3: Use of Isotope Analysis in Measuring Climate Change

Absolute and proxy measurements for CO₂ and temperature

It's simple to see how CO₂ levels over the last 800,000 years have been determined from ice cores since ancient air is trapped in bubbles in them. The bubbles can be liberated by melting/shattering and analyzed. However, how can we infer temperature changes or actual temperature from ice core samples and both CO₂ and temperatures from ocean sediment cores? Scientists use "proxies" to determine temperature as modern thermometers and temperature scales were invented only recently (early- and mid-1700s by Fahrenheit and Celsius). These proxies include tree rings, growth bands in coral, pollens found in core samples, and calcite shells from marine organisms found in lake and ocean sediments. The organisms include certain types of algae, phytoplankton like dinoflagellates and diatoms, and foraminifera (single-celled protozoans with shells).

From a more chemical perspective, the analyses of the isotopic compositions of ice, using the isotope ratios of ¹⁸O/¹⁶O in water, and of marine sediments, using the ratio of ¹⁸O/¹⁶O and ¹³C/¹²C in carbonate-containing shells, have proved critical in determining temperatures back millions of years ago. Even molecules in leaf waxes (as discussed in Chapter 31.1) can be used. Isotope analyses of diatoms (with silicate shells) in lake sediments are also useful.

The analysis of proxies is quite complicated since many factors contribute to the measurement derived from proxy use. Take tree rings as an example. The width of a given ring depends not only on temperature but precipitation. Each proxy must be calibrated using alternative data, such as from tree rings. Proxy data from ice cores go back a few million years, while data from marine sediments back as far as 100 million years ago. Isotope analysis in rocks formed from marine or land sediments can go back billions of years. Proxy data (tree rings) taken from recent times can be compared to actual temperature measurements from the same time. The calibration relationships can then be used for past samples. Alternatively, proxy data can be taken across many different places and temperatures to develop calibration constants, an approach useful for pollen analysis. Organisms could also be cultured in different temperature, nutrient, and CO₂ conditions to calibrate past data. From a statistical sense, it is best to combine multiple proxies to develop past temperature records. Proxy data sets are available for over 10,000 sites around the world. Figure \(\PageIndex{1}\) below links to sites compiled by Carbon Brief, where databases with information about each study are available.

Figure \(\PageIndex{1}\): Map of over 10,000 proxy data studies and sites. Robert McSweeney and Zeke Hausfather and Tom Prater. https://interactive.carbonbrief.org/...-distant-past/

Most people might care little about climate change millions of years ago. The value in understanding climate change so far in the past is, in part, to build confidence in the data, methods of analysis, and climate model to better understand the relationship between CO₂ and temperature. Some, particularly the PAGES 2k Consortium, focus on the last 2000 years of the Common Era. Figure \(\PageIndex{2}\) shows global mean temperatures obtained from proxy data (yrs 0 - 2000) and direct observations (through the use of thermometer and satellite measurements) since around 1850.

Figure \(\PageIndex{2}\): Global mean surface temperature reconstruction (yellow line) and uncertainties (yellow range) for the years 0-2000 period from the PAGES 2k Consortium along with observations from Cowtan and Way from 1850-2017. Data available in the NOAA Paleoclimate Archive.

Note the overlap of proxy measurements and direct observations of temperatures since around 1850.

Ocean Microorganisms

Biochemistry students come from many backgrounds, but not all have a strong biology background. Let's look at a few relevant to this chapter to help with that and to develop a sense of wonder about the microorganisms that inhabit the oceans and play such a key part in the biosphere. These descriptions are probably new for those with a "chemistry-centric" focus.

Plankton

The word plankton derives from a Greek word meaning drifter or wanderer. There are two main types. One is zooplankton, which are not plants but rather microscopic animals and protozoans. They are heterotrophs that don't synthesize their food. Most have calcite shells. The other is phytoplankton, autotrophic plants that use photosynthesis for food production. Hence, they are carbon "capturers", key players in the carbon cycle in maintaining atmospheric CO₂ and producing O₂. The phytoplankton broadly includes algae (protists), cyanobacteria (also known as blue-green algae), and dinoflagellates which also fit into other groups. Here are some examples. Also, remember that protists are eukaryotic organisms, not animals, plants, or fungi.

Zooplankton

Existing shells in sea sediments from foraminifera have been critical in dating studies and determining CO₂ and temperatures over millions of years. Two major types are benthic foraminifera (which live at the sea bottom and in sediment) and a smaller group of planktonic foraminifera, which live near the surface. They are heterotrophs, but on ingesting small autotrophic phytoplankton, some can retain and sequester their chloroplasts, which can engage in photosynthesis for a while. Figures \(\PageIndex{3-5}\) below show examples of zooplankton that have been important in climate studies.

Figures \(\PageIndex{3}\) above: Benthic foraminifera:

Live mostly at sea bottoms and in sediment); capture carbon indirectly through the carbon cycle by incorporating CO₃^2- into their shells. Living benthic foraminifera in the Bohai Sea, showing normal specimens and abnormal individuals (indicated by arrows).

Figures \(\PageIndex{4}\) above: Planktonic foraminifera

They live near the surface but are buried in ocean sediments after death. (a–h) Nano-CT scan of planktonic foraminifera specimens with a color map of test thickness, warm colors indicating areas of relatively thicker shell; (a,b) Globigerinoides ruber (Tara), (c) Globigerina bulloides (Tara), (d) Neogloboquadrina dutertrei (Tara), (e) G. ruber (Challenger), (f) Trilobatus trilobus (Challenger), (g) N. acostaensis (Challenger), (h) N. dutertrei (Challenger); (i–p) SEM images of selected planktonic foraminifera specimens; (i) T. trilobus (Tara), (j) G. ruber (Tara), (k) G. ruber (Challenger), (l) G. bulloides (Challenger), (m,n) G. ruber test cracked to reveal wall texture (Tara), (o,p) G. ruber test cracked to reveal wall texture (Challenger).

Figures \(\PageIndex{5}\) Above: Radiolaria .single-cell protists that secrete silica

Figures \(\PageIndex{3-5}\): Examples of zooplankton. Benthic foraminifera: https://commons.wikimedia.org/wiki/F...raminifera.p; Planktonic foraminifera: Creative Commons Attribution 4.0 International License. Fox, L., Stukins, S., Hill, T. et al. Quantifying the Effect of Anthropogenic Climate Change on Calcifying Plankton. Sci Rep 10, 1620 (2020). https://doi.org/10.1038/s41598-020-58501-w. http://creativecommons.org/licenses/by/4.0/.; Radilaria: https://commons.wikimedia.org/wiki/F...ria-sp2_hg.jpg

Phytoplankton

Phytoplankton are microscopic plants that use photosynthesis, capture CO₂, and produce O₂. Hence they are primary autotrophs. We will consider three types: diatoms, photosynthetic dinoflagellates, and coccolithophores. Diatoms and photosynthetic dinoflagellates are the major ones and are prey for the zooplankton. They are described in Figures \(\PageIndex{6-8}\) below.

Figures \(\PageIndex{6}\) Above: diatoms

Single-celled eukaryotic algae surrounded by a silica shell (test). These can reach 1 mm in diameter and form an assortment of shapes. Some can form multicellular chains. They engage in high-efficiency photosynthesis, resulting in carbohydrate synthesis. They are found in coastal and cold waters, which contain lots of nutrients.

Figures \(\PageIndex{7}\) Above: photosynthetic dinoflagellates

Algae with a single shell. They are smaller than diatoms. Most have two flagella for motion. They have a cellulose shell, which degrades on death. Hence, they don't have shells that enter the sediment. Some are nonphotosynthetic and are considered zooplankton.

Figures \(\PageIndex{8}\) Above: coccolithophores

Coccolithus pelagicus is covered with a CaCO₃ shell, a coccosphere. These are very small single-cell algae that form interlinked calcium carbonate circulate plates that cover the surface.

Figures \(\PageIndex{6-8}\): Some phytoplankton. Diatoms: https://commons.wikimedia.org/wiki/C...le:Diatom2.jpg; photosynthetic dinoflagellates: https://commons.wikimedia.org/wiki/F...lagellates.jpg; coccolithophores: https://commons.wikimedia.org/wiki/F..._pelagicus.jpg

Along the coast in summer, nutrient-rich upwelling of water can lead to the explosive growth of dinoflagellates, causing the water to become red-gold (often called a red tide). Some species in these blooms produce neurotoxins such as saxitoxin (an inhibitor of sodium channels), which can produce paralytic shellfish poisoning if shellfish from the bloom area are eaten, and brevetoxin (which stimulates voltage-gated sodium channels in nerve and muscle).

Ice cores from Antarctica and Greenland can extend to over 3.4 km (2.1 miles) in depth and yield direct information on CO₂ and indirect temperature measurements. Until recently, continuous ice core records covered 130,000 years in Greenland and 800,000 years in Antarctica. In a remarkable accomplishment, a single continuous ice core 2,800 meters or 1.74 miles long that extended to the bedrock in Antarctica dates back 1.2 million years is now available. Analysis of this core might reveal the cause of the extension of the glacial cycle from 40,000 to 100,000 years, starting around 1 million years ago.  Data going back 2 million years is available using discontinuous cores. Concentrations of trapped CO₂ as a function of time and the temperature of each layer in these cores can be determined. Figure \(\PageIndex{9}\) shows a section of an ice core from the West Antarctic Ice Sheet Divide (WAIS Divide).

Figure \(\PageIndex{9}\): The dark band in this ice core from the West Antarctic Ice Sheet Divide (WAIS Divide) is a layer of volcanic ash that settled on the ice sheet approximately 21,000 years ago. Credit: Heidi Roop, NSFhttps://icecores.org/about-ice-cores

The remains of plankton shells described above are found in cores from sea sediments. Analyses of shells, especially from foraminifera, have provided climate data going back tens of millions of years ago. For ice and ocean sediment cores, isotope analyses have been the key to obtaining CO₂ and temperature data.

Isotope Analyses

In analyzing ice and ocean sediment cores, three things are needed: the age of the layer, a direct or indirect measure of the atmospheric CO₂ at the time the layer was deposited, and an indirect measurement of the temperature at the deposition time. As shown in the figure above, ice core samples have rings, similar to trees, that can be used to count backward in time. The rings get harder to distinguish the further back you go. Figure 3 above shows a visible dark band deposited by volcanoes 21,000 years ago. Ultimately, isotopic analyses of H₂O and CO₂ in ice samples and of carbonates in minerals and deposited microfossils in ocean samples are critical in determining past CO₂ and temperature values.

Most readers are familiar with ¹⁴C radioisotope dating and ¹³C-NMR. Metabolic pathways have been elucidated using ²H (deuterium), ³H (tritium), ¹³C, and ¹⁴C to label specific atoms in substrates and follow their flow into products. These same isotopes have been used in kinetic experiments to determine enzyme reaction mechanisms. In this section, we will explore isotopes in some detail.

Use of unstable radioactive isotopes

¹⁴C radioisotope dating is limited in climate analyses, given its short half-life (t_1/2 = 5730 years). In contrast to most isotopes made in stellar nucleosynthesis or by the radioactive decay of a precursor radioactive elements to an isotope of another element, ¹⁴C is made continually in the atmosphere when high energy neutrons (n) from solar radiation react with atmospheric nitrogen (N). The neutron kicks out a proton to form ¹⁴C, as shown in the nuclear reaction below.

\[ n+{ }_7^{14} \mathrm{~N} \rightarrow{ }_6^{14} \mathrm{C}+p
\]

¹⁴C becomes oxidized to form ¹⁴CO_2,which can enter the carbon cycle and the organic carbon pool through uptake by photosynthetic organisms and organisms that consume them. It can also form inorganic bicarbonates and carbonates, which could enter into shells.

All living things take in ¹⁴C until their death, after which ¹⁴C decays through converting a neutron to a proton, a beta particle (electron), and an antineutrino, forming stable ¹⁴N. Hence, ¹⁴C in dead organisms or their remains diminishes with a t_1/2 = 5730 years in a process unaffected by temperature or pressure. ¹⁴C dating can be used in samples dating back about 55,000 years, a timespan representing 9.6 half-lives. Only 0.13% of the original ¹⁴C would be left. Data by this method give the age of death of the organism.

Carbon-14 dating depends on the assumption that its amount in the environment is constant, but the burning of fossil fuels and detonation of nuclear weapons have altered its amount (see box below). Changes in solar activity and resulting changes in high-energy neutrons also affect the amount of carbon 14. Also, given the relatively short time frame used in ¹⁴C dating, differences in CO₂ based on its sequestration and circulation in the oceans are factors. Oceans cover 61% of the Northern Hemisphere compared to 81% of the Southern Hemisphere. Books of calibration factors can be used to control for these effects. The calibrations are based on data on tree rings, lake and ocean sediments, corals, and stalagmites, which date back to 55,000 years ago.

The decay of other "unstable" radioactive isotopes is used for dating samples and determining their age of burial:

• ³⁹Ar, an extremely rare isotope (t_1/2 =269 yr), has been used to date ice cores from the Tibetan Plateau over the last 1,300 years.
• ⁴⁰K (t_1/2 =1.25 billion yr) decays to ⁴⁰Ar (stable), so their ratios can be used to determine how much time has passed since magma solidification into rock, based on diffusion rates of the resulting stable ⁴⁰Ar.
• The ratio of ²⁶Al/¹⁰Be in buried samples is used in dating analyses. The two isotopes are rare and produced in a fixed ratio (6.75/1) when formed in surface quartz by solar radiation (much like the formation of ¹⁴C). When buried through geological processes, there is no further production of the isotope, but fortunately (for those who measure the age of burial), they decay with different half-lives (t_1/2 = 717,000 yr for ²⁶Al and t_1/2 = 1.39 million yr ¹⁰Be)
• The ratios of ²¹Ne/²⁶Al and ²¹Ne/¹⁰Be can be used. ²¹Ne is a stable isotope, and these ratios are independent of the ²⁶Al/¹⁰Be rate.
• Uranium isotopes are widely used to measure age on a long-term time scale. ²³⁸U (t_1/2 =4.45 billion yr) is converted to ²⁰⁶Pb (stable) and ²³⁵U (t_1/2 =704 million yr) to ²⁰⁷Pb (stable) by parallel decay routes, which allow for multiple types of dating measurement.

Use of stable isotopes

Most of the data and graphs of CO₂ and temperature vs time (years ago) presented in Chapter 31.1 were determined using stable isotopes that do not decay. Much of the data is based on the ratio of the stable isotope pairs of oxygen (¹⁸O/¹⁶O) or carbon (¹³C/¹²C) in buried ice or ocean sediment cores. These isotopes have also been used to infer the temperature or temperature change when targets were buried in ice cores or ocean sediments. Temperatures at the time of water deposition in the ice layers are often inferred from ¹⁸O/¹⁶O ratios in the ice layers.

Ice core ¹⁸O/¹⁶O analyses

The oceans are huge and generally homogenous reservoirs that can give clues to long-term changes in climate. Short-term climate change would have limited effects on the oceans. The ¹⁸O/¹⁶O ratios in Greenland and Antarctic ice cores have allowed dating and temperature reconstruction over geological time since the ratio is determined by the ¹⁸O/¹⁶O in the liquid oceans at the time of ice formation.

The % natural abundance of ¹⁸O (0.205 %) and ¹⁶O (99.757 %) gives a ratio of the two isotopes of 0.0021, which is so small that the exact ratio is inconvenient for routine use. Rather, a comparison of the ratios in a target sample vs a universal reference, the δ¹⁸O value, is determined using a mass spectrometer. The δ¹⁸O value is calculated by the following equation:

\[ \delta^{18} O=\left[\frac{\left(\frac{18}{16} O\right)_{\text {sample }}}{\left(\frac{18}{16} \mathrm{O}\right)_{\text {reference }}}-1\right] * 1000
\]

Similar δ values are determined for D/H ratios (δ²H) and for ¹³C/¹²C (δ¹³C)

• The reference for δ¹⁸O calculations is the Standard Mean Ocean Water (SMOW or V-SNOW)
• The δ²D reference value is also based on SMOW or V-SNOW
• The standard for the analogous δ¹³C value is the Cretaceous Peedee Belemnite (an extinct order of squid-like cephalopods with an internal cone skeleton) sample from the Peedee belemnite (PDB) formation in South Carolina, USA. This standard is no longer available so an alternative, NBS 19, a carbonate material, is used in a new V-PDB (Vienna-PDB) scale.

Note that the ratios of the standards, and likewise of the samples, are small. Also note that in the equation for δ¹⁸O, the bracketed term is multiplied by 1000.If multiplied by 100, the value for δ¹⁸O would be a percentage. Instead, it's multiplied by 1000 to convert it to permill (per mil or %o) or parts per thousand (just like percent is parts per 100). Hence 1%o is 1 part per 1000 or 0.1%.

Equations can be just collections of letters with little intuitive meaning, or they can be deconstructed by the user to make intuitive sense. To help understand this equation, which most readers have likely never encountered, given its importance in climate studies, let's look at 3 sets of conditions. If ...

• ¹⁸O is enriched in the sample compared to the reference, then (¹⁸O/¹⁶O)_sample/(¹⁸O/¹⁶O)_reference is >1, so subtracting 1 from it makes the bracketed term +, along with δ¹⁸O;
• ¹⁸O in the sample is equal to that in the reference, then (¹⁸O/¹⁶O)_sample/(¹⁸O/¹⁶O)_reference =1, so the bracketed term = 0, and δ¹⁸O = 0;
• ¹⁸O in depleted in the sample compared to the reference, then (¹⁸O/¹⁶O)_sample/(¹⁸O/¹⁶O)_reference <1, so the bracketed term is -, along with δ¹⁸O.

A sample with a higher ¹⁸O/¹⁶O ratio (enriched in heavier isotope) than the SMOW reference will have a positive (+) δ value. If the ¹⁸O/¹⁶O ratio of the substance is lower (depleted in heavier isotope) than the SMOW reference, the δ value will be negative (-). The δvalues of SMOW (O and H isotopes) and PDB (C isotopes) are zero as they are compared to themselves.

Now let's examine ice and ocean δ¹⁸O values and see how delta values are used as a proxy for temperature or change in temperature at deposition.

The ice shields come from snow from water evaporated from the oceans. Water can have multiple isotopic compositions, but from the % abundancies, the most likely ones are H₂¹⁶O and H₂¹⁸O, heavier than the light form, H₂¹⁶O. H₂¹⁶O evaporates more readily from the mid-latitudes of the oceans, and when it reaches the poles, condenses to form snow and eventually ice enriched in H₂¹⁶O. Urey showed that the vapor pressure of H₂¹⁸O is about 1% less than that of H₂¹⁶O between 46.35°C and 11.25°C.

In addition, the H₂¹⁸O that evaporates at lower latitudes is more likely to condense and be removed in rain, leaving the southern oceans enriched in H₂¹⁸O. This effect is quite significant as water, in the form of ice, in Greenland and Antarctica has about 5% less H₂¹⁸​​​​​O than water at 20⁰C from midlatitudes, making the δ¹⁸O for water a proxy for temperature but even better as measures of ice volume and from that sea levels.

Now consider the Ice Ages, when oceans were enriched in H₂¹⁸O. On glacial melting of H₂¹⁶O-enriched ice, with the melting flowing into the oceans, the H₂¹⁸O would get diluted with H₂¹⁶O. At the same time, the salinity of the sea would decrease since ice condenses without ocean salts. These differences in H₂¹⁸O/H₂¹⁶O values in ice core samples are converted to δ¹⁸O values, which can be positive or negative, as shown in Figure \(\PageIndex{10}\).

Figure \(\PageIndex{10}\): Typical δ¹⁸O values (in permil). Andreas Schmittner. https://eng.libretexts.org/Bookshelv...A_Paleoclimate

Surface ocean water has δ¹⁸O values of around zero. Due to fractionation during evaporation, less heavy isotopes make it into the air, which leads to negative delta values of around -10 ‰ for the evaporated water vapor. Condensation prefers the heavy isotopes, as described above. In this example, the first precipitation thus has a δ¹⁸O value of about -2 ‰,(more positive than the first vapor). The remaining water vapor will be further depleted in ¹⁸O relative to ¹⁶O, and its δ¹⁸O value will become more negative (-20 ‰). Any subsequent precipitation event further depletes ¹⁸O. This process is known as Rayleigh distillation and leads to very low δ¹⁸O values of less than -30 ‰ for snow falling onto ice sheets. Thus, ice has very negative δ¹⁸O of between -30 and -55 ‰. Deep ocean values today are about +3 to +4 ‰. During the last glacial maximum, as more water was locked up in ice sheets, the remaining ocean water became heavier in δ¹⁸O by about 2 ‰. Foraminifera build their calcium carbonate (CaCO₃) shells using the surrounding seawater. Thus, they incorporate the oxygen isotopic composition of the water into their shells, which are then preserved in the sediments and can be measured in the lab. Bralower and David Bice. https://www.e-education.psu.edu/earth103/node/5. Creative Commons Attribution-NonCommercial-ShareAlike 4.0 International License(link is external).

In summary, in cold conditions, Greenland and Antarctic ice is enriched in ¹⁶O since H₂¹⁶O preferentially evaporates, condenses, and freezes into ice at the poles. In addition, as we will see below, deep sea Foraminifera shells contain more ¹⁸O in shells since water is enriched H₂¹⁸O in cold conditions, since less evaporates. Hence δ¹⁸O becomes more positive.

Ocean sediment core ¹⁸O/¹⁶O analyses

In the last section, we looked at δ¹⁸O values in ice core layers and their use as proxies for land and ocean temperatures and, more fundamentally, as ice volume measures, affecting sea levels and ocean salinity. It's amazing what we can surmise about past climate based on the fact that H₂¹⁶O evaporates more readily than H₂¹⁸O and that H₂¹⁸O that did evaporate condenses at low and mid-latitudes more readily than H₂¹⁶O. These two factors lead to the enrichment of H₂¹⁶O in polar ice and H₂¹⁸O in low and mid-latitude ocean water. Remember that evaporation and condensation are physical reactions involving a state change.

Yet ice cores go back only so far in geological time. To go back further, scientists analyze shells buried in ocean sediments. More specifically, they analyze the isotopic ratio of ¹⁸O/¹⁶O in calcium carbonate (calcite, CaCO₃) in buried shells of organisms like Foraminifera. (Calcite is a stable anhydrous form of CaCO₃, but under high pressure, it can form different calcite phases.) Now, we are dealing with a new atom, C, in the carbonates. Hence, the interpretation of δ18O of buried carbonate depends on many factors compared to the δ¹⁸O values of solid and liquid water. It must include these factors

• the chemical reactions of inorganic carbonate formation (precipitation) and dissolution (compared to the physical reactions of water evaporation and condensation)
• the biological formation of CaCO₃ in shells
• an understanding of the carbon cycle
• the temperature at which the CaCO₃ was deposited
• the salinity at deposition as the formation of CaCO₃ from its ions as the overall ionic strength of the medium in which CaCO₃ formed would influence the thermodynamics and kinetics of how the separate ions approach each to form the solid
• equilibrium and kinetic controls of the precipitation reactions.

To truly understand biochemistry, we must include both biological and chemical aspects. Biochemistry requires synthesizing knowledge from many disciplines, including introductory chemistry (where precipitation reactions were likely covered for the first time) and analytic chemistry (which delves more deeply into precipitation reactions). Hence we don't apologize for bringing back your previous knowledge of precipitation reactions. At the same time, our description of the use of δ¹⁸O values in buried ocean sediments is very simplified.

We must consider two reactions to understand δ¹⁸O values in calcite Foraminifera shells. The first describes how ¹⁸O gets into CO₃^2- in the first place. The second describes the seemingly simple reactions to form CaCO₃ from its ions.

Summary

This chapter delves into the sophisticated methods used to reconstruct past atmospheric CO₂ levels and global temperatures over millions of years using isotope analyses. It highlights how scientists extract and interpret data from ice cores and ocean sediment cores—key records that have preserved information in the form of isotope ratios of oxygen (δ¹⁸O) and carbon (δ¹³C). The chapter emphasizes the following points:

• Isotope Fundamentals and Their Significance:
  The study begins by outlining why isotope effects are crucial not only for understanding biochemical structures and reactions but also for decoding Earth’s climate history. The partitioning of isotopes into water and biomolecules involves both equilibrium and non-equilibrium processes that mirror many biochemical pathways.

• Proxies for Past Climate Reconstruction:
  Since direct measurements of temperature and CO₂ are available only from recent centuries, scientists rely on “proxies” such as tree rings, coral growth bands, and the shells of marine organisms (e.g., foraminifera, diatoms). These proxies, calibrated against modern observations, allow researchers to infer past temperature variations and atmospheric CO₂ levels extending back millions to billions of years.

• Isotope Analyses in Ice and Sediment Cores:
  The chapter explains how isotopic ratios—especially ¹⁸O/¹⁶O—in ice cores from Antarctica and Greenland serve as indicators of past temperature changes and ice volumes, while similar analyses of carbonate shells in marine sediments provide complementary data on ocean temperatures and global climate trends.

• Chemical Reactions and Fractionation Processes:
  Detailed discussions focus on how isotopic fractionation occurs during the evaporation and condensation of water and the precipitation of calcium carbonate (CaCO₃) in foraminifera shells. The theoretical framework, including the Kim and O’Neil calibration equations, illustrates how temperature influences isotope partitioning and thus the δ¹⁸O values measured in climate proxies.

• Interdisciplinary Integration:
  Finally, the chapter underscores the importance of integrating knowledge from chemistry, biology, and geology. By linking the study of isotope effects in biochemical systems to their application in climate science, students gain a comprehensive understanding of how ancient climate records are constructed, validated, and used to inform models of future climate change.

Overall, this chapter provides junior and senior biochemistry majors with both the theoretical background and practical applications of isotope analysis in reconstructing Earth’s climatic history, demonstrating how rigorous scientific methods underpin our understanding of global climate dynamics.