courtesy of David Bice, Pennsylvania State University. Used with permission.
Studies of the carbon cycle are aimed at understanding how the atmospheric carbon content is regulated by geological and biological processes. Over the past century, fossil fuel burning has overwhelmed natural processes, quickly transferring a large amount of buried carbon (in the form of organic matter, petroleum, coal, and natural gas in sedimentary rock formations) into the atmosphere as CO2. On longer timescales, natural processes (e.g., volcanism, subduction, chemical weathering, sedimentation, metamorphism, glaciation, wildfires) also shift carbon between the atmosphere, oceans, sedimentary formations, soils, plants, and deep interior. These processes produce cycles of increasing and decreasing atmospheric CO2 that occur over timescales of thousands, millions, and billions of years. At each timescale, different processes are primarily responsible for the changes.
Over the past century and through the next, changes in the greenhouse gas content of the atmosphere are the most important factor affecting climate, although changes in atmospheric particulates and clouds are also important. Burning coal, oil, and natural gas continues to add greenhouse gases and aerosols to the atmosphere, reducing emissions of infrared radiation to space and causing Earth’s global mean surface temperature to rise. The amount of increase depends on feedbacks in the climate system, especially the (poorly known) feedback from clouds. On even shorter timescales (years to decades), changes in atmospheric particle loading, notably sulfate aerosols, can affect climate, in
OCR for page 74
Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 3.2 Example of oxygen isotope data measured on carbonate shells of a single species of foraminifer separated from a 10-m core of deep-sea sediment. Glacial-interglacial cycles are evident. Higher δ18O values represent times when bottom water temperature was lower and the volume of continental glaciers was larger. Modern time (depth = 0, age = 0) corresponds to an “interglacial” period. Upper graph shows depths where age can be estimated and the estimated age. SOURCE: Data from SPECMAP, <http://www.ngdc.noaa.gov/mgg/geology/specmap.html>.
part countering the effect of increased CO2. The 1991 eruption of Mount Pinatubo, for example, caused slightly cooler-than-average global temperatures for about a year.1 Despite the uncertainties and feedbacks, a doubling of CO2 from fossil fuel burning is now predicted to increase the mean surface temperature 2°C to 4.5°C by about the middle of this century (IPCC, 2007a).
As the period of time under consideration lengthens, more diverse processes that can affect climate come into play. Over thousands of years, variations in Earth’s orbit around the Sun (Milankovitch forcing) affect how solar energy is distributed around the globe and lead to changes in mean annual temperature, precipitation, and seasonality. Earth’s orbital cycles are responsible in part for the oscillations between ice ages and interglacial periods that characterize the past 3 million years of Earth history (Figure 3.2). Over thousands of years the oceans are important as well; for example, excess CO2 in the atmosphere should dissolve into the oceans after about 1,000 years. And in glacial times the increased ice cover on Earth changes the albedo. If ice caps start to grow as a result of cooling over thousands of years, they can reflect more sunlight and enhance cooling.
The role of tectonic processes (volcanoes, mountain building, continental drift) becomes dominant at timescales of a million years or longer (Figure 3.3). Volcanoes, for example, tend to move CO2 from the deep Earth to the atmosphere, whereas erosion of mountain ranges and the associated chemical weathering of minerals tend to remove CO2 from the atmosphere and ocean and bury it as calcite and organic matter in sediments on the ocean floor. Plate motions, which rearrange the continents and oceans, affect atmospheric
1
<http://data.giss.nasa.gov/gistemp/2005/>.
OCR for page 75
Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 3.3 Global deep-sea oxygen and carbon isotope variations associated with major climatic, tectonic, and biotic events, based on data compiled from more than 40 ocean drilling holes. The temperature scale refers to the temperature of typical water near the ocean bottom and applies only to the time period from 70 million to 34 million years ago. Presently the bottom water temperatures are typically about 2°C and global mean surface temperature is 15°C. Fifty million years ago, bottom water temperatures were about 10°C to 12°C, which corresponds to a global mean surface temperature of about 25°C. From the early Oligocene to the present, about 70 percent of the variability in the δ18O record reflects changes in Antarctica and northern hemisphere ice volume. The vertical bars provide a rough qualitative representation of ice volume in each hemisphere relative to the Last Glacial Maximum, with the dashed bar representing periods of minimal ice coverage (≤ 50 percent), and the full bar representing close to maximum ice coverage (> 50 percent of present). SOURCE: Zachos et al. (2001). Reprinted with permission from the American Association for the Advancement of Science (AAAS).
and oceanic circulation, which in turn changes the efficiency of heat transport from low to high latitudes. These connections can be seen where geological events are correlated with major climate shifts. Volcanic activity that occurred as North America broke away from Europe and the large outpouring of lava that produced the Columbia River plateau about 15 million years ago are both associated in time with globally warm temperatures. The opening of the Tasmanian and Drake passages as continental drift separated Antarctica from neighboring continents is close in time to the first growth of continental glaciers on Antarctica. Factors on the Antarctic continental shelf, such as the elevation of the Vostok and Gamburtsev Mountain regions, may have played an important role in initiating glaciations as well. Continental drift, combined with volcanism, also closed the Panama Seaway, which once connected the Pacific and Atlantic oceans, drastically changing ocean
OCR for page 76
Origin and Evolution of Earth: Research Questions for a Changing Planet
circulation patterns and probably triggering glaciation in the northern hemisphere about 3 million years ago (Zachos et al., 2001).
Some also think that the continental collision of India with Asia, which formed the Himalayas, the Tibetan plateau, and related mountain ranges, has been a primary cause of Earth’s gradual cooling to glacial conditions over the past 50 million years (Raymo and Ruddiman, 1993). The growth of those massive mountain ranges is hypothesized to have accelerated erosion and weathering, yielding dissolved calcium that was carried to the oceans by rivers. This calcium was used by organisms to build shells of calcium carbonate (calcite or aragonite), some of which accumulated as sediment on the ocean floor. This well-known example of sequestering carbon by burying it on the seafloor has drawn broad interest among those searching for a practical way to reduce atmospheric CO2 (IPCC, 2005). Also, when sedimentary conditions and ocean chemical conditions are right, as has happened many times in the past 500 million years, large amounts of carbon can be held as organic matter within silicate and carbonate sediment on the ocean floor. It is this process that produced the rock formations that we now exploit for fossil fuel.
Why Has Climate Stayed in a Hospitable Range?
The luminosity of the Sun may be an important regulator of climate on timescales of billions of years. Stellar evolution models indicate that the Sun’s power output has increased by about 40 percent since it first became a star. The lower solar luminosity 4.5 billion years ago would correspond to an Earth surface temperature about 35°C lower than the present—well below the freezing point of water—if other conditions on the early Earth were similar to those of today (Kasting and Catling, 2003). And yet there is evidence from 3.8-billion-year-old rocks and more controversially from 4.4-billion-year-old zircons (Valley et al., 2002) that the earliest Earth had liquid water at its surface (see Question 2). How can this be possible?
How Earth has remained within the temperature limits for liquid water and life for over 4 billion years is a central question about our planet (Box 3.1). A feedback involving volcanism and weathering may provide a partial answer (Walker et al., 1981; Berner et
BOX 3.1
A Hospitable Climate
We know that during the past 4 billion years Earth’s climate has varied enough to contribute to the extinction of many species. And yet the variations have been mild enough that life has always rebounded quickly. So how hot is too hot, and how cold is too cold for humans?
Earth’s nearest planetary neighbors have both stronger and weaker greenhouse effects, with climates either too hot or too cold for life as we know it. A thick cloak of CO2 heats the surface of Venus to about 470°C, whereas the thin atmosphere of Mars keeps the mean annual surface temperature at about −56°C (Consolmagno and Schaefer, 1994). In comparison, Earth’s mean global surface temperature is currently about 15°C. In our present climate state, the mean annual temperature is about 27°C in the equatorial regions and below freezing (and perennially ice covered) at high latitudes. If we imagine an Earth with a global mean temperature just 10°C higher, the equatorial regions might have temperatures as high as 35°C (depending on how much the tropics widen due to water vapor feedback), unusually hot by human standards; there would be no permanent ice cover in polar regions, and most high-latitude precipitation would fall as rain rather than snow. If the global mean temperature were 10°C lower than today, Earth would be covered with ice to the midlatitudes, more extensively than in the ice ages of the past few million years.
The geological record suggests that climate has stayed within these extremes throughout Earth’s history, except for geologically brief “snowball Earth” episodes in the Precambrian. But even much smaller fluctuations in temperature can have a significant impact on human settlement. For example, the Medieval Warm Period (about AD 1000 to 1270) brought extensive drought that may have caused indigenous peoples to abandon the great cliff cities in the western United States (Herweijer et al., 2006); at the same time it made Greenland habitable to the Vikings until the Medieval glaciations of the early and mid-14th century (Barlow et al., 1997). Even with modern technologies, the coldest and warmest areas on Earth support only small populations.
al., 1983). According to this model, weathering slows as climate cools, allowing volcanic CO2 to accumulate in the atmosphere. The added CO2 warms the climate again, causing weathering to accelerate and prevent further warming. The same feedback loop may have allowed more CO2 to accumulate in the atmosphere early in Earth’s history, compensating for the lower solar luminosity and keeping temperatures above freezing. This stabilizing feedback mechanism would operate slowly and so would be effective only over millions of
OCR for page 77
Origin and Evolution of Earth: Research Questions for a Changing Planet
years; it would not significantly temper the effects of rapid CO2 increases over the next 100 years. In addition, there is still uncertainty about the effectiveness of this weathering-volcanism feedback because of the competing effect of water–crust interactions as a sink for CO2 and because of increasing evidence (discussed below) that weathering rates do not depend mainly on Earth’s surface temperature. If temperature is not the primary determinant of weathering rates, atmospheric CO2 could vary rapidly and the fluctuations may be even more difficult to predict because they would depend on global factors such as the rate of mountain building due to continental collisions.
Long-term climate regulation may also involve other processes and other greenhouse gases. During the first half of Earth’s history, when atmospheric O2 levels were low (Holland, 1984; Farquhar et al., 2000), reduced gases may have been more abundant in the atmosphere. Methane (CH4), for example, could have been present at concentrations of 1,000 ppmv or more, compared to only 1.6 ppmv today (Kharecha et al., 2005). At such high concentration, CH4 could have contributed 10°C to 20°C of greenhouse warming (Pavlov et al., 2003). Disappearance of much of this CH4, which must have happened when atmospheric O2 levels rose at about 2.4 Ga (billion years ago), could explain why Earth became glaciated at that time. This hypothesis is attractive, but it has not been tested directly with data from the geological record. Below we discuss what types of information are available from detailed sampling of this record.
What Caused Exceptionally Warm and Cold Periods in Geological Time?
The geological record of climate change, written in ice cores, sediments, fossils, and rocks, provides clues about how much climate has varied over the past 4 billion years (Box 3.2) and the future habitability of Earth. From this record geologists have been able to identify some of Earth’s more extreme climates and the factors that may have triggered them.
One of the warmest extended periods in the geological record occurred in the Cretaceous period, about 120 million to 90 million years ago (Barron and Washington, 1982), when large areas of the continents were flooded with shallow seas (Figure 3.4). At the end of that time, polar temperatures up to 14°C were high enough to support evergreen vegetation, dinosaurs, turtles, and crocodiles north of the Arctic circle (Tarduno et al., 1998). Equatorial temperatures were 3°C to 5°C warmer than today (Wilson and Norris, 2001), and the deep-ocean temperature may have reached 20°C (Huber et al., 2002) as compared to 0°C to 5°C today. Models and proxy studies suggest that the atmospheric CO2 concentration during the Cretaceous was 2 to 10 times higher than it is today (Caldeira and Rampino, 1991; Ekart et al., 1999; Haworth et al., 2005), although these estimates are still highly uncertain and we do not know how variable the CO2 concentration was on shorter timescales during the Cretaceous.
The causes of Cretaceous warming are still unknown. Volcanic activity and hence the input of CO2 to the atmosphere were probably unusually high, as suggested by the plethora of volcanic mountains and plateaus of that age on the western Pacific Ocean floor. The weathering that removes CO2 from the atmosphere may have been reduced by two processes: (1) the higher sea level would have reduced the continental area subject to weathering, and (2) this period lacked the major continental collision zones that make mountains, which weather more rapidly than flatter terrain. The paucity of sea ice would also have decreased albedo. The clustering of continents could have changed atmosphere and ocean circulation patterns, increasing the poleward transport of heat and thus making the polar regions warmer relative to the tropics. Whatever the primary causes, the middle Cretaceous is our best example of a greenhouse Earth. However, the geography and ocean circulation are so different today that a future greenhouse may look very different.
The coldest period we know of occurred in the Neoproterozoic. This period is particularly interesting for climate scientists. Conditions then were so drastically different from those today that they strain our understanding of how the climate system works. Between 750 million and 580 million years ago, Earth’s surface, including all of the oceans, may have frozen over completely for several brief intervals (Hoffman et al., 1998), creating a “snowball Earth.” This hypothesis is vigorously disputed (e.g., Hyde et al., 2000)—not the anomalous cold but its cause, duration, and severity. The cold was almost certainly triggered by transient lowering of greenhouse gas concentrations, and the
OCR for page 78
Origin and Evolution of Earth: Research Questions for a Changing Planet
BOX 3.2
How Do We Estimate Climate Variables in the Past?
Historical accounts of climate are available for only the past few hundred years, so information about older climate events must be gleaned from alternative archives, such as tree rings and isotopic compositions of ice cores and ocean sediments. These records provide an indirect (or proxy) measure of climate variables, such as temperature and CO2. Proxies tend to respond to more than one factor in the climate system, so multiple measures are needed to interpret them. The further back in time we go, the fewer kinds of proxy records are available, the more limited their spatial coverage, and the greater the uncertainty in what they mean. Thus, a major effort is being made to expand the collection of proxy observations in space and time and to develop new kinds of proxies (Henderson, 2002).
Proxy Measurements Used to Estimate Climate Variables
Variable
Age Range
Proxy Measurement
Mean temperature
Centuries
Glacier length
Ground surface temperature
Centuries
Borehole temperature measurements
Summer temperature
Few millennia
Tree rings, pollen analysis
Land temperature, precipitation
Millennia
Lake sediments (O isotopes)
Mean annual temperature, precipitation
Millennia
Speleothems (O isotopes)
Sea surface temperature
Millennia
Corals (O isotopes, Sr/Ca, and U/Ca)
Atmospheric temperature
Hundreds of thousands of years
Ice cores (O and H isotopes)
Sea surface temperature
Millions of years
Foraminifera (O isotopes, Mg/Ca)
Land or ocean temperature
Millennia to hundreds of millions of years
Fossils, evidence of ice, sedimentary structures (evidence of water)
CO2 and ocean pH
Tens of millions of years
Foraminifera (B isotopes, Ca isotopes)
CO2
Hundreds of millions of years
Soil carbonate (C isotopes), stomatal indices in plant leaves
FIGURE 3.4 Physiographic representation of North America, Europe, and North Africa 90 Ma when climate was warm and sea level was high. The land area of the continents was substantially smaller because oceans had risen above the edges of the continents and flooded the interiors. North America was still close to northern Europe, and the North Atlantic Ocean was barely connected to the other oceans. The South Atlantic Ocean (not shown) had not yet formed. SOURCE: <http://jan.ucc.nau.edu/~rcb7/090NAt.jpg>. Courtesy of Ron Blakey, Northern Arizona University. Used with permission.
OCR for page 79
Origin and Evolution of Earth: Research Questions for a Changing Planet
actual cause may have been the different locations of the continents. The continents were all situated at low to midlatitudes where temperatures are warmest, allowing silicate weathering to proceed rapidly and draw down CO2 levels, even as the global surface temperature dropped and polar ice accumulated (Marshall et al., 1988; Donnadieu et al., 2004). Alternatively, CH4 concentrations may have been high during the mid-Proterozoic and then dropped as O2 levels increased (for a second time; see Question 8) near the end of this time (Pavlov et al., 2003). In either case, as ice cover increased, the albedo and thus cooling would have increased until the planet plunged into an extreme “icehouse” condition. Surface temperatures calculated for this hard snowball Earth are about −20°C at the equator and about −40°C averaged ver the globe (Pollard and Kasting, 2004).
The existence of a snowball Earth must be inferred from geological evidence. Translation of such evidence into a hypothesis about Earth’s climate and evaluation of the hypothesis using modern climate models and concepts provide an interesting example of the scientific challenges inherent in reconstructing Earth’s past conditions. The rock assemblage now considered indicative of the snowball period was initially difficult to decipher. There are marine glacial deposits that formed near the equator, suggesting glaciation in the tropics and hence exceptionally cold conditions; banded iron formations, suggesting anoxic conditions in the oceans; and stratigraphically above and below the glacial deposits there are limestones, which suggest warm conditions (Figure 3.5; Hoffman and Schrag, 2000). In some cases there are nonmarine deposits, which suggest that sea level dropped, and there is carbon isotopic evidence suggesting that photosynthesis all but stopped.
The warm conditions following the snowball Earth period may have arisen because volcanism would have continued through the snowball period, contributing CO2 to the atmosphere that could not be removed by rock weathering because the rocks were covered with ice. Once extreme levels of CO2 were reached (~400 times the modern preindustrial level; Caldeira and Kasting, 1992), the greenhouse effect would have been strong enough to overcome the high albedo, melt the ice, and swing Earth to exceptionally warm conditions (~40°C global average in this model) before weathering processes could catch up and remove the atmospheric CO2. The temporarily high atmospheric CO2 would probably have made the rain especially acidic, enhancing chemical weathering and causing a large amount of calcium to be delivered to the oceans by rivers; this may explain the unusual, rapidly deposited limestone layers that cap most Neoproterozoic glacial deposits (Hoffman and Schrag, 2000). A recent three-dimensional climate simulation by Pierrehumbert (2004) has cast doubt on this scenario, however. The new calculations indicate that even 0.2 bars of CO2 (700 times the preindustrial level) could not have deglaciated a hard snowball Earth. Given the many uncertainties involved in applying climate models to the Proterozoic Earth, it is not yet clear whether the hypotheses or the models are incorrect.
Indeed, there are many arguments against the snowball Earth hypothesis. Even supporters of this theory disagree about significant issues. One is the survival of photosynthetic algae through the plunge in temperatures. How was this possible if the ice was a kilometer thick everywhere as some models have it? Could photosynthetic life have survived in local volcanic hot spots, like modern Iceland? Or did other refuges exist? One variant of the snowball hypothesis, the so-called thin-ice model (McKay, 2000), suggests that the ice in the tropics was only about 1 to 2 m thick, allowing enough penetration of sunlight for photosynthesis. In addition, there would likely be leads and lanes of open water in very thin ice. This model allows Earth to deglaciate at a much lower CO2 level, only about 30 times the present level (Pollard and Kasting, 2005). However, there are questions as to whether such a solution can be stable, given that sea ice can flow from the poles to the equator, where it would melt (Goodman and Pierrehumbert, 2003). Clearly, much more work is required if the snowball Earth hypothesis is to become an established chapter in Earth’s climate history. Nevertheless, even the most moderate of interpretations of the Neoproterozoic evidence for glaciation suggest that it was the coldest period in the past 2 billion years. By comparison, the glaciations that have affected Earth in more recent times have had comparatively little effect on the global carbon cycle.
What Triggers Abrupt Climate Change?
Abrupt climate events are unusual, but they provide insights on the rates at which the climate system is
OCR for page 80
Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 3.5 Example from Namibia of the rock record of extreme climate change in Earth’s past. These 750-million-year-old sedimentary rocks have been tilted by tectonic movements, but the time sequence is preserved, progressing from lower right to upper left. The Ombaatjie formation is a limestone deposit formed in shallow ocean water; near its top, isotopes indicate that a glaciation was starting, and above that level the rocks are wind-blown sand dunes, indicating that sea level dropped due to glaciation. Above the dune deposits are limestones deposited after the glaciation ended. The “crystal fans” are a rare type of limestone that is hypothesized to form when inorganic carbonate is rapidly precipitated from the oceans. The time duration represented by this rock sequence is not known, but estimates suggest a few million years. SOURCE: Halverson et al. (2002). Copyright 2002 American Geophysical Union. Reproduced with permission.
capable of change. Much like the extremes of warm and cold discussed above, the rapidity of abrupt climate events provides additional clues about how climate is controlled by Earth processes. Abrupt climate events also serve as important time lines, enabling the correlation and analysis of fragmentary stratigraphic records from around the world. Examples of abrupt climate change include the Permian-Triassic boundary (see Question 8), the Paleocene-Eocene Thermal Maximum, and Dansgaard-Oeschger events in the more recent Pleistocene Epoch.
Dansgaard-Oeschger events, named after the geochemists who first documented them, refer to rapid climate fluctuations that occurred about every 1,500 years during the last ice age, especially the interval between 60,000 and 25,000 years ago (Figure 3.6). Each oscillation is characterized by gradual cooling followed by abrupt warming, typically over just a few decades. Even though these changes are rapid, their magnitude is large—annual temperature swings of up to 16°C are recorded in Greenland ice cores. A number of mechanisms have been invoked to explain them, including solar influences (Bond et al., 2001). Some of the coldest events are thought to be related to massive discharge and melting of icebergs, which would have delivered fresh water to the North Atlantic and possibly changed ocean circulation (reviewed in Hemming, 2004). Similarly dramatic but temporary events almost certainly
OCR for page 81
Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 3.6 Record of δ18O, a proxy for mean annual temperature, of Greenland ice from the GISP2 ice core. A change of five units of δ18O corresponds to a change in temperature of 14°C at the GISP site. The important features of this record are the rapid shifts between 60,000 and 25,000 years ago, when temperatures oscillated by 10°C to 15°C over periods as short as 100 years, and the unusual stability of climate over the past 10,000 years. SOURCE: Data from Grootes and Stuiver (1997).
occurred during earlier glacial periods, although high-resolution ice core and marine sediment records are not available to confirm this.
The most extreme abrupt global warming event recorded in geological history was the Paleocene-Eocene Thermal Maximum, which occurred 55 million years ago (Figure 3.3; reviewed in Zachos et al., 2001). In less than 10,000 years, deep-sea temperatures are estimated to have increased by 5°C to 6°C and sea surface temperatures by as much as 8°C at high latitudes (Stoll, 2006). This warming event was associated with changes in global carbon cycling, oceanic and atmospheric circulation, and the extinction of many marine organisms. Detailed chronology of the interval suggests that it took about 170,000 years to flush the excess 12C from the ocean and atmosphere through burial of carbonate and organic carbon in deep-ocean sediments (Röhl et al., 2000).
The cause of this abrupt event is still debated. At least seven possible triggers have been proposed, including a catastrophic release of 1,050 to 2,100 gigatons of carbon from seafloor methane hydrate reservoirs (Zachos et al., 2005). A significant shift in osmium isotopes suggests that continental weathering increased substantially (Ravizza et al., 2001), possibly as a result of increased CO2 in the atmosphere as well as higher temperature and humidity (Zachos et al., 2001).
Can Earth’s Past CO2 History Be Determined?
The connection between atmospheric CO2 levels and climate is generally accepted, but there are still few reliable data confirming the relationship through Earth’s history. The examples above show that additional or alternative factors, including other greenhouse gases like CH4, may be required to explain some temperature changes. For example, estimated concentrations of atmospheric CO2 are too low to explain some of the warmest times of the Cenozoic (Fedorov et al., 2006; Stoll, 2006), and CO2 concentrations were believed to be very high in the Ordovician and Jurassic, despite evidence of episodically cool climate (Kump et al., 1999; Veizer et al., 2000). Confirming a correlation between periods of warm climate and high atmospheric CO2 levels during the Phanerozoic remains a major objective. Other key questions include whether other greenhouse gases were important in the more distant geological past, and whether other causes of climate change besides greenhouse gas forcing can be inferred from the geological record.
Much of the work on deep time has focused on proxy studies of marine sedimentary rocks, which record the evolving chemistry of the ocean. Since the ocean and atmosphere are roughly at chemical equilibrium over timescales longer than 10,000 years, and because most of the available carbon is stored in the oceans, reconstructing past changes in ocean chemistry would help establish how atmospheric CO2 has changed. But because the chemistry of the oceans is so complicated, available data are still insufficient for the task. Further complicating the picture are isotopic data suggesting that steady state models of the carbon cycle are applicable in the Cenozoic (0 to 65 Ma), but not the Neoproterozoic (1,000 to 543 Ma; Rothman et al., 2003). Some aspects of ocean chemistry at least confirm that the oceans undergo major shifts in composition. For example, there is evidence that the ratio of Ca to Mg and the ratio of carbonate (HCO3−) to sulfate (SO42−) have changed markedly and systematically (Figure 3.7).
Similarly, the rates of past volcanism and weathering cannot be measured directly, and better estimates
OCR for page 84
Origin and Evolution of Earth: Research Questions for a Changing Planet
changes in the types of rocks exposed to constant rates of weathering (e.g., Harris, 1995). The relative contributions of these factors will have to be sorted out before we can determine how to translate the Sr isotopic data into a quantitative estimate of global weathering rates. Other possible proxies for weathering include Os, Ca, and Mg isotopes, although these elements are still in early stages of study.
Summary
The geological record teaches us that Earth’s climate has always been changing, but remarkably the surface temperature has remained within a range suitable for life for the past 3.5 billion to 4 billion years. The primary factors responsible for this relatively benign climate are believed to be volcanic emissions of carbon dioxide to the atmosphere, removal of CO2 by weathering of surface rocks, and more subtle effects, such as the positions of the drifting continents, the patterns of ocean currents, the orientation of Earth’s rotational axis and orbit around the Sun, and the luminosity of the Sun. Other chemical and biological effects are also likely to be important, such as the oxidation state of the atmosphere and the concentrations of other greenhouse gases. Interspersed in this vast and mostly life-supporting history are a few periods when Earth was considerably warmer than it is at present, and completely ice free, and a few times when Earth might have been extremely cold and completely ice covered.
At present the greenhouse gas content of the atmosphere is increasing rapidly. The greenhouse gas content of the atmosphere is the most important determinant of climate on geologically short timescales, and models can be used to predict how climate will change over the next decades and centuries. Over longer geological time periods, natural geological processes control the greenhouse gas content of the atmosphere, and other geological and astronomical factors are influential. We have a good qualitative understanding of the factors that contribute to Earth’s natural climate states, but we still lack a comprehensive model that can account for the climate changes of the past or predict climate changes into the distant future. Better models for both the volcanic and weathering components of the climate cycle, more quantitative descriptions of erosion and its relation to weathering, and the incorporation of inputs from the biosphere and other factors will likely lead to a more accurate understanding of Earth’s climate and climate history.
QUESTION 8:
HOW HAS LIFE SHAPED EARTH—AND HOW HAS EARTH SHAPED LIFE?
It is not surprising that many Earth scientists have viewed the geological evolution of Earth as a fundamentally inorganic process—dominated by titanic mechanisms such as mantle convection and plate tectonics. After all, virtually all of Earth’s organic mass exists as a veneer of frail and short-lived creatures within a few vertical miles of the outermost surface, a seemingly insignificant afterthought to this massive planetary body of rock. And yet this multitude of organisms—most of them microscopic packages composed primarily of carbon, hydrogen, nitrogen, and oxygen—determines major features of the atmosphere, oceans, and continents. Biologically influenced processes like erosion and weathering, for example, continually shape and reshape Earth’s surface. And as we have seen in Questions 4 and 5, the erosion and weathering influenced by life forms affect not only the topography and composition of continents but also the chemical composition of subducted crust and therefore the mechanism of plate tectonics and the composition of the mantle.
Life scientists, in the same spirit, have regarded the evolution of life as a fundamentally biological issue, dependent primarily on time, chance, and competition to trend toward increasing diversity and complexity. We now know that Earth itself is not the mere substrate or background for life’s activities as once supposed but rather an active partner in evolution. Geological processes and astronomical events have strongly and repeatedly influenced the story of life on Earth and often determine the kinds of life that can survive and flourish.
The interconnectedness of life and the environment has been a subject of continuing research and debate. An extreme view is that life controls Earth’s surface environment and does so in ways that are most beneficial to the continuation of life (Lovelock, 1979). But evidence in the geological record, especially of mass extinctions, suggests that life cannot always maintain
OCR for page 85
Origin and Evolution of Earth: Research Questions for a Changing Planet
conditions favorable to life. We are far from understanding how much of evolution is purely biological and how much has been forced by Earth processes; nor do we know exactly how much of Earth’s environment is determined by the presence of life. And yet these questions have suddenly become more urgent as we find ourselves in an era when—presumably for the first time—Earth’s surface environment can be manipulated by a single dominant life form, Homo sapiens, that is capable of making choices about the effects of its actions.
How Does Life Affect Geological Processes?
Life affects Earth’s planetary processes in several ways. At the microscopic scale, life is an invisible but powerful chemical force. Organisms can catalyze reactions that would not happen in their absence, and they can accelerate or slow other reactions. The chemical reactions they enhance have a specific character; in general they extract energy from Earth and from sunlight to fuel life processes. These reactions, compounded over immense stretches of time by a large biomass, can generate changes of global consequence. An example of this global influence is the processing of carbon and oxygen. Weathering reactions on land, combined with organic precipitation of carbonate shells in the oceans, remove carbon from the atmosphere and convert it to carbonate minerals on the seafloor (Question 7). Photosynthesis also extracts carbon from the atmosphere, converting carbon dioxide into oxygen plus organic material. Some of this organic carbon is stored in soils, ocean sediments, and the living biomass of the continents and oceans, while the oxygen is delivered to the atmosphere. Larger animals and plants also have physical effects on Earth, such as promoting soil formation and moderating erosion.
Beyond these generalities, we understand little about the details of biologically mediated chemical processes in the environment, especially those of the distant past. Like many fields of science, however, this one is being revolutionized by powerful new analytical tools and computational techniques. For example, new ultrahigh-resolution microscopes can now be used to observe microorganisms in the environment and in laboratory experiments (Figure 3.10). Synchrotron X-ray techniques can be used to study the chemical processes of these microorganisms. Innovative isotopic techniques are being used to help understand the complicated chemical processing that organisms can achieve. DNA sequencing methods have brought a new dimension to studies of microbiological processes. In the past it was difficult to identify the organisms in natural samples because many could not be cultured. Today, organisms do not need to be cultured; their identity can be determined directly from their DNA. Computational chemistry (see Question 6) also shines a strong new light on natural biochemical processes, bringing the possibility of calculating from quantum mechanical theory how atoms and molecules will behave in the microenvironments surrounding tiny organisms.
Soils represent a particularly clear example of how multiple fields, including inorganic chemistry, physics, and hydrology, can wrest new insights from geobiological processes. Inorganic weathering of minerals and organic carbon in the soil environment releases nutrients and carbon. The rate of release and the types of nutrients define the environment in which life can exist and control the range and abundance of life forms that can survive. In addition, the roots of land plants, as well as bacteria, fungi, and animals such as earthworms, can accelerate the weathering of mineral and organic matter in soils. Such biological catalysis of weathering processes can enhance the suitability of soil for life and
FIGURE 3.10 High-resolution images of (A) a cell (outer cell wall indicated by white arrows) and associated mineralized filaments (white) and nonmineralized fibrals (gray) and (B) FeOOH-mineralized filaments filtered from water. SOURCE: Chan et al. (2004). Reprinted with permission from AAAS.
OCR for page 86
Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 3.11 The history of life, based on geological evidence, along with long-term oxygen, ice ages, and mass extinctions. Molecular data suggest that eukaryotic organisms (protozoans, algae, fungi, plants, and animals) share a common ancestor with Archaea.
also speed up the weathering that would have gone on in the absence of life. The ultimate control on the soil environment is probably climate; insufficient rainfall, for example, limits how fast both inorganic and organic chemistry can proceed. But on a global basis we now know that soil chemistry is powerful enough to affect climate by helping to regulate atmospheric carbon dioxide.
Similarly, we know that vascular plants have an enormous effect on Earth’s environment. Life on Earth originated nearly 4 billion years ago, but land plants are found in the geological record only during the past 400 million years or so (Figure 3.11). Several lines of geological evidence suggest that diversifying land vegetation changed the nature of continental weathering, erosion, and sedimentation, changing the physical stability of stream banks and even influencing the composition of Earth’s atmosphere (Berner and Kothavala, 2001). Roots break up rock and help transform it into soil. Deep roots also contribute carbon dioxide to soils, resulting in concentrations of soil CO2 that are 10 to 100 times higher than the modern atmosphere. The high CO2 concentrations in soil gas act to acidify soil water, which leads to increased rates of dissolution of minerals. Deeply rooted plants can also extract water from well below the surface and return it to the atmosphere via evaporation from leaves. This evapotranspiration has an important cooling effect on the land surface, as does the shade provided by the leaf canopy.
There is ample evidence that plants and animals also influence erosion rates, but there is still uncertainty about how important they are in the long-term evolution of continental surfaces and how their effects should be represented in new landscape evolution models (Dietrich and Perron, 2006). Erosion itself affects habitat conditions and can strongly influence biodiversity and ecosystem processes. Hence a central question is the extent to which life and landscape evolution are related. For example, do hillslope shapes and river forms reflect the presence of life, or would Earth’s land surface be more or less the same shape if
OCR for page 87
Origin and Evolution of Earth: Research Questions for a Changing Planet
the planet were lifeless? A related question is whether topography, which is created by mountain uplift and erosion, affects the structure of ecological communities. The prospect of changing climate in the near future brings up other, possibly more urgent, questions. For example, we would like to know whether rates of erosion will change with changing climate and whether climate-induced variations in vegetation will reduce or enhance the response of erosion rates to climate change. To answer these questions, we need much better models of the effects of biota on weathering, erosion, and sediment transport rates. Biotic diversity needs to be linked directly to changes in material strength (resistance to erosion), mass loss, and sediment mobilization from hill slopes. Similarly, ecological theory needs to include explicit physical effects that influence food web processes. These issues lie at the interface of ecological and Earth sciences. Since the locus of life is typically found in the soil that cloaks the landscape, there is a great opportunity to integrate the fields of pedology, hydrology, geobiology, geochemistry, and geomorphology into a new understanding of this life-supporting system (NRC, 2001).
How Long Has Life Fostered a Habitable Surface Environment?
Because organisms on Earth help maintain a life-supporting surface environment today, it is natural to ask whether they have always done so. This question proves surprisingly difficult to answer, however, largely because we have so little evidence from the early geological record. The parts of early organisms richest in information are organic, namely proteins and nucleic acid, but these are also the most reactive and appealing to a gauntlet of other organisms bent on using them as food. Even biomolecules that reach the seafloor after death are usually broken down by decay processes within the sediments. Therefore most of our paleobiological information is gleaned from the bones, skeletons, and other hard parts preserved as fossils in sedimentary rocks.
For the interval of Earth history that begins with the Cambrian Period (542 Ma), paleobiologists have abundant fossils of plants, animals, and selected algal and protozoan groups that preserve a compelling record of ancient diversity, ecology, and evolutionary pattern. Fossils of microorganisms, including the geologically important cyanobacteria, provide at least an impressionistic view of evolution and diversity that extends much deeper into our planet’s early history (see Question 3). Not all organisms, however, produce the mineral skeletons and tough organic materials preserved as conventional fossils, and this is particularly true of the microorganisms whose metabolic capabilities define much of the interface between the physical and biological Earth.
Fortunately, a new set of tools has become available to establish the presence and infer the biological activities of microorganisms in Earth’s history. Most organic compounds in microbial (and, indeed, all) cells decay quickly after death. The exception is lipid molecules found in cell membranes. These hardy compounds, commonly called biomarker molecules, can survive long-term burial in sedimentary rocks and so record aspects of the diversity, environmental setting, and metabolic workings of microorganisms spanning more than 2.5 billion years of our planet’s history. Biomarker molecules led geologists to the understanding that petroleum has a biological origin (Triebs, 1936); they have shown how microbial communities responded to transient oxygen depletion in Mesozoic ocean basins (Kuypers et al., 2002), illuminated the nature of life and environments in Proterozoic oceans (Brocks et al., 2005), and provide our earliest evidence for the presence of life’s great evolutionary branches in late Archean ecosystems (Brocks et al., 2003). Preserved organic molecules have even been reported as biomarkers in 3.5-billion-year-old rocks from Australia (Marshall et al., 2007). Much remains to be learned about the sources, function, and biosynthesis of biomarker molecules, but new research that combines microbiology, genetics, and emerging technologies for analysis (e.g., Brocks and Pearson, 2005) promises unprecedented insights into evolution and environmental history both in the marine realm (e.g., Grice et al., 2005) and on land (Freeman and Colarusso, 2001).
How Did Organisms Influence the Oxygenation of the Atmosphere and Oceans?
Perhaps the most obvious and vital link between life and Earth systems, at least from a human point of view, is the maintenance of abundant atmospheric oxygen, a feature whose development we still do not
OCR for page 88
Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 3.12 (Top) Phanerozoic history of O2 and CO2 inferred from models. SOURCE: Berner (2006). Copyright 2006 by Elsevier Science and Technology Journals. Reproduced with permission. (Right) CO2 inferred from chemical analysis of soil carbonates. SOURCE: Ekart et al. (1999). Copyright 1999 by the American Journal of Science. Reproduced with permission.
fully understand. Today’s atmosphere contains about 21 percent oxygen and only about 0.03 percent carbon dioxide, yet multiple lines of evidence indicate that the atmosphere contained little or no O2 for the first 2 billion years of Earth’s history (Bekker et al., 2004) and may have contained much more CO2. In the past 500 million years, the O2 content of the atmosphere seems to have varied from perhaps 10 to 30 percent and the CO2 content from as low as 0.02 percent to as high as 0.7 percent (Figure 3.12). Oxygen is necessary for many bacteria and nearly all forms of eukaryotic life and is critical to our concept of planetary habitability. Photosynthesis provides the only plausible source of this oxygen, and so oxygenation of the atmosphere and oceans constitutes an essential example of how life has profoundly influenced Earth’s surface conditions. Oxygenic photosynthesis also links atmospheric oxygen with atmospheric carbon, in that the O2 comes mostly from extracting oxygen from CO2 and making reduced carbon in the form of organic molecules.
OCR for page 89
Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 3.13 Schematic of the rise of atmospheric O2 concentrations. The two curves indicate the approximate range of values allowed by available data. Photosynthetic bacteria evolved no later than 2.7 Ga and perhaps as early as 3.8 Ga. Whether the diversity of early photosynthetic bacteria included the oxygen-producing cyanobacteria remains uncertain. Geological evidence indicates that whether or not oxygen-generating photosynthesis evolved in Archean oceans, O2 did not become significant in the atmosphere and surface oceans until about 2.4 Ga. Geological evidence also suggests that there was a further delay before O2 levels became significant in the deep ocean. Subsequently, there were times when the deep ocean became oxygen poor, even though there was appreciable oxygen in the atmosphere. SOURCE: Holland (2006). Reprinted with permission.
Oxygen began to accumulate in the atmosphere and oceans 2.3 billion to 2.45 billion years ago, but the abundance remained quite low for another 2 billion years (Figure 3.13; e.g., Brocks et al., 2005; Canfield, 2005). Considering that oxygen-generating photosynthetic bacteria were already present 2.7 billion years ago, the long delay in oxygenation of the atmosphere is hard to understand (Kopp et al., 2005). Why didn’t the radiation of cyanobacteria—the only bacteria to evolve oxygenic photosynthesis and the progenitors, via endosymbiosis, of chloroplasts in algae and land plants—spread O2 rapidly through surficial environments to produce an atmosphere like the one we have today? Part of the answer is biological: organisms that respire aerobically, from bacteria to humans, gain energy from the reaction of oxygen with organic molecules, reversing the chemistry of photosynthesis. The growth of atmospheric oxygen becomes possible when rates of oxygen production exceed those of aerobic respiration and other reactions that consume O2. For example, burial of organic material (reduced carbon produced by photosynthesis) by sediments inhibits aerobic respiration, paving the way for oxygen accumulation in the atmosphere and oceans.
On the early Earth other processes could also have contributed to the production of molecular oxygen. For example, if the early atmosphere contained much higher amounts of both CO2 and H2O than it does today, substantial hydrogen could have been lost from the upper atmosphere to space. This process would have had the effect of converting H2O to O2. Considering the evidence that only tiny amounts of oxygen were present in Earth’s early atmosphere, however, this process could not have been very efficient (Catling et al., 2001; Tian et al., 2005). Photochemical destruc-
OCR for page 90
Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 3.14 Sulfur isotope data that indicate the atmosphere was effectively devoid of oxygen until about 2,400 Ma. Δ33S represents the mass independent sulfur isotope fractionation, which occurs at high ultraviolet radiation levels. Nonzero Δ33S before 2 Ga implies that ozone (and therefore O2), which absorbs ultraviolet radiation, had very low concentrations. SOURCE: Farquhar and Wing (2003). Copyright 2003 by Elsevier Science and Technology Journals. Reproduced with permission.
tion of methane in the Archean atmosphere could also produce hydrogen that would escape from Earth, again facilitating the oxidation of Earth’s surface (Catling et al., 2001).
However oxygen accumulated in the atmosphere, its consequences were immense. Some bacteria evolved a mechanism to gain energy from the reaction of oxygen gas with organic molecules (aerobic respiration), and the ancestors of modern eukaryotes appropriated this mechanism by capturing respiring bacteria and reducing them to the metabolic slaves we know as mitochondria. It has been proposed that the bacterial ancestors of mitochondria were only facultative respirers, conducting anoxygenic (nonoxygen-producing) photosynthesis in oxygen-free environments (Woese, 1977). If true, the original basis of ecological interaction could have been photosynthetic—but its lasting legacy was unquestionably aerobic respiration in nucleated cells.
We do not yet understand how biological, tectonic, volcanic, and atmospheric processes combined to produce the episodic rise in the amount of oxygen in the atmosphere. In fact, we are only now developing the analytical tools needed to read Earth’s long-term environmental record at high resolution. For example, the recent discovery that the isotopic composition of atmospheric sulfur was subtly different before 2.5 billion years ago (Figure 3.14) confirms that the concentration of O2 in the atmosphere back then must have been less than 10−5 of the present level (Farquhar et al., 2000; Pavlov and Kasting, 2002)—effectively oxygen free.
The rise of atmospheric oxygen also eventually produced a rise in the level of atmospheric ozone, which shields Earth’s surface from ultraviolet radiation that is detrimental to life on land. The ozone concentration that is sufficient to provide full ultraviolet shielding is surprisingly small—an atmospheric O2 concentration about 1 percent of the modern level (Kasting et al., 1985), a level that was probably reached soon after 2.5 billion years ago.
We do not understand whether the initial evolution of cyanobacteria triggered the first round of oxygen accumulation or preceded it by hundreds of millions of years; nor do we understand why oxygen levels remained low through most of the Proterozoic Eon (2,500 to 542 million years ago) or what processes drove the renewed increase that paved the way for animal diversification. For that matter, we do not
OCR for page 91
Origin and Evolution of Earth: Research Questions for a Changing Planet
understand why the modern atmosphere contains the amount of oxygen it does. Oxygen-related questions are sufficiently complex that they will require expanded research interactions among Earth scientists, atmospheric scientists, and biologists. For example, we need better paleontological resolution of when oxygenic photosynthesis first evolved and when the eukaryotic organisms that now dominate primary production rose to global prominence (Falkowski et al., 2004). There is still no reliable geochemical proxy for ancient oxygen abundances (Berner et al., 2003), and models relating deep-Earth processes to surface conditions do not yet take account of historical patterns and feedbacks from physiology, tectonics, and atmospheric chemistry.
Major questions of oxygen history are not limited to its long-term trajectory. During the Paleozoic and Mesozoic eras, wide portions of the oceans beneath the surface mixed layer became essentially “oxygen deserts,” a condition known as anoxia. Geologically transient but globally distributed oceanic anoxic events are well documented from Early Jurassic, Early Cretaceous, and Late Cretaceous rocks (Jenkyns, 2003). In modern oceans, O2 can fall to low levels and, locally, may decline to zero in the ocean-minimum zone just below the well-mixed surface water mass in which most photosynthesis takes place. What makes the oceanic anoxic events stand out is the large spatial scale of anoxic water masses. So far, we know that these events coincide with perturbations in the carbon cycle, as deduced from records of the isotopic composition of marine carbon and the strontium isotopic composition of seawater (e.g., Jones and Jenkyns, 2001; see Question 7). But we do not know what tipped the redox balance, causing anoxia to spread repeatedly through Mesozoic oceans; nor do we know why there are similar (but less well documented) events in Paleozoic oceans but none from the Cenozoic Era. And we do not know whether these events were produced entirely by inorganic geological processes or whether organisms exacerbated, ameliorated, or otherwise responded to these events.
As oxygen levels increased in the atmosphere and oceans, new forms of life became possible. Animals that move about in search of food have an elevated need for oxygen, and so it is not surprising that the first evidence for large animals with high oxygen demands met (initially) by diffusion through tissues coincides with geochemical evidence for elevated O2. Oxygen levels may have reached historically high levels (perhaps as much as 30 percent of the atmosphere, by volume) some 300 Ma, potentially explaining how pigeon-sized dragonflies could fly above tropical forests of the day (Dudley, 2000). Sharp, if transient, depletion of oxygen in ocean waters had the opposite effect, reducing animal diversity and size in widespread areas of the seafloor—a particularly widespread episode of marine anoxia is associated with mass extinction at the Permian-Triassic boundary, some 250 Ma (Wignall and Twitchett, 1996).
Other Interactions Between Earth and Life
Oxygen provides a compelling example of rapidly unfolding research on the interactions between the physical and biological Earth, but it is hardly the only example. Carbon dioxide is also intimately related to biological activity, not only through climate and the carbon cycle (Royer et al., 2001) but also because it affects the ability of marine organisms to form carbonate skeletons (Kleypas et al., 2006). The physiological link between skeletons and carbon dioxide may help explain some major biological changes of the past. For example, accelerating physiological research on the biological consequences of ocean acidification illuminates Earth’s greatest mass extinction at the end of the Permian period (252 Ma) when marine ecosystems collapsed (e.g., see below). Similarly, current increases in carbon dioxide raise concern about the future of reef corals and other organisms that form carbonate skeletons in the shallow ocean. The other side of this interaction—how biological and physical processes interact to govern CO2 on both short and long timescales—is also a major issue in Earth history, one of immense importance as we debate the consequences of current human activities (Question 7). As in the case of oxygen, deeper understanding of the feedbacks between life and carbon dioxide levels, on scales from the local and ephemeral to those governing the long-term history of the planet, will require better geochemical and paleobiological proxies for ancient CO2 abundances, more nuanced understanding of the biological processes that influence carbon dioxide levels (especially those related to microorganisms and plants), and increasingly sophisticated models that account for both biological and physical parameters.
OCR for page 92
Origin and Evolution of Earth: Research Questions for a Changing Planet
FIGURE 3.15 Number of marine animal genera through time, showing the five major times of “mass depletion” in biological diversity. Only three drops—end-Ordovician, end-Permian, and Cretaceous-Tertiary—are driven primarily by increases in extinction rates, rather than declines in rate of origin. SOURCE: Bambach et al. (2004). Copyright 2004 by the Paleontological Society, Inc. Reproduced with permission.
Problems as diverse as the influence of rainforests on Earth’s hydrological cycle, the role of vegetation in stabilizing the land surface, the relationship between nutrient availability and diversity, and the oceanwide biogeochemical consequences of deep-water anoxia engage a wide range of Earth scientists because they have both deep-time evolutionary components—how did the diversification of woody plants change Earth’s surface?—and topical applications—what will be the consequences for the Earth system of rainforest clear-cutting, increased soil erosion, and seafloor anoxia linked to fertilizer-spiked nutrient flows from agricultural lands to the ocean? Earth scientists have almost limitless opportunities to join with biologists to fashion both a new picture of our planet’s history and a clearer picture of our future.
What Caused Mass Extinctions?
Nothing illustrates how heavily life depends on a favorable surface environment as clearly as a sharp change in that environment—which has occurred several times during the past 500 million years, causing the mass extinction of species (Figure 3.15). In particular, the great extinctions at the end of the Permian (252 Ma) and Cretaceous (65 Ma) periods influenced the course of biological evolution as much as all the accumulated genetic changes during the 187 million years between them. But what specific events or environmental changes precipitated the great mass extinctions, and what aspects of biology influenced the patterns of survival and recovery, are not known. Most Earth scientists agree that a meteorite impact caused the end-Cretaceous extinction of dinosaurs, ammonites, and myriad other plant, animal, and microscopic species (Alvarez, 1997), but the actual kill mechanisms unleashed by this trigger remain poorly understood. The relative importance of coincident environmental perturbations, including an interval of oceanographically driven global change, extensive extrusion of flood basalts, and the particular location of the impact on a tropical continental platform, are simply not known.
Although a single plausible event may account for the end-Cretaceous extinctions, the cause of the end-Permian mass extinction, which may have erased as many as 90 percent of marine species and many terrestrial species (Erwin, 2006), is still debated. Support for an extraterrestrial cause is limited, with growing interest in direct and indirect effects of massive volcanism in what is now Siberia. An emerging view is that massive flood basalts, intruded through thick carbonates and extruded onto thick peat deposits, produced unusually high emissions of carbon dioxide and thermogenic methane, resulting in global warming, acidification of the oceans, depletion of oxygen in ocean waters below the mixed layer, and enhanced production of hydrogen
OCR for page 93
Origin and Evolution of Earth: Research Questions for a Changing Planet
sulfide by bacteria living in those oxygen-depleted water masses. Physiological research on modern marine organisms, aimed at understanding current environmental change (e.g., Pörtner et al., 2005), allows Earth scientists to predict the biological consequences of such an event on end-Permian biological diversity. Indeed, paleobiological data show that extinctions did not affect all Permian animals equally. For example, groups whose living relatives were vulnerable to the physiological consequences of sea acidification disappeared at rates much higher than those physiologically well buffered against such environmental perturbations. Extinctions on land are consistent with the predicted effects of rapid climate change (summarized in Knoll et al., 2007). Continuing research on Earth’s great intervals of biological upheaval will increasingly integrate insights from paleobiology, stratigraphy, high-precision geochronology, and geochemistry with physiology and models generated to help understand current issues of global change.
What Governs the History of Biological Diversity?
Major extinctions have clearly influenced the history of plant and animal life, but what, fundamentally, controls the observed pattern of diversity increase from the Cambrian to today (Figure 3.15)? Quantification of diversity change through time on land and in the oceans remains a subject of active research and debate, but many Earth scientists would agree that the modern world (at least in preindustrial times) harbors more species of land plants, more species of land animals, and more species of marine animals than any previous moment in our planet’s history (e.g., Benton and Emerson, 2007). Attempts to model diversity history employ logistic equations, which imply biologically or physically imposed limits to diversification (e.g., Sepkoski, 1984), or exponential equations, which imply persistent diversity increases, episodically knocked back by mass extinctions (Stanley, 2007).
The tension between these classes of models focuses attention on a great and unsolved problem. What are the relative roles of genetic innovation, ecology, and physical Earth history in governing the long-term history of life? The answer certainly requires macroecological insights from biologists, but the questions are necessarily framed by paleontologists. And rapidly emerging insights into the physical history of the Earth surface system provide, for the first time, the proper environmental framework to address the issue. Has primary production increased through time, and if so what have been its consequences? What are the consequences of sea-level change, episodically flooding and exposing continental interiors, on species origination and extinction in the marine realm (Peters, 2005)? Did the rules of community construction change when flowering plants evolved the capacity to use animals to ensure the faithful spread of pollen from one plant to the next? How did the ecological relationships that undergird community diversity reform following episodes of mass extinction? Detailed analyses of community organization in systems as disparate as Pleistocene coral reefs, Cenozoic mammals, and Carboniferous forests promise important insights into ecology and evolution that cannot be made solely on the basis of the short-term observations and experiments available to biologists (Jackson and Erwin, 2006).
Summary
Earth’s surface environment is obviously altered by large-scale geological processes (Questions 4 and 5), but it is also affected continuously and pervasively by the activities of life forms. Likewise, Earth’s geological evolution and infrequent catastrophic events, such as meteorite impacts, have clearly affected the evolution of life. But even when we can document extinctions and major evolutionary changes, we cannot yet sort out the causes. To what extent were they caused by geological as opposed to biological processes? Which environmental conditions were responsible for which extinctions or changes in biological form and function? We know that the composition of Earth’s atmosphere, especially its high concentration of oxygen, is a major consequence of the presence of life, one that made possible the evolution of more complex organisms. But exactly how other geological events have affected evolution, and how much control life has had on climate, are still topics of debate.
Life processes and Earth processes also interact locally. Erosion rates, climate, and weathering rates affect the habitability of specific regions of Earth, and the ecosystems themselves in turn affect erosion rates, climate, and weathering processes. Understanding the
OCR for page 94
Origin and Evolution of Earth: Research Questions for a Changing Planet
interrelationships between surficial processes that shape the land and the life that inhabits it presents a critical challenge for managing land resources and becomes even more important as we attempt to forecast the effects of future climate change.
Understanding how Earth’s life and geological environment arrived at their present state, and how they interacted in doing so, constitutes a major intellectual challenge. Meeting this challenge will help us understand how life will respond to present-day environmental change, but Earth scientists will have to develop new research and educational partnerships with biologists and atmospheric scientists. The search for life on extrasolar planets will similarly depend on better understanding of biogeochemical influences on atmospheric composition here at home.