Molecular fossils from phytoplankton reveal secular Pco₂ trend over the Phanerozoic

Fossil remains from algae are used to reconstruct 500 million years of atmospheric carbon dioxide concentrations.

Abstract

Past changes in the atmospheric concentration of carbon dioxide (Pco₂) have had a major impact on earth system dynamics; yet, reconstructing secular trends of past Pco₂ remains a prevalent challenge in paleoclimate studies. The current long-term Pco₂ reconstructions rely largely on the compilation of many different proxies, often with discrepancies among proxies, particularly for periods older than 100 million years (Ma). Here, we reconstructed Phanerozoic Pco₂ from a single proxy: the stable carbon isotopic fractionation associated with photosynthesis (Ɛ_p) that increases as Pco₂ increases. This concept has been widely applied to alkenones, but here, we expand this concept both spatially and temporally by applying it to all marine phytoplankton via a diagenetic product of chlorophyll, phytane. We obtained data from 306 marine sediments and oils, which showed that Ɛ_p ranges from 11 to 24‰, agreeing with the observed range of maximum fractionation of Rubisco (i.e., 25 to 28‰). The observed secular Pco₂ trend derived from phytane-based Ɛ_p mirrors the available compilations of Pco₂ over the past 420 Ma, except for two periods in which our higher estimates agree with the warm climate during those time periods. Our record currently provides the longest secular trend in Pco₂ based on a single marine proxy, covering the past 500 Ma of Earth history.

INTRODUCTION

Carbon dioxide shapes climate, breathes life into the biosphere, and turns the cogs of the carbon cycle both in the present and in the past. The past atmospheric concentrations of carbon dioxide (expressed in partial pressure; Pco₂) are reconstructed from indirect measurements (i.e., proxies) such as stomatal densities and indices in plant fossils, the boron isotopic composition of marine carbonate, and the stable carbon isotopic composition (δ¹³C) of marine phytoplankton, paleosols, and liverworts (1). Each proxy has strengths and limitations, such as its time span of application, associated estimation error, and sensitivity to specific Pco₂ levels (2). The reconstruction of secular trends of Pco₂ over long time scales [>10 million years (Ma) ago] often relies on compiling many different proxies to generate a continuous record (1). Thus, a single well-constrained proxy that spans the Phanerozoic may strengthen and support our understanding of Pco₂.

The stable carbon isotopic fractionation associated with oxygenic photosynthesis (Ɛ_p) is a proxy that has the potential to span the Phanerozoic. Isotopic fractionation occurs when the CO₂-fixing enzyme Rubisco (ribulose 1,5-biphosphate carboxylase oxygenase) favors ¹²C over ¹³C during inorganic carbon fixation, making the photosynthates’ isotopic composition (δ¹³C) depleted in ¹³C compared to its surrounding environmental CO₂ (3). Higher CO₂ concentrations lead to greater fractionation and vice versa, resulting in a dynamic δ¹³C of photoautotrophic biomass (4, 5). This concept is reverse engineered to reconstruct past Pco₂ by calculating Ɛ_p from the δ¹³C of organic matter (OM) derived from photoautotrophic biomass and the δ¹³C of CO₂ derived from fossilized carbonates (e.g., planktonic foraminifera) (6).

Ɛ_p has been extensively tested as a Pco₂ proxy since it was first estimated using the δ¹³C of geoporphyrins (7) and later using the δ¹³C of bulk OM (8). In subsequent studies, factors that influence Ɛ_p other than CO₂ concentrations have been explored in laboratory cultures [e.g., growth rate (9) and cell size (10)] and environmental conditions, such as seasonality, light, and temperature (11). In addition, brought to the forefront in more recent studies, alkenones (and theoretically other phytoplankton) may underestimate Pco₂ due to other factors such as cell size and carbon acquisition strategies (12–14). The impact of some factors remains difficult to constrain, such as the assumption that the primary source of carbon is passively diffused CO_2[aq] into the cell; under low CO₂ conditions, many phytoplankton implement active uptake of bicarbonate (15), a potential concern given the substantial δ¹³C difference between bicarbonate (0‰) and CO₂ (−8‰) (16) and even further complicated by active uptake elevating CO₂ at the site of carboxylation.

The δ¹³C of total organic carbon (TOC) to calculate Ɛ_p, in principle, provides a long-term record for Pco₂ (8). Using TOC does raise concerns regarding isotopic heterogeneity in different organisms due to kinetic isotope effects and Rayleigh distillation effects with branching points in biosynthetic pathways, leading to distinct δ¹³C values for carbohydrates, proteins, and lipids (17). These δ¹³C differences among biosynthetic products can be further influenced by diagenetic conditions, such as carbohydrate sulfurization (18), and mixing with terrestrial OM. Abating concerns of using TOC, compound-specific isotope analysis is used on shorter time scales, primarily relying on alkenone biomarkers, the long-chain unsaturated methyl and ethyl n-ketones produced by a select group of Haptophytes. However, Ɛ_p of alkenones only reconstructs Pco₂ during the evolutionary history of alkenone-producing Haptophytes, which are not common in the geologic record until the mid-Miocene (19).

To extend the Pco₂ reconstruction over the Phanerozoic, we estimated Ɛ_p here using the general phytoplanktonic molecular fossil phytane. Phytane is derived from chlorophyll-a, the omnipresent photoautotrophic pigment that absorbs and transfers light into chemical energy during oxygenic photosynthesis and that has been present for at least the past 2.15 billion years (Ga) (20). Phytane has been found in similarly ancient rocks and petroleum (21). Furthermore, all photosynthetic phytoplankton will contribute to this general biomarker, thereby averaging the Ɛ_p of the phytoplankton community at the time of synthesis. The Ɛ_p calculated from phytane has been previously explored as a proxy for Pco₂ at selected sites during specific time periods (22–25) and has been shown to mimic Pco₂ trends. Here, we explore its potential for reconstructing secular trends of Pco₂ over the Phanerozoic.

RESULTS

We generated δ¹³C values of phytane (δ¹³C_phytane) from 41 oils and 29 sediments. Furthermore, we compiled δ¹³C_phytane values from the literature. New and compiled data yielded 308 data points in total (table S1).

Only marine sediments and oils were used for our compilation to constrain the δ¹³C_phytane to marine phytoplankton in a more stable and homogenous environment, avoiding the potential decoupling of Pco₂ that may occur in local carbon cycles of terrestrial and lacustrine settings. By using only marine settings, this also excludes the additional confounding influence of C₃ and C₄ higher plants; chlorophyll breaks down relatively quickly, eliminating effective transport of terrestrial phytol to the ocean. Immature oils lacking signs of biodegradation were selected on the basis of the confidence in source rock identification to constrain age. Furthermore, these oils were selected on the basis of the lack of terrestrial biomarkers (e.g., oleanane, taraxastane, and bicadinanes) and the lack of local environmental irregularities (e.g., high salinity) to minimize spurious influences on the overall baseline signal for Pco₂ (for more details, see Supplementary Text). To attain the general baseline trend for the δ¹³C_phytane from marine phytoplankton over the Phanerozoic, short-term isotope anomalies were excluded [e.g., carbon isotope excursion events (CIEs) with isotopic spikes of ≥2‰ in less than 100 thousand years (ka)] such as the negative CIE of the Paleocene/Eocene boundary (26). Data before and after CIEs (when the excursion has a clear end point) are included in this compilation.

In our dataset, most δ¹³C_phytane is from extractable free phytane. Sulfur-bound phytane (i.e., phytane released from sulfur-bound moieties present in sediments that were deposited in anoxic environments) is also included. Sulfur-bound phytane is different than free phytane in that during early diagenesis, inorganic reduced sulfur species selectively react with labile functionalized lipids such as phytol or phytadienes (27). That is, sulfur-bound phytane is an excellent addition to this record: It may more accurately reflect the δ¹³C of the original phytol, whereas free phytane may have small influences by fluctuating inputs of terrestrial OM or archaeal-derived ether lipids (25, 28).

Our compilation shows that over the Phanerozoic, values for the δ¹³C_phytane range from −34.7 to −23.2‰ (Fig. 1). During the Late Ordovician (455 to 450 Ma), there is a marked negative shift from −28.3 to −34.2‰, followed by a data-scarce Silurian. A gradual positive trend during the Devonian was observed from −33.9‰ at ca. 380 Ma to −28.7‰ at ca. 355 Ma. The Carboniferous into the Early Permian lacks substantial data from which to describe a trend. There is a large decrease from the Permian through the Triassic, from −26.4‰ at ca. 261 Ma to −33.2‰ at ca. 242 Ma. Then, a smaller increase in the Jurassic δ¹³C_phytane, fluctuating between ca. −33 and −30‰, is observed through the Cretaceous. A decrease and a rapid increase are observed in the Late Cretaceous, from −33.0 to −26.8‰ between ca. 98 and 93 Ma. The Paleogene shows a similar decrease, followed by an increase from −34.7 to −32.6‰. There is a data gap between 52 and 30 Ma, after which the overall trend continues positive from −33.0 to −25.3‰ at 0.1 Ma, the most positive value in the record of −23.2‰ at 14 Ma.

Fig. 1. δ¹³C_phytane.

Phanerozoic compilation of the δ¹³C_phytane from literature (pink) and data from this study (blue), and from sediment (square) and oil (circle). Age uncertainties are shown in the horizontal error bars.

DISCUSSION

Phytane-derived Ɛ_p

To calculate Ɛ_p, the δ¹³C of the photosynthetic biomass (δ_p) and the δ¹³C of dissolved CO₂ (δ_d) have to be estimated. δ_p is derived from the δ¹³C_phytane, correcting for the isotopic offset between phytol and biomass. The latter factor was estimated by compiling culture studies from 22 phytoplankton species, yielding an average of 3.3 ± 1.3‰ SD (fig. S1 and Supplementary Text). δ_d is estimated from δ¹³C of carbonate, correcting for the carbon isotopic fractionation between dissolved CO₂ with respect to HCO₃⁻ (16). Where available (dataset S1), the δ¹³C of carbonate is derived from planktonic foraminifera at the same (or nearby) site as the δ¹³C_phytane. Where unavailable, the average δ¹³C of carbonate is obtained from the global compiled average of δ¹³C of marine planktonic foraminiferal carbonate at the time of deposition (8, 29). Uncertainty for marine carbonate was assigned ±0.4‰ with uniform distribution. The correction for the isotopic fractionation between dissolved CO₂ with respect to HCO₃⁻ requires sea surface temperature (SST). This information was obtained from SST proxies (preferably δ¹⁸O from planktonic foraminifera, but otherwise from other proxies such as U^k′₃₇ or TEX₈₆) measured from each site or nearby site (dataset S1) and assigned a ±4°C SD of uncertainty. Where SST data are unavailable, temperature was estimated by adjusting the modern site for its paleolatitude (using www.paleolatitude.org), finding the SST at that location (e.g., seatemperature.org), and then correcting the present-day SST for global temporal SST anomalies [i.e., 0 to 56 Ma (30, 31) and 65 to 455 Ma (32)]. For further details on the calculations and uncertainty in each parameter on calculated Ɛ_p, see Supplementary Text.

Figure 2 shows that calculated Ɛ_p ranges from ca. 11 to 24‰. The vertical error bars indicate Monte Carlo simulations of uncertainty to 1 SD (68%), the culmination of the aforementioned uncertainties within each calculation parameter. The calculated Ɛ_p shows similar trends to the δ¹³C_phytane in Fig. 1 (side-by-side trends in fig. S2) due to the relatively minor variations in the estimated δ¹³C of dissolved CO₂. In this Phanerozoic record, Ɛ_p does not surpass 25‰. This observation matches the theoretical assumption (33) and culture-based observations (9, 34–36) that maximum fractionation (Ɛ_f) for phytoplankton is 25 to 28‰. Because our Ɛ_p is derived from a common phytoplankton biomarker, this 25‰ limit suggests that Ɛ_f is relatively similar among the major taxa. Furthermore, this limit suggests that Ɛ_f has not notably changed over the course of the Phanerozoic, despite the fact that Ɛ_f of Rubisco, when measured in vitro, has found to be substantially lower (e.g., 11‰ in Emiliania huxleyi) (37). Young et al. (38) show the positive selection of the chloroplast gene that encodes large Rubisco subunits appearing in the evolutionary lineage of ecologically important species (e.g., Chromista, Haptophyta, and Bacillariophyta), likely due to environmental stressors (i.e., during periods of marked Pco₂ declines). Considering that our observed Ɛ_p does not surpass 25‰ over the Phanerozoic, these evolutionary changes to Rubisco may not have made noticeably large changes to Ɛ_f.

Fig. 2. Ɛ_p calculated from phytane.

Phanerozoic Ɛ_p calculated from the δ¹³C_phytane and δ¹³C of dissolved CO₂ estimated from δ¹³C of foraminifera from literature (pink) and data from this study (blue), and from sediment (square) and oil (circle). Horizontal error bars indicate dating uncertainty in sample age. Vertical error bars indicate 1 SD (68%) uncertainty in Ɛ_p estimation based on Monte Carlo simulations, culminating the uncertainty in δ¹³C of the photosynthetic biomass (based on uncertainty in δ¹³C of phytane ± 0.5‰ uniform distribution and the uncertainty in offset between biomass and phytane of 1.3‰ SD) and the δ¹³C of dissolved CO₂ (based on uncertainty in δ¹³C of planktonic foraminifera ± 0.4‰ uniform distribution and uncertainty in SST ± 4°C SD).

Estimates of Pco₂ based on phytane-derived Ɛ_p

To estimate the dissolved carbon dioxide (CO_2[aq]) from Ɛ_p, we use

$${\text{CO}}_{2[\text{aq}]}=\mathit{b}/({\mathrm{Ɛ}}_{\mathrm{f}}-{\mathrm{Ɛ}}_{\mathrm{p}})$$ (1)

a relationship developed by Hayes (17) and Francois et al. (39) and that is a modification of the relationship developed for higher plants from Farquhar et al. (40). This concept has been successfully tested in laboratory cultures for CO_2[aq] ranging over 0.4 to 79 μmol kg⁻¹, covering CO₂ concentrations lower than the glacial cycles to CO₂ much higher than inferred from the past (10, 36, 41).

The term b accounts for all species-specific factors that may influence isotopic fractionation, in particular cell carbon allocation and bicarbonate uptake, as well as cell geometry and growth rate (9), and influencers of growth rate such as nutrient availability (e.g., b was found to be empirically related to phosphate concentrations) (42). The factor b has almost exclusively been studied in laboratory cultures of Haptophyte algae via alkenones, a relationship then extended into the modern environment (42). In marine surface sediments and suspended matter containing alkenones, b ranges from approximately 70 to 240‰ kg μM⁻¹ with a mean of 165 ± 53 (42). Given that phytane is a general biomarker averaging the entire phytoplankton community, as opposed to the select group of Haptophytes for alkenones, we calculated b from the δ¹³C of total OM in diverse modern marine surface sediments (Supplementary Text, table S2, and references therein). Over these 19 study sites, the average for b is 168 ± 43‰ kg μM⁻¹, consistent with the alkenone studies and with the b value used in previous phytane-based Pco₂ estimations (22, 23). A mean value of 170‰ kg μM⁻¹ with an assigned SD of ±60 is used throughout the record. Sensitivity plots (fig. S3A) show that a 1% change in b results in a 1% change in Pco₂ estimation. For details on these calculations and uncertainty estimations, please see Supplementary Text.

Ɛ_f is the maximum isotopic fractionation associated with photosynthetic carbon fixation, generally ranging from 25 to 28‰ for algae in modern oceans and laboratory experiments (43, 44). Given that phytane is a general phytoplankton biomarker, the exact percentages of each species in the phytoplankton composition contributing to the phytane pool are needed to estimate the value Ɛ_f, something that cannot be practically achieved for ancient sediments. Thus, we use the average of the laboratory culture Ɛ_f range (26.5 ± 1.5‰ uniform distribution) for the entire phytane-based reconstruction of Pco₂. Sensitivity tests are conducted in Supplementary Text and shown in fig. S3B.

To estimate the atmospheric concentration of carbon dioxide from the CO_2[aq], we used

$$\mathit{P}{\text{CO}}_2=[{\text{CO}}_{2(\text{aq})}]/{\mathit{K}}_0$$ (2)

based on Henry’s law, where the solubility constant K₀, expressed in M/atm, is

$$\text{ln} {\mathit{K}}_0={\mathit{A}}_1+{\mathit{A}}_2(100/\mathit{T})+{\mathit{A}}_3\text{ln}(\mathit{T}/100)+\mathit{S}‰[{\mathit{B}}_1+{\mathit{B}}_2(\mathit{T}/100)+{\mathit{B}}_3{(\mathit{T}/100)}^2]$$ (3)

where A and B are constants, T is temperature in Kelvin, and S‰ is salinity in ‰ (45). The constants used here are A_1–3 (−58.0931, 90.5069, and 22.2940) and B_1–3 (0.02777, −0.02589, and 0.00506), respectively (45). Temperatures are obtained as described above. Salinity is estimated to be 34‰ and assigned a ±2‰ SD uncertainty. Figure 3 shows the consideration of these factors in the error bars of these Pco₂ estimations. The vertical error bars show 1 SD (68%) uncertainty in Pco₂ estimation based on Monte Carlo simulations, culminating the uncertainty in b (±60‰ kg μM⁻¹ SD), Ɛ_f (±1.5‰), and Ɛ_p (combined uncertainties of δ¹³C of phytane ± 0.5‰ uniform distribution, the offset between biomass and phytane of ±1.3‰ SD, δ¹³C of planktonic foraminifera ± 0.4‰ uniform distribution, and SST ± 4°C SD). The impact of uncertainties in these parameters on the final estimated Pco₂ is discussed in Supplementary Text and shown in fig. S3C.

Fig. 3. Phanerozoic Pco₂ from phytane.

Estimated Phanerozoic Pco₂ (on a log scale) from literature (pink) and data from this study (blue), and from sediment (square) and oil (circle). Horizontal error bars indicate uncertainty in age. Vertical error bars indicate 1 SD (68%) uncertainty in Pco₂ estimation based on Monte Carlo simulations, culminating the uncertainty in b (±60‰ kg μM⁻¹ SD), Ɛ_f (±1.5‰ uniform distribution), and Ɛ_p (combined uncertainty in δ¹³C of phytane ± 0.5‰ uniform distribution, the offset between biomass and phytane ± 1.3‰ SD, δ¹³C of planktonic foraminifera ± 0.4‰ uniform distribution, and SST ± 4°C SD). Plotted for comparison, the Foster et al. compilation shows the Monte Carlo resampling and LOESS fit of ca. 1500 data points from the five most robust Pco₂ proxies: δ¹³C of long-chain alkenones, δ¹¹B of marine carbonate, δ¹³C of paleosols, stomatal densities and indices in plants, and δ¹³C of liverworts. Sixty-eight percent and 95% confidence intervals are shown in gray and light gray, respectively. The light blue bars represent glacial paleolatitude, as determined by the literature compilation of glaciogenic detritus (46).

The resulting Pco₂ values based on δ¹³C_phytane range from ca. 250 to 1700 μatm (Fig. 3). The estimated Pco₂ shows similar trends as δ¹³C_phytane and Ɛ_p; side-by-side trends in fig. S2 show the similarity of these three different trend lines over the Phanerozoic. For further context, we included the glaciation paleolatitude as determined by glaciogenic detritus compiled by Cather et al. (46) as an indicator of climate (Fig. 3). Last, Fig. 3 includes context for the phytane record by incorporating the compilation of Foster et al. (1), which averages the five most robust Pco₂ proxies in current literature: δ¹³C of long-chain alkenones, δ¹¹B of marine carbonate, δ¹³C of paleosols, stomatal densities and indices in plants, and δ¹³C of liverworts. A comparison between the phytane-based record and the Foster et al. compilation is also shown by time frame (Neogene, Paleogene, Cretaceous, and Phanerozoic) in fig. S4. The phytane-based record here contains ca. 310 estimations, fewer than the ca. 1500 data in the five-proxy Foster et al. compilation, although it does extend more than 50 Ma beyond the current record and has the potential to extend further.

The phytane-based proxy and the Foster et al. compilation show very similar values throughout the entire Phanerozoic. During the Late Ordovician (ca. 460 to 440 Ma), the phytane-based record jumps from ca. 450 to 700 μatm, a more marked shift than seen in the δ¹³C_phytane and Ɛ_p trends mostly due to the low estimates for temperatures in the Ordovician (ca. 10°C) relative to the estimates for the Devonian (ca. 23°C). The glaciation paleolatitude for this time interval extends to 60° (46), suggesting a cold climate, which agrees with the relatively low Pco₂. From the Devonian into the Early Carboniferous, Pco₂ drops from 1400 to 300 μatm, amplified from the trend seen in phytane-based Ɛ_p but a trend that is similar to the Foster et al. estimations. This significant drop in Pco₂ is further supported by the glaciation paleolatitude, where it significantly drops to 60° at the start of the Carboniferous and moves up to 30° by the end of the Carboniferous into the early Permian (46). Then, Pco₂ increases from the Late Permian at 450 μatm through the Triassic at 1600 μatm. The Jurassic exhibits a gradual decrease from 1000 μatm during the Toarcian to 600 μatm in the Tithonian. From the Late Jurassic to the mid-Cretaceous, there is a gradual increase to 1300 μatm. The Cenomanian starts at 1500 μatm, the highest Pco₂ values for the δ¹³C_phytane-based Phanerozoic record, which then rapidly drops to 600 μatm from ca. 98 to 85 Ma. The high values during the Cenomanian are much higher than those based on the Foster et al. compilation. This may also be attributed to the important role that temperature has when converting raw δ¹³C values from biomarkers to Pco₂ (see Supplementary Text and fig. S3). However, considering that this period is marked with extremely high SSTs (47), the high phytane-based Pco₂ estimations may be appropriate. A second increase and a second drop in the record then occur in the early Paleogene from ca. 56 to 54 Ma, dropping from 1400 to 7500 μatm. Here, our Pco₂ estimates are much higher than those of Foster et al. Our high estimates agree with high SST records during this time (48). Last, a decrease in Pco₂ from ca. 1000 to 250 μatm is observed from the late Paleogene toward the Holocene (ca. 30 to 0.1 Ma), the lowest estimate for the Phanerozoic. This lowering of CO₂ is supported by the glaciation paleolatitude, which extended as far as 40° (46), and in agreement with the overall cooling observed in bottom water temperatures and the descent in the so-called icehouse world (49).

CONCLUSION

Our Phanerozoic Pco₂ record based on the δ¹³C_phytane is, to the best of our knowledge, one of the longest reconstructions based on a single proxy, extending the known Pco₂ record. As a spatially and temporally ubiquitous compound, phytane is one of the most abundantly available phytoplanktonic biomarkers suitable for Pco₂ reconstructions, more so considering that both sediments and oils can be used. Among marine-based proxies, this phytane record is the longest reconstruction for Pco₂. Phytane-based Pco₂ reconstruction yields similar estimates as compilations of Pco₂ proxies, giving the potential to yield a more robust and consistent Pco₂ record from a single biomarker.

MATERIALS AND METHODS

The isotopic composition of phytane was measured in 70 marine sediments and oils derived from marine source rocks. Marine oils were processed at Shell Global Solutions International B.V., The Netherlands. Crude oil was eluted over a AgNO₃-impregnated silica gel column using three column volumes of cyclohexane to yield saturated hydrocarbon fractions. To remove n-alkanes, the saturated fractions remained in cyclohexane when two layers of 0.5-Å molecular sieve were added to the samples and saturated overnight. The remaining branched/cyclic fractions were injected splitless on gas chromatography–flame ionization detector (GC-FID) at 35°C for 5 min, ramped to 325°C at 4°C/min for 15 min, and held isothermal for another 15 min. A silica capillary column (Ultra-1, 50 m × 0.22 mm; d_f, 0.11 μm) was used with helium as a carrier gas at a constant flow of 25 cm/s. GC–isotope ratio mass spectrometry (IRMS) was conducted using a DB-1ms column (60 m × 0.32 mm; d_f, 0.25 μm). The samples were injected at 220°C into a 70°C oven for 1 min and ramped to 250°C at a rate of 4°C/min and then to 300°C at a rate of 20°C/min for 20 min at a flow rate of 30 cm/s using helium as a carrier gas. The reference gas was normal CO₂ with a predetermined isotopic composition.

Twenty-nine marine sediments from Deep Sea Drilling Project Site 467 offshore of southern California from the Middle Miocene to Lower Pliocene (50) were processed at NIOZ Royal Netherlands Institute for Sea Research, The Netherlands. Powdered sediments (15 to 20 g) were extracted with dichloromethane (DCM):MeOH (9:1, v/v) on a Dionex 250 accelerated solvent extractor at 100°C, 7.6 × 10⁶ Pa. Extracts were eluted over Na₂SO₄ to remove excess water and then over an alumina-packed column to separate polar fractions (DCM:MeOH, 1:1, v/v). Polar fractions were desulfurized using Raney nickel (51), eluted over alumina oxide into an apolar fraction (hexane:DCM, 9:1, v/v), and hydrogenated. Desulfurized apolar fractions were injected on a GC-MS to identify the presence of phytane and on a GC-FID to determine quantity before injection on IRMS for the isotopic composition of phytane. GC-FID, GC-MS, and GC-IRMS all had a starting oven temperature of 70°C and ramped to 130°C at 20°C/min and then to 320°C for 10 min at 4°C/min. GC-IRMS was conducted using a CP-Sil 5 column (25 m × 0.32 mm; d_f, 0.12 μm) using a constant flow of He carrier gas.

SUPPLEMENTARY MATERIALS

Supplementary material for this article is available at http://advances.sciencemag.org/cgi/content/full/4/11/eaat4556/DC1

Supplementary Materials and Methods

Supplementary Text

Fig. S1. Isotopic offset between biomass and phytol.

Fig. S2. Trends from reported δ¹³C_phytane data to Pco₂.

Fig. S3. Uncertainties associated with equation parameters.

Fig. S4. Pco₂ from phytane over the Phanerozoic in time slices.

Table S1. Isotopic offset between biomass and phytol.

Table S2. Estimating b from marine OM in modern-day oceans.

Table S3. Estimating b from phytol across an equatorial Pacific Ocean transect.

Code S1. Python code used for Monte Carlo simulations to calculate uncertainty for Pco₂ estimations by considering every parameter involved in the equations.

Data S1. All data used to reconstruct Pco₂ from the δ¹³C_phytane.

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