Oxygen and Carbon Isotopes in Marine Carbonates: A Biogenic Climate Archive Built Upon Disequilibria

 1811-5209/25/0021-0112$2.50 DOI: 10.2138/gselements.21.2.112

Keywords: marine biogenic carbonates; oxygen and carbon isotopes; paleoclimate; vital effects

MARINE CALCIFYING ORGANISMS: THE PRIMARY ARCHIVE

As a water planet that hosts life, a significant proportion of Earth’s climate history, from Phanerozoic to paleoseasonal timescales, is archived in calcium carbonate (CaCO3) minerals built by various marine organisms. Over the last nearly seven decades, studies of the stable oxygen and carbon isotope compositions (denoted δ18O and δ13C, respectively, which measure the relative deviations in 18O/16O and 13C/12C ratios of a sample from an international standard) of foraminifera shells in deep-sea sediments have produced an iconic record of Earth’s climate history, as worlds of dramatically different climates in the past are recorded in grains of CaCO3 sand made by these single-celled protists (Emiliani 1955; Zachos et al. 2001; Westerhold et al. 2020).

The benthic foraminifera record over the last 66 million years is nearly continuous, revealing trends, rhythms, and aberrations in Earth’s climate, as well as a complex yet intriguing relation between climate and the carbon cycle (Fig. 1A, 1B; Zachos et al. 2001; Westerhold et al. 2020). While the benthic δ18O record primarily reflects changes in deep-sea temperature and global ice volume, the δ13C record reflects shifts in the distribution of carbon among different reservoirs as a result of tectonic, climate, ocean circulation, and ecological changes. A global composite of the benthic δ18O record shows a long-term cooling trend over the Cenozoic, as the Earth transitioned from an ice-free Eocene hothouse to a Pleistocene icehouse with regular glaciations (Fig. 1A; Zachos et al. 2001; Westerhold et al. 2020). This transition was likely accompanied by a systematic shift in the interactions between climate and the carbon cycle, as the δ18O and δ13C records switch from positive correlations in warm climates to negative correlations in cold climates (Fig. 1C; Kirtland Turner 2014; Westerhold et al. 2020).

In addition to δ18O and δ13C, advances in clumped isotope geochemistry have provided a new tool for investigating the evolution of Earth’s climate system (Ghosh et al. 2006; Meckler et al. 2022). Unlike δ18O and δ13C, which have multiple sources of variability, the clumped isotope composition of marine carbonates (denoted Δ47, which measures the preferential formation of 13C-18O bonds in carbonate groups in the mineral relative to random distribution of the rare isotopes) is almost purely a function of temperature (Ghosh et al. 2006). The Δ47 record of benthic foraminifera has not only confirmed the overall Cenozoic cooling trend indicated by δ18O, (Fig. 1D, 1E), but has helped deconvolve changes in deep ocean temperature from changes in seawater δ18O over the Cenozoic. A combination of the two records also revealed previously underappreciated factors that may affect the δ18O record, including seafloor spreading rates, ocean circulation, and seawater pH (Meckler et al. 2022). Despite lower temporal resolution, the benthic Δ47 record also shows larger temperature variability in the early Cenozoic compared with the δ18O record, suggesting a less stable climate system in a hothouse world (Meckler et al. 2022).

While foraminifera give us a nearly continuous record of climate evolution over millennial and longer timescales, other biogenic carbonates offer the advantage of interannual to sub-annual (seasonal, monthly, or even higher) temporal resolution afforded by their high growth rates, most notably coral skeletons and giant mollusk shells. For instance, the δ18O, δ13C, and Δ47 of shallow-water corals exhibit regular seasonal cycles (Fairbanks and Dodge 1979; Ghosh et al. 2006; Saenger et al. 2012). However, the δ18O and Δ47 in corals do not always agree on the magnitude of temperature fluctuations (Saenger et al. 2012), while the δ13C can vary with or against δ18O seasonally under different circumstances (Fairbanks and Dodge 1979). It is a longstanding and ongoing challenge to unravel the fundamental mechanisms underlying these complex relationships during the formation of biogenic carbonates.

VITAL EFFECTS: DISEQUILIBRIUM IN BIOGENIC CARBONATES

The link between measured isotopic changes and inferred environmental signals is based on principles of equilibrium fractionation. However, all biological processes, including those that make CaCO3 minerals (i.e., biomineralization), involve disequilibrium. Deviations from isotopic equilibrium are called “kinetic isotope effects,” while disequilibrium associated with biogenic carbonates are referred to as “vital effects” (Urey et al. 1951). When using carbonates as proxies for paleoenvironment, the hope is that these effects are either minimal or systematic enough to allow for reliable corrections to be made.

The magnitudes and patterns of vital effects have been characterized for various calcifying organisms, revealing similarities and differences that offer clues to their origin. In biogenic carbonates, more extreme deviations from equilibrium almost always result in lower (more negative) δ18O and δ13C values (Fig. 2A, 2B). If photosynthesis is involved, the δ13C of some biogenic CaCO3 can be elevated above equilibrium values, while the δ18O values are sub-equilibrium, with the exception of some coccolithophores (Fig. 2B). Interestingly, a ubiquitous phenomenon in spatially sampled biominerals is that δ18O and δ13C are linearly correlated, even in the absence of significant environmental changes, which suggests a common causal mechanism (Fig. 2A, 2B; McConnaughey 1989a; Adkins et al. 2003; Chen et al. 2018). Meanwhile, the magnitude of the observed vital effects in some organisms can be as large as the benthic δ18O and δ13C variations through the entire Cenozoic (Fig. 1C). Such variability in stable isotopic composition is difficult to explain by environmental changes alone.

Some organisms, by contrast, exhibit minimal or constant deviations from isotope equilibrium, making them more attractive targets for paleoclimate reconstructions. For example, although depleted from equilibrium, the δ18O of different foraminifera species generally show the same temperature sensitivity as inorganic CaCO3, making species-specific offset corrections an effective way to remove the vital effects (Fig. 2C; Shackleton et al. 1984; Marchitto et al. 2014). A similar observation can be made for the clumped isotope thermometer, where the Δ47 temperature sensitivities of different organisms are similar to inorganic CaCO3 (Fig. 2D; Ghosh et al. 2006; Meinicke et al. 2020). However, significant scatter exists in the empirical calibrations that limit the accuracy and precision of this thermometer in some cases (Fig. 2D).

Species-specific empirical calibrations in paleoclimate reconstruction are justified because they are effective at removing vital effects from paleoclimate records (Westerhold et al. 2020; Meckler et al. 2022). The downside of treating vital effects as a nuisance is that it draws attention away from some important research avenues: What causes the linear δ18O–δ13C correlations (Fig. 2)? What drives δ18O and δ13C offsets between organisms? Why are some organisms more reliable recorders of their environment than others? Are species-specific offsets related to certain environmental conditions, and could those have changed through geologic time? The last question is particularly relevant for the early Cenozoic climate record, as a significant portion of the data come from an extinct Nuttallides species (Fig. 1). The accuracy of our interpretation of the biogenic carbonate archive critically depends on a better mechanistic understanding of these vital effects.

INORGANIC CARBONATES: A USEFUL REFERENCE FRAME

Evaluating vital effects requires a reference frame for equilibrium and non-equilibrium (kinetic) fractionation processes. For decades, researchers have grown carbonate minerals in the laboratory under controlled conditions to establish such a benchmark. These efforts showed, among other things, that stable isotope partitioning between carbonates and water depends on whether the crystals grow slowly, under near-equilibrium conditions, or rapidly, under far-from-equilibrium conditions. Under near-equilibrium conditions, isotope partitioning depends on temperature and is independent of the reaction mechanisms and pathways of crystal growth, allowing straightforward interpretations of equilibrium stable isotope values. Such conditions, however, are difficult to achieve in experiments and are the exception more than the rule in nature (Coplen 2007; Daëron et al. 2019).

The recognition that even laboratory-grown, inorganic carbonates precipitate out of isotopic equilibrium has prompted efforts to determine how the different reactions that form carbonate minerals influence their isotopic compositions. Consider, for example, carbonate growth from a CO2-fed solution (Fig. 3A), which may be similar to some calcifying organisms. The incoming CO2 reacts with H2O and OH− to form HCO3− and CO32−, which are the anion “building blocks” of the CaCO3 mineral. The stable isotope composition of these building blocks depends not only on the isotopic composition of the various reactants (CO2, H2O, and OH−) but also on the tendency of isotopically light molecules of these compounds to react faster than their heavier counterparts, forming light isotope-enriched HCO3− and CO32−. The quantity needed to describe the degree of light isotope enrichment in the resultant dissolved species is the “kinetic fractionation factor,” which is equal to the ratio of isotope-specific (i.e., mass-dependent) rate constants for a particular reaction. Isotopes are addition ally fractionated as these building blocks transfer to and from a mineral’s surfaces, and again, there are kinetic fractionation factors attending these reactions. The kinetic isotope effects associated with the formation of HCO3− and CO32− and their attachment to mineral surfaces tend to manifest most strongly at high-pH conditions, when the hydration/hydroxylation and mineral growth reactions are less reversible.

Characterizing how each reaction step contributes to the resulting isotopic composition of CaCO3 is complicated, but significant progress has been made using a variety of approaches. For instance, to isolate the crystal growth reaction (black arrows in Fig. 3A), the enzyme carbonic anhydrase has been used to facilitate isotopic exchange between CO2 and H2O (Uchikawa and Zeebe 2012) so that the isotopic compositions of the anion building blocks are equilibrated and therefore known. Under these special conditions, the kinetic fractionation factors between calcite and HCO3−/CO32− can be determined by comparing model-derived curves to experimental determinations (Fig. 3B–3D; Watkins and Devriendt 2022 and references therein). Meanwhile, the kinetic fractionation factors of the hydration and hydroxylation reactions (blue arrows in Fig. 3A) have been adequately determined only recently through carefully designed laboratory experiments (Yumol et al. 2020) and field studies (Christensen et al. 2021) that minimize the effects of the reverse reactions. A key outcome of these efforts is a model of the equilibrium and kinetic isotope fractionations within the inorganic DIC–CaCO3 H2O system, which constitutes a powerful tool in the quest to demystify vital effects and more accurately decipher the carbonate archive.

BIOMINERALIZATION: CONTROLLED DISEQUILIBRIUM BY CALCIFYING ORGANISMS

The reference frame built from inorganic experiments requires at least two conditions to achieve equilibrium fractionation in the carbonate mineral: (1) sufficient time for isotope equilibration between DIC and water (seconds to minutes for δ13C, hours to days for δ18O in alkaline conditions; McConnaughey 1989b; Uchikawa and Zeebe 2012); (2) sufficiently slow mineral growth to allow equilibration at the mineral–fluid interface. Neither condition is likely to be met for relatively fast-growing biogenic carbonates. Marine calcifying organisms can additionally deviate from equilibrium with seawater by modifying the chemical and isotopic composition of the fluid from which the CaCO3 minerals grow, with further complications from involvement of organic carbon through photosynthesis/respiration. To understand the effects each factor imposes upon the CaCO3 mineral, a comprehensive biomineralization model is needed that takes into account both the biological calcification and physicochemical isotope fractionation processes.

After decades of in-depth study from biophysical, morpho logical, and geochemical perspectives, an integrated model for CaCO3 biomineralization has been developed that is applicable to a wide variety of calcifying organisms (Fig. 4A; Gilbert et al. 2022 and references therein). The key features of this model include: (1) Calcification takes place in a space in the organism (either intercellular or extracellular) from fluid that is derived from but differs from ambient seawater. (2) This space contains calcifying fluid (CF) from which CaCO3 precipitates, with the chemical composition of the CF modified by the organism, tending to increase its pH as well as Ca2+ and DIC concentrations to facilitate growth of the biomineral. (3) Ions used for calcification come from active pumping by ion channels on the membranes around the CF and/or paracellular seawater transport. An additional source of carbon is by diffusion of neutral CO2 molecules through membranes that can be influenced by metabolic processes. (4) CaCO3 biomineralization occurs through two major pathways, particle attachment (PA) and ion attachment (IA). PA takes place through rapid formation of amorphous calcium carbonate (ACC) nanoparticles in the highly saturated CF, which are the building blocks and precursors to the final crystalline mineral. In contrast, IA is similar to inorganic precipitation that is typically much slower than PA, and acts to fill the space left unfilled by the ACC particles.

The key question that presents itself is: How do these processes act to mechanistically explain the stable isotope deviations from equilibrium described above? The CF essentially operates as a CO2-fed solution, similar to the inorganic calcite experiments shown in Figure 3. However, it includes an “alkalinity pump” where Ca2+ (and/or other cations) enter the solution in exchange for H+ ions. This active removal of H+ can counteract the acidifying effects of CO2, leading to an increase in pH (de Nooijer et al. 2009; Venn et al. 2011). This pH increase is crucial, as it significantly impacts the stable isotope composition of the final mineral in several ways: (1) It slows down isotopic equilibration between DIC and water, promoting light isotope enrichment in HCO3– and CO32–; (2) It leads to greater supersaturation, promoting faster crystal growth that can preserve larger kinetic isotope effects; (3) It increases the concentration gradient of CO2(aq) between the calcifying fluid and cell cytoplasm, and shifts the carbon source from predominantly seawater DIC to metabolic carbon with much lower δ13C (Fig. 4B; McConnaughey 1989a; Adkins et al. 2003; Chen et al. 2018).

A quantitative version of these concepts shows promise for demystifying vital effects. Focusing on deep-sea corals, Chen et al. (2018) wrote a set of equations that includes cross-membrane CO2 transport, seawater infiltration, alkalinity pumping, and the physicochemical rules of DIC–H2O and DIC–CaCO3 isotope exchange reactions. Numerical solutions to these equations can reproduce the δ18O–δ13C correlations in coral individuals (Fig. 4B), and in so doing, reveal probable values of the pH of the CF and strength of the alkalinity pump. A key insight is that the δ18O–δ13C slope is sensitive to the activity of carbonic anhydrase in the CF as well as the δ13C and residence time of cross-membrane CO2(aq), a possible explanation for why different calcifying organisms exhibit different slopes (Fig. 2A, 2B). For example, photosynthesis incorporates 13C-depleted carbon into organic matter, which elevates the δ13C of cross-membrane CO2(aq) and decreases the δ18O δ13C slope (McConnaughey 1989a), while incorporation of respiratory CO2 can cause depletions in both δ13C and δ18O of the biomineral (Erez 1978). Meanwhile, the residence time of CO2(aq) and carbonic anhydrase activity in the CF can both influence the amount of kinetic isotope effects preserved in the CaCO3 mineral and, therefore, δ18O–δ13C slopes (Chen et al. 2018; Chauhan and Rickaby 2024).

Given the similar biomineralization processes observed in different calcifying organisms, the coral-based quantita tive model can be adapted to other calcifying organisms by adjusting key parameters, such as the strength of alkalinity pumps and the seawater turnover rate, which helps to reveal and quantify the key biological processes that cause vital effects. An open question relevant for foraminifera-based paleoclimate records is whether offsets in the absolute δ18O and δ13C values among different species are caused by organism-specific pH up-regulation or growth-rate dependent kinetic fractionation at the mineral–fluid interface (or both), and whether these effects are themselves dependent on environmental conditions such as seawater pH. Another frontier regards the growth mechanisms through PA versus IA in different organisms, and how these affect the isotopic composition of CaCO3. It will also be interesting to see how such models perform for clumped isotopes and whether they can account for the vital effects in various organisms.

OUTLOOK

Quantitative paleoclimate reconstructions were born with the first measurements of δ18O on carbonate minerals around 75 years ago. Improved sampling and measurement techniques, along with the development of new methods such as clumped isotope thermometry, continue to refine our knowledge of climate change, from seasonal to Phanerozoic timescales. The records are derived largely from the remains of once-living organisms, and so-called “vital effects” have long complicated the interpretation of stable isotope data. Although empirical corrections have been applied with great success, there is much to be gained in the coming years from simplified but useful process based models. Understanding how organisms respond to environmental changes will lead to exciting new discoveries into their adaptive strategies as well as more accurate models for past, present, and future climate scenarios.

ACKNOWLEDGMENTS

We thank Michael Bender for sharing his perspective on early drafts. We gratefully acknowledge contributions from Don DePaolo, an anonymous reviewer, and the editors of the issue, including principal editor Tom Sisson during the review process, all of which improved the quality and clarity of the manuscript considerably. SC acknowledges financial support from National Key R&D Program of China 2023YFF0806100 and NSFC grant 42103081. JW was supported by NSF-CAREER Grant no. EAR1749183.

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