Carbon Dating

There is a third carbon isotope, one with 8 neutrons, designated as C- 14. It is not as stable as C- 12 and C- 13, and because nature always strives to achieve stability, it is eventually converted to an isotope with a more stable nucleus. The conversion takes place through the following nuclear reaction:

C- 14 → N- 14 + e– . [3.1]

In this reaction, one of the neutrons in C- 14 is converted to a proton. This increases the number of protons by one and hence converts carbon with an atomic number of 6 to nitrogen with an atomic number of 7. However, a neutron is not electrically charged, whereas a proton is. In order to conserve charge, the conversion must include a negative charge so as to compensate for the positive proton charge. The negative charge is consequently carried by the electron emitted in the reaction. This emission is an example of radioactive decay. Historically, the emitted electrons are called beta particles and the radiation of such electrons is called beta radiation.

The decaying of C- 14 to N- 14 and the other similar nuclear transformations are very robust processes. They follow kinetics in which they decay in an exponential way: the amount of C- 14 decreases by a fixed percentage over a given amount of time. The time that it takes to reduce an amount by 50% is designated as the half- life of the particular radioactive isotope. The half- life of C- 14 is 5700 years. Because the half- life is a nuclear property, it is not affected by external environmental factors such as temperature and the chemical environment.

C- 14 is formed in the outer atmosphere when nitrogen is bombarded by neutrons coming primarily from the sun’s cosmic radiation. The nuclear reaction is essentially a reverse of the reaction in equation 3.1:

N- 14 + n → C- 14 + p+ . [3.2]

The rate of formation and the rate of decay determine the steady- state concentration in the atmosphere. C- 14 is rapidly oxidized to carbon dioxide, labeled in C- 14 and dispersed into the atmosphere. The rate of this dispersal was determined by measurements of radioactive carbon produced from hydrogen- bomb testing.The carbon dioxide is removed from the atmosphere either by being assimilated by the biosphere through the photosynthetic process or by being dissolved in the ocean in one form or another. The flow of carbon dioxide will be fully discussed in the next chapter. Once C- 14 is removed from the atmosphere, it can only decay. If the carbon is blocked from access to new atmospheric carbon, then the ratio between C- 14 and stable C- 12 will decrease with time.This ratio serves as an accurate clock to determine the time since the carbon was removed from access to the atmosphere, which usually is the time in which the photosynthetic organism dies and is no longer able to assimilate carbon dioxide.

The approximately constant concentration of C- 14 in the atmosphere is small: approximately one C- 14 for a trillion C- 12. However, this amounts to approximately 50 billion C- 14 atoms for 1 g of atmospheric carbon. Our sensitivity of detecting radioactive radiation is very high— better than one particle per minute. Fifty billion C- 14 atoms with a half- life of 5700 years correspond to a decay rate of 10 particles per minute. When a sample is isolated from atmospheric exchange for 5700 years, the number of C- 14 atoms (per gram of carbon) is reduced by half, bringing it to 2.5 billion and the corresponding rate of radiation to 5 particles per minute. After 5700 × 2 = 11,400 years (two half- lives), the number of atoms is reduced to 1.25 billion and the rate of radiation to 2.5 particles per minute, and so on.

Radiometric Dating

In radiometric dating, we compare the measurements of the ratio of an unstable isotope with that of a much more stable one. Radioactive dating determines how long ago the object being dated became a closed system and was no longer exchanging material with the environment.

In the case of rocks, closure to the environment occurs when the temperature goes below the closure temperature— that is, the lowest temperature at which the system stops exchanging material with the environment. For biological materials, closure usually occurs with the death of the photosynthetic organisms responsible for the assimilation of the atmospheric carbon dioxide. In the case of air bubbles in permanent ice, closure comes when enough pressure accumulates to seal the pores in the ice.

An essential step for accurate radioactive dating is to determine the isotopic concentration at the time of closure. For carbon, it means the atmospheric isotopic ratio at the time when the organism was alive, and for other elements, it means verification of the absence of daughter products that should result from the nuclear transformations at the time of closure. Carbon dating is useful to about 70,000 years (about 12 half- lives).

Recent remains of biological systems are mainly dated through the ratio of C-14 to C-12, whereas longer time scales are estimated through the decay of other elements. Table 3.2 summarizes the principal isotopes used to date geological formations.

Isotopic Fractionation

The rate of most physical processes that involve movement of atoms or molecules is somewhat influenced by the particular isotope that participates in the process. The reason for this is that, in most cases, the rate of the process depends on the velocity of the atoms or molecules that participate and the velocity depends on the mass of the participants. This is similar

Table 3.2.

Radioactive isotopes used to determine the age of rocks

to two runners who are equal in every respect except that one of them carries a bit of extra weight. If the weight is completely passive (i.e., does not add strength), then the runner with the extra weight is at a slight disadvantage. In chemical terms, the result is that at any particular time, the product will be slightly richer with the lighter isotope than with the heavier isotope. The slightly different rates will result in an isotopic composition of products that will depend slightly on the mechanism of the formation of these products and on the prevailing conditions at the time of the product formation. These small differences in isotopic composition can often serve as fingerprints for the conditions under which the products were formed. A useful parameter used to characterize this difference is shown in Box 3.2, using carbon as an example.

⎤ ⎥ ⎥ ⎥ ⎥⎦

⎞⎟⎠⎞⎟⎠⎛⎜⎝

−

⎞⎟⎠⎛⎜⎝

Box 3.2

Table 3.3.

delC-13 for various carbon sources (per mil)

Source: Trumbore and Druff el (1995).2

oceans because photosynthetic planktonic organisms prefer lighter carbon from the shallow water, which is used to make everything but their shells. W hen these organisms die, they sink to the deeper ocean, where organic remains decompose, returning the carbon to the water. Th e difference in the delC- 13 between shallow and deep water is a measure of the effi ciency of the biological pump. We will go further into this topic in the next chapter. Th is diff erence can be measured by studying the surface- dwelling shelled organisms and deep-sea- dwelling shelled organisms. In addition, limestone tends to measure and preserve the carbon isotopic signature of ocean water. In Table 3.3 we see that methane is very negative. This is because it is the end product of a number of biochemical steps, each one with a preference for the lighter carbon.

DATA

Based on the tools that we have described, we will trace the history of average global temperature and the composition of the atmosphere as far back as 5 million years ago. At that time, humans did not use thermometers to measure temperature because the genus Homo can be traced back only as far as 2.5 million years ago. Even if our early ancestors had had these capabilities, we would not know much about the results because of the lack of lasting communication in the form of writing. This is approximately the time that descendents of African apes started to walk on two feet instead of four. In the absence of direct information, we need proxy measurements to tell us the story. We need physical phenomena that will “remember” history. I will describe a few of these proxies; each proxy provides the information based on a specific set of assumptions, and a check for the validity of many of these assumptions lies in our ability to get a coherent picture from diff erent proxies. We will start by using corals as proxies because we can test the data collected on them with direct measurements. The corals will lead us to small shells of planktonic organisms that, after the death of the plankton, are deposited on the ocean floor. Coring of these ocean floors can provide us with information about the climate’s earliest history. I will then proceed with ice core measurements as proxies because they provide us with the ability to get annual information from as far back as nearly a million years ago, with the added advantage that they can store isolated air bubbles that can provide us with information about the history of the composition of the atmosphere. I will continue with other proxies such as tree rings, fossil pollen, and ocean and land sediments. In all cases the strength of the data lies in our ability to use a combination of the layer’s relative dating with absolute dating.

Corals

Corals build their hard skeletons from calcium carbonate synthesized when lime (CaO) reacts with carbon dioxide according to the following reaction:

CaO + CO2 → CaCO3. [3.3]

Calcium carbonate is the main constituent in limestone, a sedimentary rock that forms from the sinking of the shells of corals and other aquatic organisms. Limestone forms the largest carbon depository on Earth. This issue will be discussed in the next chapter when I discuss the carbon cycle. Corals will come up again when I discuss advance warnings of global warming. in Chapter 15.

Coral skeletons that form in the winter have different densities from those that form in the summer because of the variations in the skeleton’s growth rate in relation to temperature and cloud cover. This manifests itself in the form of annual growth bands that form an excellent basis for relative timing, as previously described. Corals can live for several hundred years; after they die, the skeletons sink into the ocean, where they can be dated. The dating of the corals in the deposit can be based on their C-14 content, as well as their content of Th - 230, which incorporates in small amounts into the skeleton. The half- life of Th- 230 is 75,200 years. Because the skeletons form in the shallow water where corals grow, the isotopic content of the oxygen in the skeleton reflects the isotopic content of the oxygen in the shallow water and hence the temperature of the shallow water.

Figure 3.1 shows an annual comparison of the sea surface temperature (SST) directly recorded in the Galápagos Islands with the delO-18 of oxygen. The data are given on a quarterly basis from 1936 to 1982. The data in Figure 3.1 were first normalized on a yearly basis (taking the four 3- month periods, adding them, and dividing by four). This was followed by taking the average for the whole period and presenting the deviations from this average in Figure 3.1.

The fit is not perfect, but the trends follow each other closely: almost every peak and valley of the surface temperature is matched by a peak and valley in the fractionation measurement of O-18. The largest deviation in the normalized amplitude (between 1945 and 1950) amounts to 1°C.

Plankton

The earliest records for ancient climate come from the coring of deep- ocean sediments of sunken ocean shells with the same carbonate chemistry as in equation 3.3. These shells belong to small, nonphotosynthetic ocean drift ers—zooplankton. The ones most commonly used for measurements are the foraminifera, also known as “forams.” They are approximately the size of a grain of sand, and they deposit under favorable conditions to build layers at a rate of 2.5 cm/10,000 years. As with the corals, the isotopic compositions of their shells represent the water conditions at the time that the shells were formed. Water temperature is one of the most important parameters that determine the shells’ composition, and thus the shells serve as a good proxy to the temperature of the water at the time of the formation of the shells.

Ice Cores

At present, there is an international effort to probe the ice sheets of Antarctica and Greenland— and any other location in the world where thick, old ice deposits have the potential to reveal their secrets as to the climatic conditions at the time that the snow of a given era fell.4,5 Scientists are investigating global climate information by looking at the correlations of the climate in other locations with the ice core locations.6,7 The most investigated locations for ice cores being used for this purpose are in Antarctica— one in Vostok that covers about 400,000 years and another, more recent survey conducted by the European Project for Ice Coring in Antarctica (EPICA) that focuses on East Antarctica, which was able to provide high- resolution data for a period going back more than 800,000 years. The isotopic composition of the hydrogen and oxygen in the ice provides information about the temperature during the time that the ice was formed. The timing is determined by counting the layers of snowfall. Here, however, there is an additional benefit because during the consolidation of the ice layers, air bubbles, isolated from the outside world, remain behind. The gases trapped in the bubbles reflect the atmospheric composition at the time the bubbles were sealed. Here I present the data from the Vostok site. The data describe isotopic fractionation of oxygen; nitrogen of the trapped air; the isotopic fractionation of the hydrogen and oxygen of the ice; and carbon dioxide, methane, and other atmospheric gases. Deuterium and O-18 are used to trace the climate. Cooling of 10°C results in a decrease of 9 per mil in the deuterium fractionation ratio. Figure 3.2 shows the temperature profile deduced from the deuterium fractionation at the Vostok ice core in east Antarctica. Figure 3.3 shows the profile of the concentration of carbon dioxide, and Figure 3.4 shows the timeline of the atmospheric methane concentration.

Tree Rings

Tree growth is influenced by climatic conditions— patterns in tree rings such as width, density, and isotopic composition reflect variations in climate. In regions with distinct growing seasons, one usually finds annual ring formation and thus a record of the climate for that year. Trees can grow to be hundreds, even thousands, of years old. The oldest one is reputed to be a bristlecone pine that grows in California. It is more than 4700 years old and was given the name “Methuselah” for obvious reasons. Such trees can record annually resolved climate for long periods. Perhaps their greatest utility is to calibrate the C- 14 atmospheric concentrations for extended periods of time to enable dating of specimens that died long ago. This is necessary because C-14

dating is based on the assumption that the atmospheric concentration of C- 14, which the organism assimilates while alive, is known. The concentration is approximately constant, but it fluctuates because of the variations in solar radiation responsible for the formation of C- 14.

Fossil Pollen

Flowering plants produce pollen; therefore the analysis of pollen grains that have been preserved in sediment has the potential to identif y specific eras in which these plants existed. One can also infer the climate from the time period and nature of the plants embedded in the sediment. One of the issues that will be discussed in Chapter 14, where we will also discuss the warning signs for global warming, is the need for people to change the seed mixes used in their home gardens because of perceived climatic changes.

ANCIENT HISTORY AND MORE RECENT HISTORY

In Chapter 1, I argued that ever ything that marks the human experience in relationship with the global environment is dominated by unique events. We can learn about climate change neither from our predecessors nor from our contemporaries on some other planet

(in the case of the latter, we have no idea if they exist). However, we are trying to learn from past changes in the physical environment that had major impacts on the biosphere. Going back to Table 3.1, the last 500 million years show the history of multicellular organisms. It is not a smooth evolutionary history. It was punctuated by five known global events that caused the extinction of most of the life- forms on Earth. These punctuations took place at the end of the Ordovician period (443 million years ago), the end of the Devonian period (374 million years ago), the end of the Permian period (251 million years ago), and the end of the Triassic period (201 million years ago); the most well- known extinction took place at the end of the Cretaceous period (65 million years ago) and caused the extinction of the dinosaurs, most likely due to impact of a large meteorite, a remnant of which can be found in the Yucatán Peninsula in Mexico. It was an astronomical event that can obviously repeat itself, but fortunately not too often. Some of these extinctions and the recoveries from them (through major changes in the species distribution) were fast on the geological time scale, and some were much slower. The events that led to the earlier extinctions are still being intensely researched. Most of the methods discussed previously do not extend over such long periods, and therefore good data are hard to come by. Information about atmospheric composition and temperature depends on the interpretation of geological formations. The available paleoclimatological data of the earlier extinctions do not favor another astronomical event but rather enhanced volcanic activity that resulted in a major increase of atmospheric carbon dioxide. Numbers such as 1000 ppmv (3 times the present concentrations) and a temperature increase of 5°C are reported. The review by Peter Ward in Scientifi c American summarizes the scientifi c evidence for the extinctions and discusses an interesting mechanism that involves an explosion of population of anaerobic bacteria that gained advantage due to the high temperature and the oxygen that comes with a volcanic eruption.9 The bacteria release large quantities of hydrogen sulfi de that essentially poison the planet. As we will see, the business- as- usual scenario of energy use that relies on fossil fuels is predicted to cause similar conditions toward the end of the century.

Going forward in history, the Vostok data in Figures 3.2– 3.4 are qualitatively reproducible in other locations in which such measurements are feasible. These data are all very “noisy,” with strong repeating peaks of intervals of about 100,000 years. In terms of temperature, at present we are at the height of such a peak. The period of time between now and the previous peak is the recent ice age. The EPICA data described previously and that had extended the ice core dating to close to a million years ago matches the Vostok data, but for times earlier than 400,000 years, the periodicity loses its sharpness. As one goes even further back in history, the deep- ocean coring of planktonic shells shows these changes in

periodicity in an even more pronounced way. The astronomical origin of these cycles will be discussed in Chapter 7.

Human civilization can be traced approximately to 10,000 years ago, which marks the warming period from the last ice age. Therefore it is obvious that these profiles have absolutely nothing to do with human activities. The temperature profile of this period is shown in Figure

3.5. These data were taken by simply expanding the Vostok data (Fig. 3.2). However, as we trace the temperature to more recent times, more proxies are available, and the number of places that one can monitor climate change extends well beyond Antarctica and Greenland, where ice cores are preserved. Qualitatively, the features shown in Figure 3.5 repeat in other places to such a degree that we give names to different peaks and troughs. These are marked in Figure 3.5. The recent cooling between 1400 and 1800 AD is named the “Little Ice Age,” the warming before that, the “Roman Climate Optimum,” and so on. We can also see that none of these peaks and valleys comes even close to the temperature differences between the ice ages and the average warm periods in between the ice ages. The noisy temperature data clearly indicate that a present heat wave here or there cannot be directly traced to human activities. However, this is not the case with the chemical composition of the atmosphere with respect to gases such as carbon dioxide and methane. These can be directly traced to human activities because their concentrations are considerably higher (see Chapter 2) than any historic levels (limited to 1 million years of ice core studies) of these gases, and indirect measurements, which include atmospheric oxygen content and isotopic profiles, point directly at anthropogenic activities. This will be discussed in some detail in the next chapter. Furthermore, the rate of atmospheric buildup can be measured in terms of years and not thousands of years. In Chapter 6, I will show that such an atmospheric buildup must result in higher average temperatures.