Abstract
The isotopic compositions of plants can provide significant insights into paleodiets, ancient agricultural activities, and past environments. Isotopic compositions of charred (aka carbonized) ancient plant remains are typically preferred over those of uncharred/uncarbonized plants, both because charred plants are more commonly preserved and because early research suggested they experience less post-depositional isotopic alteration. In this paper, we re-explore the question of whether uncharred plants experience large-magnitude post-depositional changes in carbon and nitrogen isotope compositions by analyzing Terminal Pleistocene–Early Holocene plant specimens from rockshelters in the Escalante River Basin (Colorado Plateau, southeastern Utah). Several lines of evidence, including C3-CAM differences, plant-part comparisons, and dietary estimates from ancient herbivore collagen, suggest that the original carbon isotope compositions of these plants have not been significantly altered. The preservation status of plant nitrogen isotope compositions is equivocal. The direction of temporal shifts in plant δ15N matches global trends and the magnitude of the shift may have been exacerbated by the extinction of megafauna in an arid environment. However, the Pleistocene plant δ15N values are higher than would be expected based on herbivore bone collagen δ15N. Nevertheless, in contrast to previous research, the ancient uncharred plants in this study did not have exceptionally high δ15N values (> + 25‰). Overall, our research suggests that uncharred plants could be useful substrates for isotopic paleodietary and/or paleoenvironmental studies.
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Introduction
The isotopic compositions of ancient plants can provide crucial baselines for the interpretation of human and animal diets and direct evidence for paleoclimatic conditions and agricultural practices (Blinnikov et al. 2011; Bogaard et al. 2013; Casey and Post 2011; Fiorentino et al. 2015; Kohn 2010; Makarewicz and Sealy 2015; Szpak et al. 2013; Vaiglova et al. 2014a; Warinner et al. 2013; Wooller et al. 2011). Isotopic studies of ancient plants typically rely on charred/carbonized remains, in part because they are more often preserved in archaeological contexts (Charles et al. 2015; Fiorentino et al. 2015), but also because a pioneering study by DeNiro and Hastorf (1985) concluded that uncharred plants are much more susceptible to diagenetic alteration. DeNiro and Hastorf (1985) argued that charred plant remains from Peruvian archaeological sites were reasonably well preserved since their range of δ13C and δ15N values and magnitude of intra-plant variations (≤ 2‰) were similar to those of the modern plants they studied. Subsequent experimental studies supported this conclusion, demonstrating that isotopic shifts due to charring were relatively small (typically non-systematic shifts of < 1‰ for δ13C and increases in δ15N of up to ~ 2‰) (Aguilera et al. 2008; Araus et al. 1997; Bogaard et al. 2007; DeNiro and Hastorf 1985; Fiorentino et al. 2012; Fraser et al. 2013; Heaton et al. 2009; Kanstrup et al. 2012; Marino and DeNiro 1987; Nitsch et al. 2015; Poole et al. 2002; Styring et al. 2013; Tieszen and Fagre 1993; Yang et al. 2011). In contrast to charred plants, uncarbonized archaeological plant isotopic compositions were thought to be significantly altered, since (1) the δ15N values of uncharred plants were considerably (10–35‰) higher than those of modern plants, and (2) intra-plant variations in uncharred remains were much larger than those of the modern plants (up to 8‰ differences for δ13C and 21‰ for δ15N) (DeNiro and Hastorf 1985).
In this study, we utilize macroscopically well-preserved ancient plant samples from the Escalante River Basin (ERB) to re-examine DeNiro and Hastorf’s (1985) conclusion that uncharred plant C- and N-isotope values are significantly diagenetically altered. We do not seek to prove that absolutely no alteration has occurred, but rather to determine whether or not there has been significant alteration of original plant isotope compositions (i.e., shifts of > 1–2‰). We begin with the assumption that uncharred plants retain the original isotopic compositions of the plant while it was alive and seek to disprove this assumption. We consider four lines of evidence. First, we discuss whether or not ancient plant δ-values and intra-plant variations are outside the range expected for modern plants, taking into account paleoenvironmental changes. These were the original criteria used by DeNiro and Hastorf (1985) to assess preservation of their ancient plants, though the isotopic effects of paleoenvironmental change were not well understood when that paper was published. Second, we examine temporal changes in ancient Escalante River Basin plant δ13C and δ15N values to determine whether they match expectations based on local circumstances and global trends. Third, we compare mean plant δ-values with diet estimates derived from local herbivore bone collagen. Fourth, we report the %C, %N, and C/N of the plant samples and their Fourier-transform infrared (FTIR) spectra for comparison with those of modern plants.
Modern Plants
Carbon Isotopes
Since the first reports that photosynthetic pathways affect carbon isotope discrimination, the typical ranges of δ13C for C3, C4, and CAM plants have been well documented (Cerling et al. 1999; Kohn 2010; O'Leary 1988; Sternberg et al. 1984). Most plants utilize the C3 photosynthetic pathway (Calvin cycle), which results in low δ13C values (range = − 37 to − 20‰, mean = − 27‰ for modern plants) (Kohn 2010). C3 plants with δ13C values > − 25.5‰ are typically only found in areas with mean annual precipitation of < 500 mm/year (Kohn 2010). Very low δ13C values in C3 plants (< − 31.5‰) are found only in closed-canopy forest areas with low light levels and substantial recycling of 13C-depleted CO2 (Kohn 2010; Medina et al. 1991; Van Der Merwe and Medina 1991). C4 plants have much higher δ13C values, mostly within the range of − 16 to − 10‰ (Cerling et al. 1999; O'Leary 1988). Obligate/constitutive CAM plants have δ13C values similar to C4 plants or a little higher, typically ranging between about − 14 and − 10‰ (Sternberg et al. 1984). Facultative (inducible) CAM plants can switch between CAM and C3 photosynthesis and can have values that can fall within or between the C3 and C4 plant ranges (about − 30 to − 10‰), depending on the degree to which each pathway is used during tissue formation (Kluge et al. 1991; Osmond et al. 1973; Teeri and Gurevitch 1984). Recent research suggests that obligate versus facultative CAM are not truly discrete categories, but rather, that some degree of facultative control might be present in all species capable of CAM photosynthesis (Winter et al. 2008). In CAM plants, higher values are produced when photosynthesis occurs in the dark, which tends to occur in areas with hot, dry days and relatively cool nights (Osmond et al. 1973). It should be noted that any comparison of the δ13C values of modern and ancient plants must take into account the depletion of 13C in atmospheric CO2—and hence, in modern plants—that has occurred since humans began burning fossil fuels in industrial processes (i.e., the Suess Effect). This led to an approximately 1‰ decrease from ad 1880 to 1980 and a further 1‰ decrease from 1980 to 2009 (Bocherens et al. 2014; Francey et al. 1999; Friedli et al. 1986; Long et al. 2005).
Different parts of the same plant can have different δ13C values (O'Leary 1981; Tieszen 1991; Tieszen and Fagre 1993). The most consistently observed difference is a ca. 1–3‰ lower δ13C value in leaves relative to non-photosynthetic tissues such as stems and roots (Badeck et al. 2005; Ghashghaie and Badeck 2014; Hobbie and Werner 2004). These differences are produced by several factors, including (but not limited to) different biochemical compositions (e.g., low-13C lipids and lignin vs. high-13C cellulose, sugar, and starch), different formation times, and isotopic fractionations during transportation of biomolecules (Badeck et al. 2005; Cernusak et al. 2009; Ghashghaie and Badeck 2014).
Nitrogen Isotopes
It is more difficult to define a “typical” range of plant δ15N values since these values are not dependent on fixed processes like photosynthetic pathways but rather on complexities such as nitrogen sources, nitrogen availability, nitrogen losses, symbiotic associations with mycorrhizae, biochemical fractionations within plants, and environmental factors such as temperature and precipitation (see Szpak 2014 for review). In a survey of modern Peruvian plants from a variety of ecozones, Szpak et al. (2013) found that aboveground plant tissues ranged from − 5.3 to + 17.3‰. In a study of Kansas plants derived from herbarium samples dating to the past 130 years, McLauchlan et al. (2010) measured foliar δ13C values between − 7.5 and + 14.5‰. Within single genera of African plants that associate with nitrogen-fixing bacteria, Cerling et al. (2009) found δ15N values that ranged from about 0 to + 11‰ (Indogofera spp.) and + 3.5 to + 15.9‰ (Acacia spp.). In a study of modern Yukon plants (including various plant parts), Tahmasebi et al. (2017) measured δ15N values between about − 10 to + 27‰, with the majority of measurements falling between about − 5 and + 5‰. Experimental studies have resulted in plant δ15N values as high as + 25‰ when grown with pig manure and up to + 45‰ when fertilized with seabird guano (Szpak 2014; Szpak et al. 2012a). One field study demonstrated plant δ15N variations in a gradient from about + 20‰ near the nesting sites of seabirds such as penguins and albatross, to − 10‰ at more distant locations (Erskine et al. 1998). The highest values resulted from large fractionations that occur when seabird excrement is mineralized: the ammonia gas has a very low δ15N value and therefore the remaining guano becomes enriched in 15N, which increases the δ15N values of plants growing near the nests. Low δ15N values at inland sites downwind of the nesting sites suggested that inland plants utilized some of the 15N-depleted ammonia derived from the guano.
The effects of terrestrial herbivore dung and urine on plant δ15N in natural environments is less straightforward. In Yellowstone National Park, areas frequented by grazing herbivores had higher soil δ15N but lower aboveground plant δ15N than areas from which herbivores were excluded (Frank and Evans 1997; Frank et al. 2004). Herbivores increase N losses from soils through processes like ammonia (NH3) volatilization from urine/feces, which results in 15N enrichment of soils (Frank and Zhang 1997). However, plants grown in these soils may then rely more on NO3− (produced by bacteria and depleted in 15N) than NH4+ (enriched in 15N), leading to lower plant δ15N values (Frank and Zhang 1997). Lower aboveground plant δ15N can also be caused by absorption of volatile ammonia by shoots and leaves through their stomata and cuticle, as was the case in the seabird study described above (Erskine et al. 1998; Frank et al. 2004). It should be noted that the studies of herbivore effects on plant δ15N values cited above were conducted in temperate grasslands, and nitrogen dynamics in arid environments can be very different. For example, abiotic (rather than microbial) gas formation plays a key role in nitrogen loss in desert environments, particularly when soils become wet (McCalley and Sparks 2009). In combination with herbivore nitrogen inputs to soils, this could theoretically cause very large nitrogen fractionations, leading to 15N-enriched soils and plants. Herbivores can also affect nitrogen cycling in arid ecosystems by damaging/destroying (layers of cyanobacteria, lichens, and mosses) that fix atmospheric nitrogen and provide the primary nitrogen input to some arid ecosystems (Evans and Ehleringer 1993).
Environmental variables such as soil nitrogen availability and climate (temperature, aridity) can also affect plant δ15N. Numerous studies have noted a negative correlation between foliar δ15N and mean annual precipitation (Amundson et al. 2003; Austin and Vitousek 1998; Craine et al. 2009; Handley et al. 1999; Hofmeister et al. 2012). Very generally, plants growing in more arid environments tend to have higher δ15N values than those growing in wet environments. Some studies have found a “hump-shaped pattern” between soil/plant δ15N and aridity, where δ15N increases with aridity up to a threshold value, after which increases in aridity are associated with decreases in soil/plant δ15N (Díaz et al. 2016; Wang et al. 2014). Some studies have also found a positive correlation between δ15N and local temperature (Amundson et al. 2003; Craine et al. 2009; Martinelli et al. 1999), but this is generally less robust than the relationship with aridity.
Significant variations in δ15N can occur within a single plant. In many plants (especially annuals), stored proteins are mobilized to newly grown plant parts (i.e., those with reproductive functions), resulting in intra-plant nitrogen sources (e.g., leaf/stem/bark) having higher δ15N values than intra-plant nitrogen sinks (e.g., fruits, grains, and/or flowers) or whole plants (Choi et al. 2002; Crawford et al. 1982; Szpak 2014; Szpak et al. 2012a, 2013). However, some studies have found no difference between these tissues (e.g., Szpak et al. 2013; Tahmasebi et al. 2017), and others provided seemingly contradictory results—for example, that cereal grains have higher δ15N values than the rachis (Fraser et al. 2011).
Isotopic Effects of Plant Decomposition
Plant decomposition is a continuum that begins with the accumulation of plant litter and ends with the formation and stabilization of soil organic matter (Melillo et al. 1989). Numerous studies have examined changes in soil δ13C, δ15N, %C, %N, and C/N ratios with depth that are related in part to decomposition processes (but also to fractionations associated with transfer of N from soil to aboveground portions of growing plants). Carbon isotope fractionation during plant tissue degradation and soil formation is typically relatively small, and most studies have found either no change in δ13C with depth, or increases of up to a few permil (Boström et al. 2007; Brueggemann et al. 2011; Ehleringer et al. 2000; Kramer et al. 2003; Krull et al. 2016; Tiunov 2007). In contrast, biochemical reactions of the nitrogen cycle (e.g., nitrification, ammonification, uptake, and assimilation of N into plants) are associated with very large fractionations and hence much greater changes in plant/soil δ15N (Robinson 2001). Soil organic matter δ15N can increase by up to 20‰ with depth (Hogberg et al. 1996; Kramer et al. 2003; Ponsard and Arditi 2000; Scheu and Falca 2000; Steele et al. 1981; Tiunov 2007; Turner et al. 1983). The magnitude of 15N enrichment with depth seems to depend little on mean annual temperature, precipitation, or nitrification rates, but depends strongly on the dominant type of mycorrhizal fungus associated with plant roots (Hobbie and Ouimette 2009). The largest increases (~ 9.6‰, on average) occur in ectomycorrhizal systems and smaller increases (~ 4.6‰) in arbuscular mycorrhizal systems (Hobbie and Ouimette 2009; Martinelli et al. 1999). The former are more typically found in temperate regions, including forests, whereas the latter are more common in tropical areas. Typically, the total carbon, total nitrogen, and C/N ratios of soil organic matter all decrease with depth (Boström et al. 2007; Krull et al. 2016; Melillo et al. 1989; Pardo et al. 1997; Ponsard and Arditi 2000; Salazar et al. 2012; Scheu and Falca 2000; Steele et al. 1981; Vervaet et al. 2002).
Experimental studies have shown that plant decomposition is associated with mass loss and overall decreases in C/N ratios, though increases in C/N ratios can also occur (Benner et al. 1991; Connin et al. 2001; Salazar et al. 2012). The effects of decomposition on %N and δ15N are less straightforward. Benner et al. (1991) found that decomposing plant %N decreased over a period of about 120 days (with no significant change in δ15N), then increased substantially (along with a δ15N decrease of ~ 2‰). Connin et al. (2001) found no consistent relationship between changes in δ15N and %N over time in decomposing plants in an arid environment, and the magnitude of changes in δ15N were small (≤ 2.6‰). In a colder, temperate context, Tahmasebi et al. (2018) measured much larger δ15N increases of 2.5 to 10‰ in decomposing buried plant samples. Decomposing plant δ15N values were negatively correlated with C/N and %C, but no relationship between δ15N and %N was reported (Tahmasebi et al. 2018). The studies cited above suggest that decomposition processes may occur differently—and be associated with smaller isotopic shifts—in arid relative to temperate environments. For example, photodegradation of plant material (abiotic breakdown by solar radiation) accounts for a significant portion of plant decomposition and gas release in hot, dry ecosystems (King et al. 2012; Lee et al. 2011), whereas microbial degradation is the dominant process in most other ecosystems. In hot, dry, shaded areas (such as Escalante River Basin alcoves), plant litter would be protected from both photodegradation (due to lack of sunlight) and microbial alteration (due to lack of water), which would account for the exceptional preservation of samples from these locations.
Evershed et al. (1997) demonstrated that extraordinarily well-preserved archaeological plant remains from ancient Egypt had undergone Maillard reactions between proteins and sugars, and that the volatiles produced during these reactions were trapped inside the desiccated plant remains. They suggested that Maillard reactions, which form stable compounds called melanoidins, are typical of plant decay. In normal circumstances, volatiles produced by the Maillard reaction are released when plants are broken down, but ancient plants preserved in hot, dry environments retain structural and storage macromolecules that trap the volatiles within the plant (Evershed et al. 1997). Bland et al. (1998) demonstrated that desiccated archaeological seeds had undergone Maillard reactions but that the composition of proteins and polysaccharides in the samples had undergone minimal changes. Isotopic shifts associated with Maillard reactions in experimentally charred samples have been studied and are relatively minimal (< 1‰ for δ13C and up to ~ 2‰ for δ15N) (Aguilera et al. 2008; Araus et al. 1997; Bogaard et al. 2007; DeNiro and Hastorf 1985; Fiorentino et al. 2012; Fraser et al. 2013; Heaton et al. 2009; Kanstrup et al. 2012; Marino and DeNiro 1987; Nitsch et al. 2015; Poole et al. 2002; Styring et al. 2013; Tieszen and Fagre 1993; Yang et al. 2011). The magnitude of fractionations that might occur when Maillard reactions proceed at lower temperatures and/or in association with ancient desiccated plants has not been explored.
Can %C, %N, or C/N Be Used to Assess Post-Depositional Isotopic Alteration of Plants?
In isotopic studies of bone collagen, well-established ranges for carbon and nitrogen contents (%C, %N) and atomic C/N ratios are used to identify well-preserved versus diagenetically altered samples (Ambrose 1990; DeNiro 1985; van Klinken 1999). As noted by Szpak et al. (2013) and Szpak (2014), no such quality control indicators exist for ancient plant remains, in part because of the chemical heterogeneity of plant samples. For modern (uncharred) plants, mean carbon contents tend to be between ~ 40 and 47%, mean nitrogen contents are generally between ~ 1 and 3%, and mean atomic C/N ratios are typically between ~ 20 and 60 (Cerling et al. 2003; Codron et al. 2007; Fraser et al. 2011; Kristensen et al. 2011; Martinelli et al. 1999; Styring et al. 2016; Tahmasebi et al. 2017; Wooller et al. 2003). However, the full ranges of values for individual modern plant samples in any given study can be much larger. For example, Tahmasebi et al. (2017) measured atomic C/N values from 13 to 157 in modern Yukon plants. Furthermore, several studies have shown that large-magnitude changes in %C and %N associated with senescence (age-related remobilization of nutrients to other parts of the plant) do not necessarily correspond to any changes in δ13C or δ15N (Kolb and Evans 2002; Szpak et al. 2012a; Wooller et al. 2003). Thus, %C, %N, and C/N ranges cannot be used as quality control indicators for δ13C and δ15N of ancient plants.
Correlations between plant %C, %N, or C/N and isotope values can occur as a result of decomposition processes and might be used as evidence for alteration; however, similar correlations can also occur as a result of variations that occur when plants are alive. Thus, such correlations are not acceptable quality-control indicators. For example, experimental studies have demonstrated that increases in leaf litter δ15N values during decomposition may be accompanied by increases in %N, decreases in %C, and decreases in C/N (Connin et al. 2001; Salazar et al. 2012; Tahmasebi et al. 2018). However, similar correlations have been reported in freshly collected modern plants. Correlations between modern plant δ15N and %N have been reported in such diverse environments as African savanna (Codron et al. 2013), Hawaiian rainforest (Vitousek et al. 1989), recently deglaciated Alaska (Hobbie et al. 2000), southern Yukon grasslands (Tahmasebi et al. 2017), and in global compilations (Craine et al. 2009; Martinelli et al. 1999). In modern C3 plants from Elk Island National Park (Alberta, Canada), a positive correlation between δ15N and %N (r = 0.27, p < 0.01) and between δ15N and C/N (r = 0.34, p < 0.001) were observed (Metcalfe, unpublished data). These examples show that these relationships do not provide robust evidence for post-depositional alteration of ancient plant isotopic compositions. Thus, we report %C, %N, and C/N ratios for the plants examined in the current study, but do not discuss them further as quality-control indicators.
Study Area
In the mid-1980s, ancient megafaunal dung and plant samples were collected from rockshelters/alcoves in the Forty-Mile Canyon and Willow Gulch areas of the Escalante River Basin (Glen Canyon National Recreation Area). The dung samples were used to constrain the timing of Late Pleistocene megafaunal extinctions/extirpations (Mead and Agenbroad 1992) and to reconstruct the dietary habits of extinct taxa (Davis et al. 1984; Kropf et al. 2007). The plant samples were used to reconstruct changes in vegetation in response to Late Pleistocene–Early Holocene climate change (Withers and Mead 1993). Since their collection, the remaining materials have been curated at the Museum of Northern Arizona (Flagstaff) on behalf of the Glen Canyon National Recreation Area. In this study, we examine the isotopic compositions of some of these morphologically well-preserved uncharred plant samples (Fig. 1) to explore their potential for reconstructing paleoenvironments and as baselines for isotopic studies of ancient animal diets.
Representative samples of Late Pleistocene–Early Holocene uncharred plants from the Escalante River Basin on the Colorado Plateau. a Seed of Bigtooth maple (Acer grandidentatum).b Seeds of box elder (Acer negundo). c Seeds of Gambel oak (Quercus gambelii). d Leaf of cottonwood (Populus cf. fremontii). e Unidentified twig. f Seeds of Opuntia cactus. g Needles of Opuntia cactus
The Escalante River Basin is located in southern Utah, on the central Colorado Plateau. With elevations ranging from ~ 1100 m at the Escalante River to ~ 2300 m at the top of the Kaiparowits Plateau, the canyons in this area support diverse environments and plant communities (Anderson et al. 2000; Withers and Mead 1993). Pinyon–juniper woodlands dominate the upper elevations. The middle and lower elevations host a wide range of habitats, including riparian woodlands, hanging gardens, active sand dunes, and mixed shrublands/grasslands (Agenbroad et al. 1989; Withers and Mead 1993). Temperature and precipitation also vary with altitude: along the river, mean annual temperature (MAT) is 12 °C and mean annual precipitation (MAP) is ≤ 250 mm, whereas on the Aquarius Plateau MAT is 2 °C and MAP is 575 mm (Webb 1985). The Forty-Mile Canyon alcoves (Hooper’s Hollow, Grobot Grotto, BF Alcove) are located between 1100 and 1300 m elevation, about 10 m above a steep drop to the deeply entrenched modern streambed (Withers and Mead 1993). Hooper’s Hollow and Grobot Grotto are large (100–200 m wide) southwest-facing alcoves whereas BF alcove is narrower and faces northeast. The modern vegetation around these alcoves is predominantly desert grasses with a few scattered shrubs. The Willow Gulch alcoves (Shrub-Ox Alcove, Oak Haven) face northwest and west (respectively) and host a greater number and diversity of arboreal and shrub species, as well as grasses and cacti. Withers and Mead (1993) collected plant macrofossils from “vegetation mats” buried under eolian and/or colluvial sands within these alcoves. Some of these vegetation mats included dung with perfectly preserved morphology, indicating a lack of physical or chemical breakdown of material within the mats. Withers and Mead argued that the plants grew in the immediate vicinity of the alcoves and were essentially deposited in situ. Radiocarbon dates (n = 14) directly on the plant remains indicate that most of the samples were deposited between the terminal Pleistocene (ca. 13,000 14C yr. bp) and early Holocene (ca. 7500 14C yr. bp) (Withers and Mead 1993). Xerophytic grasses, shrubs, and cacti (Opuntia sp.) occurred throughout the temporal sequence. Terminal Pleistocene vegetation also included conifers and mesophytic species such as rose and water birch. Early Holocene vegetation lacked conifers and included reduced abundances of mesophytic species and increased reticulated hackberry. Terminal Early Holocene vegetation included abundant Gambel oak and prickly pear, and little else. The plant assemblage data suggest that there was a significant decrease in water availability over time, likely related to decreased stream levels rather than increased local precipitation (Withers and Mead 1993). The plant assemblage data also suggests that summers were cooler and drier, and winters were warmer and wetter, during these early time periods relative to the present (Withers and Mead 1993).
Materials and Methods
Ancient C3 and CAM plant samples were obtained from National Park Service collections housed at the Museum of Northern Arizona (Fig. 1, Table 1). The samples were derived from five alcoves in the Escalante River Basin of the Colorado Plateau (BF Alcove, Grobot Grotto, Hooper’s Hollow, Oak Haven, and Shrub-Ox Alcove). Plant samples from vegetation mats dating between 12,690 ± 180 and 7510 ± 160 14C yr. bp (Mead and Agenbroad 1992) were selected for carbon and nitrogen isotope analysis. The plant samples in this study were not directly dated, but dates were assigned based on stratigraphic association with directly dated plant specimens (Table 1) (Withers and Mead 1993). For temporal comparisons, we followed Withers and Mead (1993) in dividing the samples into three phases: Late Pleistocene (> 11,000 14C yr. bp), Early Holocene (11,000–8000 14C yr. bp), and Terminal Early Holocene (8000–7000 14C yr. bp).
Plant samples utilized in this study were dry-screened and separated from sediments but never subjected to chemical cleaning or wet processing (Withers and Mead 1993). Prior to isotopic analysis, plant parts (e.g., seeds, leaves, twigs/branches, needles) were separated, and each sample was ground to a fine powder using a Wig-L-Bug ball mill. A total of 90 samples (81 C3 and 9 CAM plants) were analyzed, including 11 different identifiable species (Table 1). Stable carbon and nitrogen isotope values of the plant samples were obtained at the University of British Columbia using an Isoprime stable isotope ratio mass spectrometer coupled to an Elementar VarioMicro Cube elemental analyzer. All plant samples were analyzed two or more times (Appendix A). USGS-40 and USGS-41 were used to calibrate the δ13C and δ15N to the VPDB and AIR scales, respectively. Analytical uncertainty was monitored using an in-house methionine standard with well-characterized isotopic compositions (δ13C = − 28.60 ± 0.09‰, δ15N = − 5.04 ± 0.14‰), and two international standards (IAEA CH-3 cellulose and IAEA CH-6 sucrose). Following the method presented by Szpak et al. (2017), precision (u(Rw)) was measured at ± 0.18‰ for δ13C and ± 0.16‰ for δ15N based on repeated measurements of calibration standards, check standards, and sample replicates. Accuracy or systematic error, u(bias), was ± 0.09‰ for δ13C and ± 0.16‰ for δ15N. The total analytical uncertainty (uc) was estimated to be ± 0.20‰ for δ13C and ± 0.23‰ for δ15N. Further details are provided in Appendix A.
Carbon and nitrogen contents of plant samples were calculated based on the amplitudes of the major carbon and nitrogen peaks relative to the weight of the sample, calibrated using USGS-40 (40.82 %C, 9.52 %N). Methionine (40.25 %C, 9.39 %N) and USGS-41 (41.90 %C, 9.76 %N) were used as check standards. Accuracy and precision were analyzed using the same approach as for isotope values (described above). For %C, precision (u(Rw)) was ± 2.78%, accuracy u(bias) was ± 0.83%, and the total analytical uncertainty (uc) was ± 2.90%. For %N, precision (u(Rw)) was ± 0.21%, accuracy u(bias) was ± 0.15%, and the total analytical uncertainty (uc) was ± 0.26% (Appendix A).
Skeletal remains are rare in Glen Canyon deposits, but five faunal samples were obtained for comparison of their δ13C and δ15N with those of the plants: a Harrington’s mountain goat (Oreamnos harringtoni) tooth, a shrub-ox (Euceratherium collinum) tooth, and three mammoth (Mammuthus sp.) bones (Table 2). Bone/dentin was demineralized using 0.5 M HCl, gelatinized at 70 °C, and ultrafiltered to obtain the > 30 kDa size fraction. Two of the mammoth bone samples did not have sufficient collagen for analysis. The C and N isotope values and element contents of the remaining three samples were measured at the Max Planck Institute for Evolutionary Anthropology in Leipzig, Germany.
FTIR was used to examine the molecular structure of selected ancient plant samples (Gambel oak seeds (Quercus gambelii), maple seeds (Acer spp.), and cactus needles (Opuntia sp.)) as well as modern samples from the same categories. The ancient samples were selected to include specimens with disparate δ13C and δ15N values. Modern oak and maple seeds were collected from the ground underneath trees growing outdoors in Vancouver, British Columbia. Cactus needles were collected from plants (one alive, one dead) grown indoors in Vancouver, BC. All samples (2 mg) were mixed with KBr (200 mg) and compressed into 12-mm pellets. Absorbance spectra between 400 and 4000 cm−1 were obtained using a Bruker Vector 22 FTIR spectrometer, scanning 16 times with a resolution of 4 cm−1. A baseline correction was applied to the region between about 3800 and 400 cm−1. All spectra were normalized to the amide II peak (~ 1610 to 1650 cm−1), as suggested by Baker et al. (2014).
Results
Carbon and nitrogen isotope values, C and N contents, and atomic C/N ratios for ancient uncharred Escalante River Basin plants are presented in Table 1 and Fig. 2. For ancient C3 plants, the mean (± 1σ) δ13C was − 25.5 ± 2.0‰ (range = − 28.8 to − 20.6‰, n = 81) and the mean δ15N was + 4.7 ± 5.5‰ (range = − 12.3 to + 13.6‰). For ancient CAM plants (Opuntia sp.), the mean δ13C was − 10.3 ± 0.5‰ (range = − 10.8 to − 9.6‰, n = 9) and the mean δ15N was + 11.7 ± 4.3‰ (range = + 8.9 to + 22.6‰). There was a statistically significant positive correlation between δ13C and δ15N for the entire ancient plant sample set (Pearson’s r = 0.54, p < 0.001, df = 88) and for C3 plants alone (r = 0.58, p < 0.001, df = 79). There was no significant correlation between δ13C and δ15N for ancient CAM plants alone.
Carbon and nitrogen isotope values of ancient plants from the Escalante River Basin. a Bivariate plot of carbon and nitrogen isotope compositions. b Frequency of measured δ13C values. c Frequency of measured δ15N values
For ancient C3 plants (excluding conifers), δ13C values of leaves (− 27.0 ± 1.1‰, n = 13) were lower than those of other plant parts (− 25.3 ± 2.1‰, n = 63), but the difference was not significant (t = 1.5, df = 30, p = 0.07). However, the δ13C of leaves was significantly lower than that of seeds, both for the ancient plant sample set as a whole (t = 4.0, df = 38, p < 0.001) and for oak alone (t = 2.0, df = 21, p < 0.05) (Table 3). Comparing the seeds of different taxa, oak specimens had δ13C values that were 2.5‰ lower than those of maple (t = 3.9, df = 25, p < 0.001) (Table 3). There was no significant difference between the δ15N of plant tissues that act as nitrogen sources (leaf/stem/bark) versus those that act as sinks (fruit/seed/flower) for the whole sample (+ 3.2 ± 5.6‰ vs. + 5.5 ± 5.1‰), oak alone (− 0.9 ± 8.0‰ vs. + 3.7 ± 6.1‰), or rose alone (+ 2.9 ± 1.7‰ vs. + 4.8 ± 5.3‰).
Early Holocene and Terminal Early Holocene C3 plants had significantly lower δ13C (F2,75 = 18.9, p < 0.001) and δ15N values (F2,75 = 28.7, p < 0.001) than Late Pleistocene C3 plants (Table 4, Fig. 3). With Gambel oak removed (because the results above indicated that oak had lower δ13C values), there was no longer a temporal difference between the Late Pleistocene and Early Holocene for δ13C (t = 1.3, df = 5, p = 0.2), but the significant difference for δ15N remained (t = 4.0, df = 5, p < 0.001).
Temporal changes in the δ13C and δ15N of ancient Escalante River Basin C3 plants, from the Pleistocene (P, > 11,000 14C yr. bp) to the Early Holocene (EH, 11,000–8000 14C yr. bp) and Terminal Early Holocene (TEH, 8000–7000 14C yr. bp). Open circles are results from individual plant samples. Closed circles are means and error bars represent one standard deviation
The ancient Escalante River Basin plants had %C of 32.8 ± 3.8%, %N of 1.9 ± 1.0%, and atomic C/N of 26.6 ± 15.9. C3 and CAM plants did not have significantly different %C, %N, or C/N values (Table 4). For the whole ancient plant sample set, δ15N was positively correlated with %N (r = 0.42, p < 0.001), negatively correlated with C/N (r = − 0.52, p < 0.001), and not correlated with %C (p > 0.05). For ancient C3 plants alone, δ15N was positively correlated with %N (r = 0.52, p < 0.001), and negatively correlated with C/N (r = −0.62, p < 0.001) and %C (r = − 0.25, p < 0.05). For ancient CAM plants alone, there were no significant correlations between δ15N and %N, C/N, or %C (p > 0.05).
For the whole ancient plant sample set, there were no correlations between δ13C and %C, %N, or C/N (p > 0.05). For ancient C3 plants alone, δ13C was positively correlated with %N (r = 0.49, p < 0.001) and negatively correlated with %C (r = − 0.43, p < 0.001) and C/N (r = − 0.55, p < 0.001). For ancient CAM plants alone, there were no correlations between δ13C and %C, %N, or C/N (p > 0.05).
FTIR spectra of modern and ancient plant samples were broadly similar, with prominent peaks representing amides, CH2, OH, and C–O, which are characteristic of plant remains (Fig. 4) (Baker et al. 2014; Styring et al. 2013). CH2 and amide I peaks were less pronounced in some ancient plant samples, but loss or diminishment of these peaks was not necessarily greater in samples with higher δ15N values (Fig. 4). There was no evidence for contamination with carbonates (870 and 720 cm−1), nitrates (3300, 1450, and 1985 cm−1), or humic salts (1080 and 1010 cm−1) in any of the samples, though we note that the detection limits for these compounds are 5–10% of the total sample weight (Vaiglova et al. 2014b).
Baseline-corrected and normalized FTIR spectra for modern and ancient a maple seeds, b oak seeds, and c cactus needles. GLCA catalog numbers and carbon and nitrogen isotope compositions for the ancient specimens are indicated
Ancient herbivore bone and tooth collagen samples had acceptable carbon and nitrogen contents (35–42% and 10–15%, respectively) and C/N ratios (3.2 to 3.6). The mammoth, mountain goat, and shrub-ox had similar δ13C values (− 19.2, − 19.4, and − 18.8‰, respectively). The mammoth had a higher δ15N value (+ 7.8‰) than the mountain goat or shrub-ox (+ 4.9 and + 4.7‰, respectively).
Discussion
FTIR
The reduction or loss of peaks attributable to CH2 and C–O groups in ancient plant samples, along with their lower %C relative to modern plants, indicates some loss of organic carbon. Styring et al. (2013) attributed reduced C–O peaks in charred cereal grains to conversion of starch into melanoidins during Maillard reactions. It is possible a similar process occurred in the ancient uncharred plant samples from the Escalante River Basin, albeit at lower temperatures (which, theoretically, might cause larger isotopic fractionations). However, as discussed above, mass losses or changes in biochemical composition do not necessarily indicate that isotopic alteration has occurred since isotopic fractionation does not always result from changes in %C, %N, or mass loss. If the high δ15N values observed in the ancient uncharred plants resulted from diagenetic alteration, we might expect the greatest changes in the FTIR spectra to be in the amide peaks of samples with the highest δ15N values, but this was not the case. A controlled study of the spectroscopic and isotopic effects of plant decomposition would be useful in further exploring the use of FTIR to identify isotopically altered plant samples.
Have the Isotopic Compositions of the Ancient Plants Been Altered?
Carbon Isotope Compositions
There is no strong evidence that the original plant δ13C values have been significantly altered. The 15.2‰ separation between C3 and CAM plants and the range of δ13C values for each photosynthetic category (Fig. 2) is similar to what is seen in modern plants. The ERB leaves had lower δ13C values than other plant parts (e.g., seeds), a phenomenon that has been reported in numerous previous studies of modern plants (Cerling et al. 2004; Cernusak et al. 2009; Codron et al. 2005; Hobbie and Werner 2004; Leavitt and Long 1982; Szpak et al. 2013). The lower δ13C values of oak seeds compared to maple is consistent with the former species occupying wetter areas, and/or relying on winter precipitation and deeper water reserves (Phillips and Ehleringer 1995; Williams and Ehleringer 2000; Withers and Mead 1993). Finally, the mean δ13C of C3 plants matches the expectation for local herbivores based on their bone collagen δ13C values. Large herbivores have a collagen-diet offset of about 5‰ (Drucker et al. 2008), so their mean collagen δ13C (− 19.1‰) suggests a diet δ13C of − 24.1‰, which is close to the mean δ13C of Late Pleistocene C3 plants measured in this study (− 24.7‰). It is interesting to note that despite their different dietary preferences (mammoths are grazers, shrub-ox are browsers, and mountain goats are mixed feeders) all three consumed a C3-dominated diet, which suggests that C4 grasses and shrubs were rare or absent in this environment.
When the mean δ13C for C3 plants (− 25.5‰) is converted to a “modern equivalent” by subtracting 1.5‰ to account for the depletion of 13C in modern atmospheric CO2 due to fossil fuel combustion (i.e., the “Suess Effect”) (Bocherens et al. 2014; Long et al. 2005), the result (− 27.0‰) is identical to the modern global mean C3 plant value (excluding closed-canopy environments) (Kohn 2010). Likewise, the “modern equivalent” for the mean CAM plant δ13C (− 11.8‰) is consistent with δ13C values of modern obligate CAM plants (Fleming et al. 1993; Sternberg et al. 1984; Winter and Holtum 2002) and is within the range of values reported for modern Opuntia cacti (Sayed 2001). The ancient C3 plant δ13C values are lower than we might expect for an arid environment, but consistent with Withers and Mead’s (1993) conclusion that the ancient plants grew in relatively wetter environments than are present at the alcoves today. Flanagan et al. (1997) reported shrub and tree δ13C values from modern Glen Canyon environments that are useful for comparison: − 26.7 ± 1.2‰ in the riparian zone, a combined mean of − 27.7 ± 1.3‰ in five hanging gardens (plants growing around water seeps in canyon walls), and − 25.4 ± 1.6‰ in the ephemeral wash (a nearby arid zone). The δ13C of the C3 plants presented in the present study (once corrected for the Suess Effect) are consistent with growth in a riparian environment and with the inference of a higher water table in the Escalante River Basin during the terminal Pleistocene–Early Holocene. The δ13C values of the CAM plants are consistent with photosynthesis in an environment with hot, dry days and cool nights. Altogether, these results suggest that no large-magnitude diagenetic shifts in ancient plant δ13C values have occurred.
Nitrogen Isotope Compositions
The comparison between ancient and modern plant isotopic compositions is more complicated for nitrogen than it is for carbon. Nitrogen isotope compositions of plants vary dramatically with environmental differences, and many/most modern plant studies obtained samples from diverse microhabitats. On the other hand, there were highly variable microhabitats in the vicinity of the Escalante River Basin alcoves from which the ancient plants were recovered (Withers and Mead 1993). Furthermore, the ancient plant samples were deposited over a period of more than 4000 years, which included major climate and ecosystem changes that may have had substantial impacts on the local nitrogen cycle, including (but not limited to) the extinction of megaherbivores, carnivores, and huge predatory birds that would have added nitrogen to soil and promoted N losses through deposition of urine, feces, and decomposing flesh. As a result, it would be difficult—if not impossible—to collect modern plant samples that could be used as reasonable analogues to the Escalante River Basin samples. In this paper, we compare our ancient plant δ15N values with previously published modern values, recognizing the limitations of this comparison.
The Escalante River Basin ancient plant δ15N values are within the ranges previously reported for modern plants from individual regions as diverse as southern Yukon and east Africa (Cerling et al. 2009; Koch et al. 1991; McLauchlan et al. 2010; Szpak et al. 2013; Tahmasebi et al. 2017). For example, within a limited region of southern Yukon, Tahmasebi et al. (2017) reported modern plant δ15N values ranging from about − 10 to + 27‰. The range of ancient C3 plant δ15N values in the present study (− 12 to + 14‰) is smaller that this modern range, despite the long temporal span and the diverse environments within the Escalante River Basin. Furthermore, Tahmasebi et al. (2017) showed that modern plants can have intra-plant differences (e.g., leaf vs. stem) of up to 12‰, so large ranges might be expected for various plant parts from multiple species over a long period of environmental change. Arid environments such as the Escalante River Basin have highly variable nitrogen dynamics and tend to have high rates of gaseous N losses, which are associated with large fractionations that can enrich soils and plants in 15N (Evans and Ehleringer 1994; Schaeffer and Evans 2005). In and of itself, the range of plant δ15N values measured for the ERB samples does not indicate post-depositional alteration.
Ancient CAM plants had significantly higher δ15N values than ancient C3 plants, even when the high-15N CAM outlier was excluded (t = 7.1, df = 36, p < 0001). Several studies have reported higher δ15N in modern CAM versus C3 plants in the same regions (Codron et al. 2005, 2006; Koch et al. 1991; Muzuka 1999), though the magnitude of the difference in these studies (e.g., + 13.2 vs. + 9.8‰ in Amboseli National Park, + 7.7 vs. + 3.4‰ in Waterberg National Park) is smaller than that observed for the ancient plants in the present study (+ 11.7 vs. + 4.7‰). Nevertheless, the higher values in ancient succulents could reflect original plant values.
The mean δ15N values for ancient Escalante River Basin plants were relatively high (+ 4.7‰ for C3 and + 11.7‰ for CAM) compared to modern leaf δ15N values of semi-arid woodland plants in south-central Utah, which averaged + 0.9 to + 2.7‰ (Evans and Ehleringer 1994), or to foliar δ15N values of hanging-garden plants (relatively wet environments) in modern Glen Canyon, which averaged − 2.3‰ (Flanagan et al. 1997). However, when the C3 plants are separated by time period, we see that the high δ15N values predominantly occurred during the Late Pleistocene (mean = + 7.2‰) (Table 4 and Fig. 3). The largest range for δ15N occurs in the Early Holocene—perhaps because of the “transitional” nature of this period—and only low δ15N values occur in the Terminal Early Holocene (Fig. 3). Furthermore, the Early Holocene and Terminal Early Holocene mean δ15N values (+ 0.1 and − 2.2‰, respectively) are comparable to those of modern plants in south-central Utah (Evans and Ehleringer 1994; Flanagan et al. 1997). It seems that either the Late Pleistocene plant δ15N values systematically increased after deposition whereas the Early Holocene values did not, or that neither were significantly altered and the Late Pleistocene to Holocene shift in plant δ15N reflects an underlying change in the ecosystems themselves.
Major changes in nitrogen cycling occurred at the end of the Pleistocene. A global study of 86 lakes on 6 continents recently showed a significant decrease (~ 2‰ on average) in lake-sediment δ15N values between about 15,000 and 7000 cal yr. bp (McLauchlan et al. 2013). The authors suggested that the decrease represented a major change in the global nitrogen cycle associated with increased terrestrial carbon sequestration, which consequently decreased nitrogen availability relative to biotic demand, leading to less open nitrogen cycles and reduced fractionation during processes associated with N loss. In a recent synthesis paper, Rabanus-Wallace et al. (2017) similarly showed a large decrease in megaherbivore δ15N in western North America at the end of the Pleistocene, followed by an increase during the Holocene. Although they interpreted these δ15N variations as resulting from moisture changes, the megafaunal extinctions themselves likely had significant effects on biogeochemical cycling, though these effects are not fully understood (Doughty et al. 2016; Gill 2014). Megaherbivores exert a disproportionate influence on their ecosystem relative to the same biomass of smaller herbivores (du Toit and Owen-Smith 1989; Owen-Smith 1988). Like modern bison, Pleistocene megaherbivores would have promoted nitrogen cycling and increased spatial heterogeneity in nitrogen availability through activities such as selective grazing/browsing, depositing large quantities of urine and feces, and returning nutrients to soils and plants through carcass decomposition (Doughty et al. 2013; Knapp et al. 1999), which may have led to larger nitrogen isotope fractionations and hence higher δ15N values during the Pleistocene. The loss of megafauna in the Late Pleistocene may also have allowed for the proliferation of cryptobiotic crusts, layers of cyanobacteria, lichens, and mosses that fix nitrogen (Evans and Ehleringer 1993), leading to lower soil and plant δ15N values in arid regions throughout the Holocene.
The results above suggest that the ancient plant δ15N values need not be rejected as altered simply because they seem relatively high. However, there is one line of isotopic evidence that does call these results into question: the plant δ15N values are higher than what would be expected based on our (small) sample of Pleistocene herbivore collagen δ15N. Given a collagen-diet offset of about 3‰ for these mammals (Caut et al. 2009), we would expect our three Glen Canyon Pleistocene herbivores to have been consuming plants with mean δ15N values of about + 2 to + 5‰. Since no ancient grasses were analyzed in this study, the most reasonable expected mean plant estimate might be derived from the browser (shrub-ox), whose δ15N value produces the lower plant δ15N estimate of about + 2‰. This is considerably lower than the measured mean plant δ15N for the Pleistocene samples (+ 7.2‰), and therefore suggests the possibility that the Late Pleistocene plant δ15N values may have systematically increased after deposition. On the other hand, this line of reasoning has critical limitations: the small sample size of herbivore bone collagen, the assumption that the ancient plant remains analyzed in this study are representative of the plants that comprised the shrub-ox diet, and the assumed 3‰ diet-tissue fractionation—which in reality can be quite variable (Caut et al. 2009).
Why Did DeNiro and Hastorf’s (1985) Peruvian Plants Have Such High δ 15N Values?
One thing is clear: the results for the ancient Escalante River Basin plants differ from those of DeNiro and Hastorf (1985), who found incredibly high δ15N values (many between + 20 and + 46‰) for uncharred plants from coastal regions of Peru. The results of the present study suggest that these high values cannot necessarily be considered representative of uncharred plants from other contexts. One possible explanation for DeNiro and Hastorf’s results is that their uncharred archaeological plants were fertilized with seabird guano, which can produce extremely high plant δ15N values (Szpak et al. 2012b). Indeed, recent studies have suggested that guano fertilization contributed to high δ15N values in ancient human bone collagen recovered from the Atacama Desert (Santana-Sagredo et al. 2015, 2017). It is notable that DeNiro and Hastorf’s uncharred plants were obtained from Peruvian coastal desert environments whereas their charred plants were found at wetter inland sites—very different depositional contexts with potentially different ancient land management practices. Guano was likely highly accessible at some of these coastal sites, and its impact on the samples studied by DeNiro and Hastorf remains to be further explored.
Conclusions
Ancient plants from the Escalante River Basin had δ13C values that were consistent with good preservation based on several lines of evidence: separation between plants utilizing different photosynthetic pathways (C3 vs. CAM), differences among plant parts (higher values in leaves vs. seeds), and species environmental preferences (oak vs. maple). Furthermore, the δ13C values of the ancient plants were consistent with expected diet values estimated from local Pleistocene herbivore bone and tooth collagen. This evidence suggests that ancient uncharred plant δ13C values did not undergo significant post-depositional alteration.
The evidence for preservation of original plant δ15N values is equivocal. Ancient plant δ15N values could not be rejected as altered based on their range relative to modern plants. CAM plants had higher δ15N values than C3 plants, as expected based on modern studies. Temporal changes in plant δ15N were in a similar direction to—though of a larger magnitude than—global changes in sedimentary δ15N related to shifts in the nitrogen cycle. In the Escalante River Basin, the mean δ15N values of Early Holocene plants were similar to those of modern plants in south-central Utah, whereas the Late Pleistocene plant values were much higher. Thus, it seems that either Late Pleistocene plants experienced systematic isotopic alteration while Early Holocene plants did not, or neither were significantly altered and the temporal change reflects an underlying change in the ecosystem itself. Late Pleistocene plant δ15N values were higher than expected based on Pleistocene herbivore collagen δ15N, which suggests the possibility that the plants experienced decay-related 15N enrichment. However, this line of reasoning relies on a small sample size and several critical assumptions that remain uncertain.
In contrast to DeNiro and Hastorf (1985), we did not find exceptionally high δ15N values (> + 25‰) in our ancient uncharred plants. The possibility that the coastal Peruvian plants had such high δ15N values because of fertilization with seabird guano should be further explored. The role of megaherbivores on nutrient cycling and variations in soil, plant, and animal δ15N—particularly in arid regions—also requires further investigation.
Plant δ15N values are important tools for understanding nitrogen fluxes in modern ecosystems, which can help predict the effects of anthropogenic climate change (e.g., Liu et al. 2017; Robinson 2001). Aside from their use in paleodietary research, ancient plant isotopic compositions could provide an important long-term record of nitrogen fluxes, if we can work out the taphonomic processes associated with their deposition and alteration (or lack thereof). %C, %N, and C/N ratios hold little promise for evaluating isotopic alterations of ancient plants, but the use of FTIR or other structural characterization techniques to identify isotopically altered samples and to characterize decay processes should be assessed in larger and more systematic studies. This study suggests that archaeologists and paleontologists should carefully consider the use of ancient uncharred plants for paleoenvironmental and paleodietary research.
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Acknowledgments
We thank Janet and David Gillette (Museum of Northern Arizona) and John Spence (National Park Service) for access to samples, Paul Szpak and Ed Grant for valuable discussions, Michael Richards and Fred Longstaffe for laboratory facilities and equipment, and Reba Macdonald, Christina Cheung, Joe Hepburn, and Grace Yau for laboratory assistance. Paul Szpak and three anonymous reviewers provided valuable comments on earlier drafts of this manuscript. Funding for this study was provided by a Killam Postdoctoral Research Fellowship, a Social Science and Humanities Research Council of Canada Banting Postdoctoral Research Fellowship, and the Larry D. Agenbroad Legacy Fund for Research from The Mammoth site of Hot Springs South Dakota.
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Metcalfe, J.Z., Mead, J.I. Do Uncharred Plants Preserve Original Carbon and Nitrogen Isotope Compositions?. J Archaeol Method Theory 26, 844–872 (2019). https://doi.org/10.1007/s10816-018-9390-2
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DOI: https://doi.org/10.1007/s10816-018-9390-2