Variable Isotope Fractionation Research Articles

Biogeochemistry of iron in coastal Antarctica: isotopic insights for external sources and biological uptake in the Amundsen Sea polynyas

Seasonal phytoplankton blooms in the Antarctic Amundsen Sea Polynyas are thought to be supported by an external supply of iron (Fe) from circumpolar deep waters, benthic sediments, and/or ice shelf meltwaters. However, largely due to the limited amount of Fe data reported for the Amundsen Sea Polynyas, understanding of the sources and processes that affect the biogeochemistry of Fe in this region (notably within the ice shelf system) remains limited. Here, we present the first investigation of dissolved Fe isotope distributions (δ56Fe) along the conveyer belt of waters into and through the Amundsen Sea, via the Dotson Ice Shelf, from samples collected during austral summer (2017–2018). Our dataset allows us to characterize and compare the dissolved δ56Fe signatures of incoming modified Circumpolar Deep Water (mCDW) and of sedimentary sources on the continental shelf. The range in dissolved δ56Fe (–1 to +0.1 ‰) observed in the Amundsen Sea close to the seafloor, coupled with elevated dissolved Fe concentrations (up to 1.6 nmol/L), suggests that Fe is released from shelf sediments via a combination of reductive and non-reductive processes, with non-reductive dissolution input being relatively more important (20–56 %) than reductive dissolution (4–12 %). Near the Dotson Ice Shelf, the δ56Fe in the mCDW inflow (–0.70 ‰) was lower than the mCDW outflow (–0.23 ‰), whereas any change in dissolved Fe concentrations was negligible. We speculate that this shift in dissolved δ56Fe underneath the ice shelf is driven by a combination of enhanced preservation (and addition) of lithogenic colloidal Fe(III) and/or complexation with Fe-binding ligands, together with a differential loss of Fe2+. We also found distinct δ56Fe signatures in surface waters of the polynya, with apparent preferential uptake of isotopically light Fe in a bloom dominated by diatoms leading to a relatively heavy remnant dissolved δ56Fe signature of +1.06 ‰, compared to a bloom dominated by haptophytes where more modest and variable isotope fractionation was observed. The different isotopic composition between the two regions could be related to the dominance of different species, but this remains speculative. Despite prominent biological uptake, we suggest that other factors such as rapid recycling (e.g., adsorption and regeneration), bacterial regeneration, and complexation with organic ligands, together with the supply of lithogenic particles also play important roles in setting surface dissolved δ56Fe in the Amundsen Sea Polynyas. Overall, this study provides a further understanding of the external Fe sources and the biogeochemical processes in the Amundsen Sea and thus a baseline on how changing conditions in Antarctica can affect Fe cycling in the Southern Ocean and beyond.

The metabolic sensitivity of hydrogen isotope fractionation differs between plant compounds

Hydrogen stable isotope analyses (δ2H) of plant derived organic compounds are a useful tool for ecological, environmental, and palaeoclimatological research. However, during organic compound synthesis, variable biosynthetic 2H-fractionation has been suggested to occur as a result of changes in plant carbon fluxes. So far, inference has been based on examining the δ2H patterns of plant compounds along environmental gradients, among plant species, and between plant organs. In an alternative approach, we used four plant species with four different types of mutations that cause impaired starch synthesis to assess whether variability in carbon metabolism affects the biosynthetic 2H-fractionation during cellulose, phytol, and acetogenic lipid synthesis. We found that mutants with impaired starch synthesis always had higher cellulose and phytol δ2H values compared to the wild type. By contrast, 2H-fractionation during acetogenic lipid biosynthesis generally did not show strong metabolic sensitivity. We rationalise these differences by considering the biosynthetic pathway of each compound and the likely source of the variable isotope fractionation. In different organic compounds, the sensitivity of variable biosynthetic 2H-fractionation to changes in C-metabolism depends on incorporation of specific H atoms from precursor molecules. As such, we determined that the similar increase in cellulose and phytol δ2H values as an effect of impaired starch synthesis most likely originates in triose-phosphates.

Sedimentary n-alkanes and n-alkanoic acids in a temperate bog are biased toward woody plants

Sedimentary plant waxes and their hydrogen and carbon (δ2H and δ13C) isotopes are important proxies for past hydrologic and vegetation change. However sedimentary waxes accumulate from diverse sources, integrating uncertainties from: (i) variable isotope fractionation among plant species, and (ii) unresolved processes controlling the transport of waxes from plants to sediments. We address these uncertainties by comparing the molecular and isotopic composition of n-alkanes and n-alkanoic acids in recent bog sediments with all major plant species growing in the catchment of Browns Lake Bog (BLB) in Ohio, USA. There are two distinct plant assemblages at BLB, including a forest dominated by trees and a bog shoreline composed of shrubs, woody groundcover, herbs and graminoids. n-Alkane concentrations in trees were 10–300 times higher than in shoreline plants, while n-alkanoic acid concentrations were generally lower and comparable across all species. The overall range of wax δ2H values among individual plants (77‰ for n-C29 alkane and 84‰ for n-C28 alkanoic acid) was likely driven by interspecies differences in biosynthetic δ2H fractionation as well as source water differences between forest and shoreline plants. A considerably smaller range of δ2Hwax values in the bog sediments (9‰ for n-C29 alkane and 11‰ for n-C28 alkanoic acid) suggests that sediments are either biased toward specific plants, or that signal averaging processes during or after deposition are constant. The combined δ2H and δ13C signatures of plant sources and sediments indicate a sediment bias mainly toward trees, with contributions from woody shrubs and groundcover growing in the bog shoreline. Within trees and woody shrubs, we observed δ2Hwax–δ13Cwax relationships of opposite sign for n-C29 alkane and n-C28 alkanoic acid, which we speculate may reflect contrasting seasonal timing of synthesis and plant metabolic status between compound classes. The net apparent δ2H fractionation between precipitation and wax (εapp) was approximately 30‰ larger for n-alkanes (−133‰) than for n-alkanoic acids (−103‰), both at the plant level and in sediments. These results demonstrate the sensitivity of sediments in a hydrologically closed basin to woody plants growing in the associated catchment and can guide εapp estimates for sedimentary records from similar depositional settings.

Chemical nitrogen fractionation in dense molecular clouds

Nitrogen-bearing molecules display variable isotopic fractionation levels in different astronomical environments such as in the interstellar medium or in the Solar system. Models of interstellar chemistry are unable to induce nitrogen fraction in cold molecular clouds as exchange reactions for 15N are mostly inefficient. Here, we developed a new gas–grain model for nitrogen fractionation including a thorough search for new nitrogen fractionation reactions and a realistic description of atom depletion on to interstellar dust particles. We show that, while dense molecular cloud gas-phase chemistry alone leads to very low fractionation, 14N atoms are preferentially depleted from the gas phase due to a mass-dependent grain surface sticking rate for atomic nitrogen. However, assuming an elementary 14N/15N ratio of 441 (equal to the solar wind value), our model leads to only low 15N enrichment for all N-containing species synthesized in the gas phase with predicted 14N/15N ratios in the range 360–400. Higher enrichment levels can neither be explained by this mechanism, nor through chemistry, with two possible explanations: (i) The elementary 14N/15N ratio in the local ISM is smaller, as suggested by the recent work of Romano et al., with a hypothetic 15NNH+ and 15NNH+ depletion due to variation of the electronic recombination rate constant variation with the isotopes and (ii) N2 photodissociation leads to variable nitrogen fractionation in diffuse molecular clouds where photons play an important role, which is conserved during dense molecular cloud formation as suggested by the work of Furuya & Aikawa.

Variable apparent hydrogen isotopic fractionation between sedimentary n-alkanes and precipitation on the Tibetan Plateau

The hydrogen isotope ratio values of sedimentary leaf wax n-alkanes (δDwax) have been considered a reliable proxy for precipitation isotope ratios (δDP) across various transects on a global basis. Records of δDwax have been used to infer past changes in hydrology, climate and altimetry, assuming nearly constant apparent hydrogen isotopic fractionation (εwax-P) between sedimentary leaf wax components and precipitation. Here, we present the results of an analysis of δDwax values from the surface sediments of 22 lakes across a precipitation and vegetation gradient on the Tibetan Plateau. The δDwax values correlated significantly with mean annual precipitation δD (r 0.90) and summer precipitation δD (r 0.91), suggesting that δDwax can record δDP in very dry regions. However, εwax-P varies with annual precipitation as well as with relative humidity. Calculation using Craig-Gordon isotope models showed that the variable εwax-P values likely resulted from intensive soil evaporation and plant evapotranspiration in this relatively arid region. This implies that variable isotopic fractionation should be taken into consideration when reconstructing paleohydrology, paleoclimate and paleoaltimetry.

Rate-Dependent Carbon and Nitrogen Kinetic Isotope Fractionation in Hydrolysis of Isoproturon

Stable isotope fractionation permits quantifying contaminant degradation in the field when the transformation reaction is associated with a consistent isotope enrichment factor epsilon. When interpreted in conjunction with dual isotope plots, isotope fractionation is also particularly useful for elucidating reaction mechanisms. To assess the consistency of epsilon and dual isotope slopes in a two-step reaction, we investigated the abiotic hydrolysis of the herbicide isoproturon (3-(4-isopropylphenyl)-1,1-dimethylurea) using a fragmentation method that allows measuring isotope ratios in different parts of the molecule. Carbon and nitrogen position-specific isotope fractionation, as well as slopes in dual isotope plots, varied linearly with rate constants k(obs) depending on the presence of buffers that mediate the initial zwitterion formation. The correlation can be explained by two consecutive reaction steps (zwitterion formation followed by dimethylamine elimination) each of which has a different kinetic isotope effect and may be rate-limiting. Intrinsic isotope effects for both steps, extracted from our kinetic data using a novel theoretical treatment, agree well with values computed from density functional calculations. Our study therefore demonstrates that more variable isotope fractionation may be observed in simple chemical reactions than commonly thought, but that consistent epsilon or dual isotope slopes may nonetheless be encountered in certain molecular fragments.

Determining high precision, in situ, oxygen isotope ratios with a SHRIMP II: Analyses of MPI-DING silicate-glass reference materials and zircon from contrasting granites

The development of new techniques and instrumentation on the ANU SHRIMP II ion microprobe has made it possible to measure the oxygen isotope ratios of insulating and conducting phases (e.g. silicates, carbonates, phosphates and oxides) on a 25 µm scale with better than 0.4‰ precision and accuracy at 95% confidence. Instrumentation changes include the installation of a multiple collector, charge neutralization using an oblique-incidence low-energy electron gun, and the addition of Helmholtz coils to counter mass dispersion by the Earth's magnetic field. A redesign of sample mounts and mount holders has effectively eliminated differences in variable isotope fractionation across the mount surface during analysis. Techniques have been developed to minimize the effect of electron-induced secondary ionization of oxygen. During a 6-minute analysis involving 100–140 s of data collection, δ 18O values can be measured on one 25 µm spot with an internal precision of better than 0.2‰ (2 standard errors). Analyses of MPI-DING silicate-glass reference material demonstrate that the external reproducibility of single spots can be better than 0.4‰ at 95% confidence, and that for matrix-matched samples and reference material, accuracy is commensurate with precision. MPI-DING glasses are acceptable ion microprobe reference materials for oxygen isotope measurements of glasses, although KL2-G is possibly heterogeneous. Zircon reference materials TEMORA 2 and FC1 appear to be acceptable as preliminary oxygen isotope reference materials. SHRIMP II analyses of FC1 indicate that it has a δ 18O value of 5.4‰ (VSMOW). Analyses of zircon oxygen isotopic compositions from a gabbro, a tonalite and a granodiorite from southeastern Australia are presented. Zircon from the gabbro has a δ 18O value of 5.6‰, the tonalite has an I-type affinity and slightly heterogeneous δ 18O values around 6.6‰, and the granodiorite has an S-type affinity and a range of igneous, melt precipitated zircon δ 18O values between 8.2 and 10.2‰. These results suggest that the gabbro is mantle-derived and slightly contaminated with crustal material, and that the I-type granodiorite has evolved in a similar manner from a mantle-derived source. The δ 18O values of the zircon from the S-type granodiorite are not only higher than from the I-type, but also more heterogeneous, consistent with partial melting of a poorly-mixed, metasedimentary source.

Calcium isotope record of Phanerozoic oceans: Implications for chemical evolution of seawater and its causative mechanisms

A total of 280 brachiopods of Ordovician to Cretaceous age, complemented by published data from belemnites and planktonic foraminifera, are used to reconstruct the evolution of calcium isotope composition of seawater (δ 44/40Ca SW) over the Phanerozoic. The compiled δ 44/40Ca SW record shows a general increase from ∼1.3‰ (NIST SRM 915a) at the beginning of the Ordovician to ∼2‰ at present. Superimposed on this trend is a major long-term positive excursion from the Early Carboniferous to Early Permian as well as several short-term, mostly negative, oscillations. A numerical model of the global cycles of calcium, carbon, magnesium and strontium was used to estimate whether the recorded δ 44/40Ca SW variations can be explained by varying magnitudes of input and output fluxes of calcium to the oceans. The model uses the record of marine 87Sr/ 86Sr ratios as proxy for seafloor spreading rates, a record of oceanic Mg/Ca ratios to estimate rates of dolomite formation, and reconstructed atmospheric CO 2, discharge and erosion rates to estimate continental weathering fluxes. The model results indicate that varying magnitudes of the calcium input and output fluxes cannot explain the observed δ 44/40Ca SW trends, suggesting that the isotope signatures of these fluxes must also have changed. As a possible mechanism we suggest variable isotope fractionation in the sedimentary output flux controlled by the dominant mineralogy in marine carbonate deposits, i.e. the oscillating ‘calcite–aragonite seas’. The ultimate control of the calcium isotope budget of the Phanerozoic oceans appears to have been tectonic processes, specifically variable rates of oceanic crust production that modulated the hydrothermal calcium flux and the oceanic Mg/Ca ratio, which in turn controlled the dominant mineralogy of marine carbonates, hence the δ 44/40Ca SW. As to the causes of the short-term oscillations recorded in the secular δ 44/40Ca SW trend, we tentatively propose that these are related to variable rates of dolomite formation and/or to changing chemical composition of the riverine flux, in particular Ca / HCO 3 - and Ca / SO 4 2 - ratios, induced by variable proportions of silicate vs. carbonate weathering rates on the continents.

Assessment of Bacterial Degradation of Aromatic Hydrocarbons in the Environment by Analysis of Stable Carbon Isotope Fractionation

13C/12C stable carbon isotope fractionation was used to assess biodegradation in contaminated aquifers with toluene as a model compound. Different strains of anaerobic bacteria (Thauera aromatica, Geobacter metallireducens, and the sulfate-reducing strain TRM1) showed consistent 13C/12C carbon isotope fractionation with fractionation factors between αC = 1.0017 and 1.0018. In contrast, three cultures of aerobic organisms, using different mono- and dioxygenase enzyme systems to initiate toluene degradation, showed variable isotope fractionation factors of αC = 1.0027 (Pseudomonasputida strain mt-2), αC = 1.0011 (Ralstonia picketii), andαC = 1.0004 (Pseudomonas putida strain F1). The great variability of isotope fractionation between different aerobic bacterial strains suggests that interpretation of isotope data in oxic habitats can only be qualitative. A soil column was run as a model system for contaminated aquifers with toluene as the carbon source and sulfate as the electron acceptor and samples were taken at different ports along the column. Microbial toluene degradation was calculated based on the 13C/12C isotope fractionation factors of the batch culture experiments together with the observed 13C/12C isotope shifts of the residual toluene fractions. The calculated percentage of biodegradation, B, correlated well with the decreasing toluene concentrations at the sampling ports and indicated the increasing extent of biodegradation along the column. The theoretical toluene concentrations as calculated based on the isotope values matched the measured concentrations at the different sampling ports indicating that the Rayleigh equation can be used to calculate biodegradation in quasi closed systems based on measured isotope shifts. A similar attempt was performed to assess toluene degradation in a contaminated, anoxic aquifer. A transect of groundwater wells was monitored along the main direction of the groundwater flow and revealed decreasing concentrations accompanied with an increase in the 13C/12C stable carbon isotope ratio of the residual toluene. Calculation of the extent of biodegradation based on the isotope values and laboratory derived isotope fractionation factors showed that the residual toluene was degraded to more than 99% by microbial activity. Calculation of the theoretical residual toluene concentrations based on the measured isotope values described the strongly decreasing concentrations along the plume. Other aromatic hydrocarbons like benzene and naphthalene which were analysed in the same course also showed decreasing concentrations along the groundwater flow path accompanied by increasing δ13C values indicating biodegradation.

Microbial influence on the oxygen isotopic composition of diagenetic siderite

Numerous early diagenetic siderite concretions previously described have been interpreted as the results of microbially mediated reactions. Interpretation of oxygen isotope data for such material requires an understanding of the effect of temperature on fractionation processes. However, whilst equilibrium fractionation of oxygen isotopes between siderite and water has been measured down to 33°C, extrapolation to lower temperatures may be invalid. Furthermore, inorganic measurements may not be applicable to microbial systems. The specific iron-reducing microorganism, Geobacter metallireducens, has been cultured anaerobically in the laboratory using acetate as an organic substrate and amorphous FeOOH as an electron acceptor. The acetate is oxidised to CO 2, with concurrent iron reduction and extracellular siderite precipitation. Rhombohedral siderite crystals up to 25 μm in size have been precipitated over a range of temperatures (18–40°C). Stable isotopic analysis of these crystals and the solutions from which they precipitate shows that this microbial precipitation of siderite has an associated isotopic fractionation different from the published equilibrium, and which is not simply a function of temperature. In all cases, δ 18O values of siderite are lower than predicted by inorganic equilibrium fractionation data. This may explain the numerous anomalously-low δ 18O values reported for early diagenetic marine siderites and previously attributed to mechanisms such as mixing with meteoric water, sediment-water interaction, recrystallisation, or variable isotopic fractionation, despite a lack of supporting evidence.

Geologic and geochemical controls of mineralization in the Southeast Missouri barite district

Residual barite deposits of the Southeast Missouri barite district were derived from the weathering of lower Paleozoic dolostones, chiefly the Upper Cambrian Potosi and Eminence Dolomites. Open-space, Mississippi Valley-type mineralization in the bedrock was controlled by the intersection of permeable lithologies with late Paleozoic near-vertical faults or linear zones of intense jointing. Linear trends of residual deposits are the surface expression of these bed-rock controls. A consistent mineralogical and barite textural zonation is disposed about the central structures. Sulfide mineralization is concentrated about the central structures, whereas barite is more widespread. Barite textures vary systematically from coarse bladed near the central structures to fine plumose more distally. Such zoning indicates outward ore fluid flow from the central faults and joint zones.Fluid inclusions in sphalerite indicate that warm Na-Ca-Cl brines (69 degrees -105 degrees C, 20-25.5 equiv wt % NaCl) were present during early mineralization. Salinity data for fluid inclusions in barite (0-14 equiv wt % NaCl) may be representative of the barite mineralizing fluid, but contamination by later fluids is possible. Sulfur isotope studies show that the paragenetically early sulfides are isotopically lighter than later barite. Individual minerals involved in mainstage mineralization have the following ranges of delta 34 S values: pyrite-marcasite, 8.4 to 34.7 per mil; galena, 3.2 to 11.8 per mil; sphalerite, 7.2 to 15.8 per mil; barite, 20.5 to 32.2 per mil. Barite delta 34 S values show enrichment in 34 S away from the central fault. This spatial distribution is due to the incorporation of sulfate from two reservoirs. Isotopically light sulfate, developed most strongly about the central fault, originated from the oxidation of H 2 S in the ore fluid emanating from the central fault. The heavy sulfate that prevailed distally was indigenous to the host dolostone formation water. The delta 18 O values of main-stage barite vary from 17.0 to 19.6 per mil and the range is attributed to mixing of two fluids of different oxygen isotope composition and to variable isotopic fractionation between sulfate and water due to changing temperature. The light sulfur and oxygen isotope composition of supergene barite (delta 34 S = 15.8ppm, delta 18 O = 9.0ppm) is due to incorporation of isotopically light sulfate produced locally during the oxidation of sulfide minerals in the soil environment.Barite precipitation was caused by fluid mixing, in which a Ba-Pb-Zn-H 2 S-bearing brine discharged from faults and joints into lower Paleozoic dolostones and mixed with an oxidizing, sulfate-bearing, dilute fluid. Increased supersaturation of barite with progressive mixing caused the systematic spatial variation of barite textures from coarse bladed near the central structures to fine plumose more distally. Similar barite-sulfide mineralization in the Central Missouri barite district probably formed in the same manner. Barite mineralization across the north flank of the Ozark dome postdates local Appalachian-Ouachita-Marathon orogenic deformation and suggests the existence of dilute sulfate-bearing water, probably shallow meteoric ground water, in the Paleozoic section during barite mineralization. Uplift of the Ouachitas in late Paleozoic time apparently initiated Mississippi Valley-type mineralization of the Ozark region.

Land use change and carbon exchange in the tropics: II. Estimates for the entire region: Comment

Estimated annual releases of carbon of 1.2 to 1.5 BMT in 1980 from tropical vegetation and soils due to land use change are presented by Detwiler and others ,(1985). These estimates are compared with the oceanic uptake of CO 2 released from forests and soils as derived from tree ring ~3C records. Detwiler and others (1985) reported that on the basis of new models of oceanic carbon cycle published by Peng and others (1983), the ocean could have accommodated at least 1.2 BMT of forest-soil carbon in 1980. They reached the conclusion that "the global carbon budget may be balanced if there is no significant release from nontropical ecosystems due to land use change and all mature ecosystems are in collective equilibrium with the atmosphere." They misinterpreted our article (Peng and others 1983) and, as a result, their conclusion concerning a balanced global carbon budget cannot be supported. In our article, a time history of the releases of forestsoil CO z, but not of oceanic uptake of this CO 2, is calculated for the period between 1800 and 1980. These release estimates are based on the tree ring xsC records summarized by Freyer and Belacy (1983). The estimated release of forest-soil carbon in 1980, derived from the model calculations involving oceanic uptake of atmospheric ~3C and the response of terrestrial ecosystems to the input of low 13C-CO2, is about 1.2 BMT C/year. The actual amount of forest-soil carbon taken up by the ocean during the whole period is not given in the article. In fact, how much of the forest-soil carbon released into the atmosphere will be taken up by the ocean has not been calculated. A comparison of the results from Detwiler and others (1985) and Peng and others (1983) can best be described as indicating a similar estimate of forest-soil carbon release for 1980. It has nothing to do with a balanced global carbon budget, and as Figure 13 of our article showed, the ocean cannot absorb the excess forest-soil CO z during the period of Mauna Loa record unless we unrealistically manipulate ocean model parameters. Developments since the publication of our article have warranted the downward revision of our estimate of forest-soil releases, weakening the correspondence with the Detwiler and others (1985) results and making their global balance calculations even more untenable. Since our estimate of the time history of forest-soil carbon release depends on tree-ring 13C records, the changes in tree ring tsC data will certainly change in estimates of carbon release history. Freyer (1986) has revised the global mean tree-ring 13C curve on the basis of additional new measurements published since his last summary in 1983 (Freyer and Belacy 1983). Deconvolution of the revised i~C record presents a very different time history of forest-soil carbon release (Peng 1985, Peng and Freyer 1986). Two maximum periods of carbon release were seen, one in 1860 and the other in 1910. For 1980, the estimate from this revised 13C data was about 0.5 BMT C/year. This value is less than half the estimate given by Detwiler and others (1985). A similar calculation of the transfer of biospheric carbon to the atmosphere has been performed by Stuiver and others (1984). Their deconvolution of the average 13C curve derived from six Pacific coastal trees indicates that the forest-soil reservoir became a carbon sink during the periods centered in 1910 and in 1960. The average forest-soil carbon release during the whole period between 1900 and 1975 is estimated to have been about 0.3 BMT C/year. In addition, Stuiver and others studied the relationships between xsC, ring area, and ring width and found a few trees showing significant correlation between ~3C and ring area. Many environmental factors could influence the isotopic composition of trees. Among those, variable isotope fractionation of the trees was the major cause of the variability in the 13C records. Use of the variation of 13C in tree-ring as an analogue of changes in atmospheric a3C has been considered to be questionable. A more reliable way of deconvolving the forest-soil carbon release history should come from an actual time history of atmospheric CO2 variations. Precise measurements of atmospheric CO 2 have been made continuously at Mauna Loa Observatory since 1958 (Bacastow and Keeling 1981). Also, the air bubbles trapped in polar ice offer the possibility of a way to reconstruct the history of atmospheric CO 2 variations before 1958. This can be done by extracting the air bubbles from the ice and measuring the CO 2 concentration in these bubbles, using very precise instruments. A recent article by Neftel and others (1985)

13C/12C ratios in tree rings and the transfer of biospheric carbon to the atmosphere

The δ13C records of 11 Pacific coastal trees have been measured. Relationships between δ13C, ring area, and ring width were investigated. Three of the six trees with records longer than 200 years yield a significant (2σ level or better) correlation between δ13C and ring area. The correlation is most pronounced for Sequoia giganteum, where the variability in ring area accounts for 60% of the δ13C variance. The environmental factors influencing the isotopic composition of trees are many. A major portion of the pre‐1850 variability in the δ13C record is most likely caused by variable isotope fractionation of the trees. Only about one quarter of the δ13C variance in the average record can be removed by normalizing on ring area. The remaining δ13C variability results in a standard deviation of 16 ppm in model‐calculated preindustrial CO2 levels. This is an upper limit for atmospheric CO2 variability induced by changes in the size of the biospheric reservoir. Average preindustrial CO2 levels are estimated to be in the 260–290 ppm range. The cumulative amount of biospheric carbon released to the atmosphere between 1600 and 1975 is approximately 150±100 gigatons C.