Stable Carbon Isotopes of C3 Plant Resins and Ambers Record Changes in Atmospheric Oxygen Since the Triassic

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Geochimica et Cosmochimica Acta 121 (2013) 240–262 www.elsevier.com/locate/gca

Stable carbon isotopes of C3 plant resins and ambers record changes in atmospheric oxygen since the Triassic

Ralf Tappert a,b,⇑, Ryan C. McKellar a,c, Alexander P. Wolfe a, Michelle C. Tappert a, Jaime Ortega-Blanco c,d, Karlis Muehlenbachs a

a Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta, Canada b Institute of Mineralogy and Petrography, University of Innsbruck, Innsbruck, Austria c Division of Entomology (Paleoentomology), Natural History Museum, University of Kansas, Lawrence, KS, USA d Departament d’Estratigraﬁa, Paleontologia i Geocie`ncies Marines, Facultat de Geologia, Universitat de Barcelona, Barcelona, Spain

Received 28 September 2012; accepted in revised form 10 July 2013; Available online 24 July 2013

Abstract

Estimating the partial pressure of atmospheric oxygen (pO2) in the geological past has been challenging because of the lack of 13 reliable proxies. Here we develop a technique to estimate paleo-pO2 using the stable carbon isotope composition (d C) of plant resins—including amber, copal, and resinite—from a wide range of localities and ages (Triassic to modern). Plant resins are particularly suitable as proxies because their highly cross-linked terpenoid structures allow the preservation of pristine d13Csignatures over geological timescales. The distribution of d13C values of modern resins (n = 126) indicates that (a) resin-producing plant families generally have a similar fractionation behavior during resin biosynthesis, and (b) the fractionation observed in resins is similar to that of bulk plant matter. Resins exhibit a natural variability in d13C of around 8& (d13Crange:31& to 23&,mean: 27&), which is caused by local environmental and ecological factors (e.g., water availability, water composition, light exposure, temperature, nutrient availability). To minimize the effects of local conditions and to determine long-term changes in the d13Cof 13 13 resin 13 resins, we used mean d C values (d Cmean) for each geological resin deposit. Fossil resins (n = 412) are generally enriched in C 13 resin compared to their modern counterparts, with shifts in d Cmean of up to 6&. These isotopic shifts follow distinctive trends through time, which are unrelated to post-depositional processes including polymerization and diagenesis. The most enriched fossil resin 13 resin samples, with a d Cmean between 22& and 21&, formed during the Triassic, the mid-Cretaceous, and the early Eocene. Exper- 13 imental evidence and theoretical considerations suggest that neither change in pCO2 nor in the d C of atmospheric CO2 can 13 resin 13 13 account for the observed shifts in d Cmean. The fractionation of C in resin-producing plants (D C), instead, is primarily influ- 13 resin enced by atmospheric pO2, with more fractionation occurring at higher pO2. The enriched d Cmean values suggest that atmospheric pO2 during most of the Mesozoic and Cenozoic was considerably lower (pO2 = 10–20%) than today (pO2 = 21%). In addition, a 13 resin 18 correlation between the d Cmean and the marine d O record implies that pO2, pCO2, and global temperatures were inversely linked, which suggests that intervals of low pO2 were generally accompanied by high pCO2 and elevated global temperatures. Inter- vals with the lowest inferred pO2, including the mid-Cretaceous and the early Eocene, were preceded by large-scale volcanism during the emplacement of large igneous provinces (LIPs). This suggests that the influx of mantle-derived volcanic CO2 triggered an initial phase of warming, which led to an increase in oxidative weathering, thereby further increasing greenhouse forcing. This process resulted in the rapid decline of atmospheric pO2 during the mid-Cretaceous and the early Eocene greenhouse periods. After the cessation in LIP volcanism and the decrease in oxidative weathering rates, atmospheric pO2 levels continuously increased over tens of millions of years, whereas CO2 levels and temperatures continuously declined. These findings suggest that atmospheric pO2 had a considerable impact on the evolution of the climate on Earth, and that the d13C of fossil resins can be used as a novel tool to assess the changes of atmospheric compositions since the emergence of resin-producing plants in the Paleozoic. Ó 2013 Elsevier Ltd. All rights reserved.

⇑ Corresponding author at: Institute of Mineralogy and Petrography, University of Innsbruck, Innsbruck, Austria. Tel.: +43 512 492 5519. E-mail address: [email protected] (R. Tappert).

1. INTRODUCTION Epstein, 1960; Brugnoli and Farquhar, 2000). In addition, some plant groups use distinctive metabolic Our knowledge of the composition of the atmosphere pathways that fundamentally affect their isotopic through time and of its impact on climate change and composition. For example, C4 and CAM plants are the evolution of life is based on evidence preserved in generally more enriched in 13C compared to C3 the fossil record. For the Pleistocene and Holocene, a di- plants (O’Leary, 1988; Farquhar et al., 1989). There- rect and continuous record of the partial pressures and fore, if stable isotopes are used in chemostratigraphic isotopic compositions of atmospheric gases is preserved studies, it is crucial to determine the organic compo- in the form of air vesicles entrapped in polar ice sheets sition of the material under investigation and its (e.g., Petit et al., 1999). With increasing age, the record botanical source. of atmospheric compositions becomes less directly accessi- (2) Environmental and ecological factors can play an ble, and inferences about the composition of the atmo- important role in the isotopic fractionation of plants, sphere in the geological past rely on indirect measures and individual isotope values of fossil or modern or proxies. The quality of such information depends di- plant material may reflect local instead of large-scale rectly on the ability of proxies to reliably capture the com- environmental conditions. position of ancient atmospheres. (3) Plant constituents can undergo substantial composi- Attempts have been made to reconstruct atmospheric tional changes during and after burial and diagenesis. pCO2 and pO2 using mass balance calculations and flux These changes have the potential to alter the isotopic modeling (Berner, 1999; Berner and Kothavala, 2001; Ber- composition of the fossil plant material, to the extent ner et al., 2007). Other approaches to estimate the paleo- of preventing a meaningful paleoenvironmental pCO2 have utilized the density of stomata on fossil leaf cuti- interpretation. cles (Retallack, 2001, 2002; Beerling and Royer, 2002) and the carbon stable isotopic composition of pedogenic car- To determine environmentally mediated changes in the bonates (Cerling and Hay, 1986; Ekart et al., 1999; Bowen d13C of fossil plant organic matter through time, we ana- and Beerling, 2004). One of the richest sources of informa- lyzed modern and fossil resins from a wide range of locali- tion on ancient atmospheric compositions comes from the ties and ages. Compared to other terrestrial plant marine sedimentary record, in particular from the stable constituents, resins have chemical properties that make isotope composition (d13C, d18O, d11B) of biogenic carbon- them particularly suitable as proxies of environmental ates (e.g., Clarke and Jenkyns, 1999; Pearson and Palmer, changes over geological time. Resins can polymerize into 2000; Royer et al., 2001; Zachos et al., 2001). This informa- highly cross-linked hydrocarbons that have considerable tion has been used to infer changes in the isotopic compo- preservation potential in the geological record (Fig. 1). This sition of atmospheric CO2 through time and to derive means that a reliable terrestrial stable isotope history can paleo-temperature estimates. More recently, the d13Cof potentially be reconstructed well into the Paleozoic, consid- marine phytoplankton biomarkers have been used to esti- ering that the earliest known fossil resins are from the Car- mate the pCO2 of Cenozoic atmospheres (Pagani et al., boniferous (White, 1914; Bray and Anderson, 2009). 2005). Additional constraints on the composition of ancient Unlike many other fossil plant organic constituents, atmospheres were obtained from the marine record using resins can often be assigned to a specific source plant fam- the biochemical cycles of sulfur (Berner, 2006; Halevy ily, and only a few plant families produce significant et al., 2012). amounts of preservable resins, all of which have C3 For terrestrial environments, attempts to estimate paleo- metabolism. Furthermore, resins are chemically conserva- 13 pCO2 have frequently exploited the d C of fossil organic tive, which means that the composition of the resins did matter, primarily plant organic matter (Strauss and Pe- not change significantly as plants evolved. These factors ters-Kottig, 2003; Jahren et al., 2008). A range of plant con- limit not only the chemical variability of fossil resins, stituents have been targeted for d13C analysis, including but also any potential family-specific biases, which could wood, foliage, and cellulose extracts, as well as the coals influence the isotopic composition of fossil resins in to which they contribute (Stuiver and Braziunas, 1987; chemostratigraphic studies. Gro¨cke, 1998, 1999, 2002; Lu¨cke et al., 1999; Peters-Kottig Although large samples of gem-quality fossil resins, or et al., 2006; Poole and Van Bergen, 2006; Poole et al., 2006; amber, are found in fewer than 50 locations worldwide, Bechtel et al., 2008). However, three key issues need to be small pieces (<5 mm) of non-precious resin, which are addressed when plant organic matter is used for stable car- sometimes referred to as resinites, are found frequently, bon isotopic analyses in geochemical and chemostrati- and are often associated with coal deposits (Stach et al., graphic studies: 1982). Previous studies of the carbon isotopic composition of resins were limited because they were conducted on only (1) Plants are composed of a complex range of organic a few localities, and in most cases, using only a few samples compounds, including polysaccharides (e.g., cellu- (Nissenbaum and Yakir, 1995; Murray et al., 1998; Nissen- lose), lignin, lipids, cuticular waxes, and photosynth- baum et al., 2005; McKellar et al., 2008; DalCorso et al., ates. Due to differential fractionation of carbon 2011). Our study is the first attempt to integrate the d13C isotopes during biosynthesis, the isotopic composi- data of resins from a wide range of localities and ages, tion of individual plant fractions typically differ from and to determine their potential as proxies for changes in each other and from that of the bulk plant (Park and atmosphere composition through geological time. 242 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

A B

D E

Fig. 1. Examples of modern and fossil resins. (A) Fresh resin oozing out of the trunk of Araucaria cunninghamii, Adelaide Botanical Garden, Adelaide, Australia. Field of view (FOV): 20 cm. (B) In situ lenticular resinite fragment (arrow) within a coal seam, Horseshoe Canyon Formation (Campanian), Edmonton, Canada (forceps for scale). (C) Selection of amber samples from the Horseshoe Canyon Formation (Campanian), Drumheller, Canada. The largest specimen is 5 cm. (D) Inclusions of foliage from Parataxodium sp. (an extinct cupressaceous conifer) in amber, Foremost Formation (Campanian), Grassy Lake, Canada. FOV: 1 cm. (E) Triassic resin droplets in carbonate matrix, Heiligkreuzkofel Formation (Carnian), Dolomites, Italy. FOV: 10 cm.

2. SAMPLES AND METHODS old), which are referred to as copal. The sample set includes specimens from many well-known deposits (e.g., Baltic am- In total, 538 d13C measurements were conducted on ber, Bitterfeld amber, Dominican amber, Myanmar or Bur- modern (n = 126) and fossil resins (n = 412) from various mese amber, and Lebanese amber), but also material from locations worldwide, ranging in age from recent to Triassic newly discovered occurrences (Table 1). (Fig. 2, Table 1). Analytical work was performed on hard- For many amber deposits, the age of amber formation ened resins that were visibly free of inclusions and vesicles. can be tightly constrained (±2 Ma or better) by indepen- Modern resins were collected primarily from the trunks of dent stratigraphic controls or radiometric dating (Table 1). native trees growing under a range of climatic conditions However, for some of the deposits, the age constraints are (i.e., tropical to subarctic), although specimens from culti- more provisional, especially for localities that show evi- vars were also analyzed in some cases. We restricted the dence of re-working or re-deposition (e.g., Baltic and Bitter- sampling of modern resins to representatives of the families feld ambers). For these deposits, the age estimates have Pinaceae (Pseudotsuga, Picea, Pinus), Cupressaceae (Meta- relatively large uncertainties (±4 Ma). sequoia, Thuja), Araucariaceae (Agathis), and Fabaceae To examine the chemical composition of the resins, rep- (Hymenaea), as these families represent the most important resentative samples of modern and fossil resins from each resin producers in modern environments and the fossil re- locality were analyzed using micro-Fourier transform infra- cord (Langenheim, 2003). red (FTIR) spectroscopy. The resulting absorption spectra The fossil resin samples include many small, non-pre- were used to (1) classify the resins into distinct composi- cious varieties, in addition to genuine ambers. To simplify tional groups (see below), and (2) determine the botanical the terminology, all fossil resins are referred to herein as aﬃnity of the fossil resins. Particular emphasis was placed amber, with the exception of sub-recent resins (>200 years on identifying the botanical source of fossil resins from R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 243

Resin Type Age pinaceous Neogene cupressaceous-araucarian Paleogene fabaceous Late Cretaceous dipterocarpaceous Early Cretaceous Triassic

Fig. 2. Locations of amber deposits and types of amber analyzed in this study.

previously undescribed localities. Absorption spectra were dominant terpenoid structures, but they can also be influ- collected between 4000 and 650 cm1 (wavenumbers) using enced by the presence of additional organic components a Thermo Nicolet Nexus 470 FTIR spectrometer equipped and the degree of polymerization. with a Nicolet Continuum IR microscope, with a spectral Four compositional types of resins are distinguished in resolution of 4 cm1. A total of 200 interferograms were this study. These resin types were produced by specific plant collected and co-added for each spectrum. Spectra were col- families and are, accordingly, classified into the following lected from inclusion-free, freshly broken resin chips placed four categories: on an infrared-transparent NaCl disc. To avoid the effects of oversaturation of the spectra, sample thickness was kept (1) Pinaceous resins are exclusively produced by conifers below 5 lm. Depending on the quality of the sample, the belonging to the family Pinaceae. The infrared spec- spot size was set to values between 50 50 and tra of pinaceous resins are characterized by distinc- 100 100 lm (square aperture). Further details concerning tive absorption peaks at 1460 and 1385 cm1 the FTIR analytical procedures are presented in Tappert (Fig. 3). The amplitude of the peak at 1460 cm1 is et al. (2011). generally lower than the peak at 1385 cm1. Another The stable carbon isotope compositions of the resins were trait is the absence or the weak development of a determined using a Finnigan MAT 252 dual inlet mass spec- peak at 2848 cm1. This peak is typically well devel- trometer at the University of Alberta. Samples of untreated oped in cupressaceous-araucarian resins (see below). resin (1–5 mg) were combusted together with 1g CuO Further details about the spectral characteristics of as an oxygen source in a sealed and evacuated quartz tube pinaceous resins can be found in Tappert et al. at 800 °Cfor12 h. The d13C analyses were normalized to (2011). Based on results of gas chromatography– NBS-18 and NBS-19, and the results are given with respect mass spectrometry (GC–MS), pinaceous resins consist to V-PDB (Coplen et al., 1983). Instrumental precision and primarily of diterpenoids that are based on pimarane accuracy was on the order of ±0.02&.Repeatedanalyses and abietane structures (Langenheim, 2003). of individual samples were within ±0.1&. (2) Cupressaceous-araucarian resins are mainly produced by conifers of the families Cupressaceae and 3. RESULTS Araucariaceae, but also by Sciadopitys verticillata, the sole extant member of the family Sciadopytaceae 3.1. Infrared spectroscopy (Wolfe et al., 2009). The infrared spectra of cupressaceous-araucarian resins are characterized by distinc- The comparison of infrared spectra from modern and tive absorption peaks at 1448, 887, and 791 cm1 fossil resins (Fig. 3) shows that resins, irrespective of their (Fig. 3). In fossil resins, the 887 cm1 peak is often age and botanical origin, produce spectra that are very sim- reduced or absent as a result of the fossilization pro- ilar. This similarity is due to the dominance of diterpenoids cess. The amplitude of the absorption peak at in the composition of the resins. However, some variability 1385 cm1 is typically lower than that of the adjacent in the spectra exists and differences in the position and peak at 1448 cm1. Further details about the spectral height of specific absorption features can be used to distin- characteristics of cupressaceous-araucarian resins and guish different resin types. These subtle spectroscopic differ- their distinction from pinaceous resins can be found ences primarily reflect variations in the structure of the in Tappert et al. (2011). Cupressaceous-araucarian 244 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

Table 1 Overview of the fossil resins analyzed in this study, including their age and botanical source. No.Sample Location Stratigraphic Age/Epoch AbsoluteCompositional ClassaBotanical References to age setting of formation Age Type (based on source and/or botanical [Ma] FTIR source spectroscopy) 1 Colombian Colombia Undetermined Pliocene– 0.0002– Fabalean Ic Hymenaea sp. Ragazzi et al. (2003), copal Holocene 2.5 Lambert et al. (1995) 2 Malaysian Borneo Merit-Pila coal Middle 12 DipterocarpaceousII DipterocarpaceaeLangenheim and Beck amber field Miocene (1965), Schlee and Chan (1992) 3 Mexican Chiapas La Quinta Fm.– Early Miocene 13–19 Fabalean Ic Hymenaea sp. Langenheim and Beck amber Balumtun Fm. (1965), Cunningham et al., 1983, Solo´rzano Kraemer (2007) 4 Dominican Dominican La Toca Fm.– Late Oligocene– 15–20 Fabalean Ic Hymenaea sp. Langenheim and Beck amber Rep. Yanigua Fm. Early Miocene (1965), Iturralde-Vinent and MacPhee (1996) 5 Baltic Baltic Blaue Erde Fm. Late Eocene– 33–40 Cupressaceous- Ia Sciadopitys sp. Langenheim (2003), amber coast (redeposited) Early Oligocene araucariansuc Wolfe et al. (2009) 6 Bitterfeld Germany Bitterfeld coal Late Eocene– 33–40 Cupressaceous- Ia Sciadopitys sp. Fuhrmann (2005), amber field (redeposited)Early Oligocene araucariansuc Wolfe et al. (2009) 7 Giraffe Canada Peat sequence in Eocene 39–38 Cupressaceous- Ib Metasequoia sp. Tappert et al. (2011) kimberlite (NT) kimberlite crater araucarian amber 8 Tiger USA (WA) Tiger Mountain Middle Eocene 47–48 Cupressaceous- Ib Metasequoia sp. Mustoe (1985) Mountain Fm. araucarian amber 9 Ellesmere Canada Eureka Sound Middle Eocene 50 PinaceousSUC V Pseudolarix sp. Francis (1988), this study Island amber(NU) Fm. 10 Tadkeshwar India Cambay shale Eocene 50–52 DipterocarpaceousII DipterocarpaceaeRust et al. (2010) amber 11 Panda Canada Kimberlite crater Early Eocene 53 Cupressaceous- Ib Metasequoia sp. Wolfe et al. (2012) kimberlite (NT) infill araucarian amber 12 Alaskan USA (AK) Chickaloon Fm. Paleocene–Early 53–55 Cupressaceous- Ib Metasequoia sp. Triplehorn et al. (1984), amber Eocene araucarian this study 13 Evansburg Canada Upper Scollard Paleocene 59–62 Cupressaceous- Ib Metasequoia sp. This study amber (AB) Fm. araucarian 14 Genesee Canada Scollard Fm. Early Paleocene 65 Cupressaceous- Ib Metasequoia sp. This study amber (AB) araucarian 15 Albertan Canada Horseshoe Campanian– 70–73 Cupressaceous- Ib Parataxodium sp.McKellar and Wolfe amber (AB) Canyon Fm. Maastrichtian araucarian (2010) 16 Grassy LakeCanada Foremost Fm. Campanian 76–78 Cupressaceous- Ib Parataxodium sp.McKellar and Wolfe amber (AB) araucarian (2010) 17 Cedar Lake Canada Foremost Fm. Campanian 76–78 Cupressaceous- Ib Parataxodium sp.McKellar and Wolfe amber (MT) (redeposited) araucarian (2010) 18 New Jersey USA (NJ) Raritan Fm. Turonian 90–94 Cupressaceous- Ib Cupressaceae Grimaldi et al. (1989), amber araucarian Anderson (2006) 19 Burmese Myanmar Hukawng Basin Late Albian– 96–102 Cupressaceous- Ib Cupressaceae Grimaldi et al. (2002), amber Early araucarian Cruickshank and Ko Cenomanian (2003) 20 Spanish Spain, Eschucha Fm. Albian 100–110 Cupressaceous- Ib CheirolepidiaceaeAlonso et al. (2000) amber various araucarian 21 Lebanese Lebanon Gre`s de Base Fm.Late Barremian–123–127 Cupressaceous- Ib Cupressaceae Nissenbaum and Horowitz amber Early Aptian araucarian (1992), Azar et al. (2010) 22 Wealden UK (Isle Hastings beds Berriasian– 138–142 Cupressaceous- Ib Cupressaceae Nicholas et al. (1993) amber of Wight) (Wealden Group)Valanginian araucarian 23 Italian Italy Heiligkreuzkofel Carnian 220–225 Cupressaceous- Ib CheirolepidiaceaeRoghi et al. (2006), amber (Dolomites) Fm. araucarian Ragazzi et al. (2003) SUC: contains succinic acid esters (succinate). a Classification based on Anderson et al. (1992) and Anderson and Botto (1993). R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 245

resins consist primarily of diterpenoids that are based Although pinaceous resins are produced in extremely on labdane skeletal structures, such as communic high abundance in northern-hemisphere forests, their pres- acid polymers (Thomas, 1970; Anderson et al., 1992). ervation potential in the fossil record is limited. The main (3) Fabalean resins are produced by the tropical angio- constituent of pinaceous resins—diterpenes with pimarane sperm Hymenaea (Fabaceae, Fabales). The infrared and abietane skeletal structures—do not possess the func- spectra of these resins superficially resemble spectra tional groups required to form stable polymers. Therefore, from pinaceous and cupressaceous-araucarian repre- they are prone to decomposition. Nevertheless, under sentatives (Fig. 3). Similar to pinaceous resins, faba- favorable conditions, even pinaceous resins can persist for lean resins produce an absorption peak at 1460 cm1, millions of years (Wolfe et al., 2009). but the amplitude of the 1385 cm1 peak is lower or In accordance with previous studies (Langenheim and equivalent to the 1460 cm1 peak. A distinctive fea- Beck, 1965; Broughton, 1974; Tappert et al., 2011), all res- ture of fabalean resin spectra is the lack of strong ins of a given type produce very similar FTIR absorption absorption features between 970 and 1050 cm1. spectra, irrespective of whether the samples are modern Fabalean resins also produce an absorption peak at or fossil, which indicates that chemical transformations 887 cm1. Like cupressaceous-araucarian resins, during burial and maturation are minor. This observation fabalean resins are primarily based on labdane struc- is important because it supports the premise that resins tured diterpenoids, with zanzibaric and ozic acid can resist isotopic exchange over geological timescales. polymers being typical components (Cunningham et al., 1983). 3.2. Carbon stable isotopes (4) Dipterocarpaceous resins are produced by members of the angiosperm family Dipterocarpaceae, which 3.2.1. Modern resins is a family of diverse and widespread tropical rainfor- The d13C values of modern resins analyzed in this study est trees. The infrared spectra of dipterocarpaceous range from 31.2& to 23.6&. These values are consistent resins are quite distinct from other resin spectra with previously published d13C analyses of modern resins (Fig. 3). The most characteristic spectroscopic fea- from comparable plant taxa (Nissenbaum and Yakir, tures are the distinctive absorption triplet between 1995; Murray et al., 1998; Nissenbaum et al., 2005). The 1360 and 1400 cm1, and an additional absorption d13C values of the modern resins follow a near Gaussian peak at 1050 cm1. Dipterocarpaceous resins are distribution, with a mean value of 26.9& (Fig. 4). It is primarily based on sesquiterpenoids, such as cadin- notable that resins from individual plant families have sim- ene, but triterpenoids can also be present (Ander- ilar d13C distributions and similar mean d13C values son et al., 1992). The resin produced by (Fig. 5). This indicates that the isotopic fractionation that dipterocarpaceous trees is colloquially referred to as occurs during resin biosynthesis is nearly identical for the Dammar. plant families considered in this study, despite differences in the terpenoid profile of individual resin types. Some of the resins, including Baltic amber and resins Fig. 4 also shows the carbon isotopic compositions of from Ellesmere Island, were found to produce an additional bulk leaf material from modern C3-plants using data com- and distinctive spectral feature at 1160 cm1, which is asso- piled by Ko¨rner et al. (1991), Diefendorf et al. (2010), and ciated with a broad shoulder or a broad peak between 1200 Kohn (2010). The leaf material was sampled from a wide and 1300 cm1 (Fig. 3). These spectral features generally re- range of plant taxa growing under diverse climatic condi- late to the presence of esterified succinic acid (succinate) as tions, from alpine and polar to tropical. Therefore, the iso- an additional component in the resin (Beck et al., 1965; topic composition should approximate the average Beck, 1986). Although succinate has been identified in some composition of modern C3 plant organic matter. However, pinaceous resins—namely in resins from Pseudolarix amabi- samples from extremely dry environments and samples that lis and in some fossil resins of Pseudolarix-affinity from the were potentially influenced by the canopy effect (see below) Canadian Arctic (Anderson and LePage, 1995; Poulin and were excluded. It is notable that the isotope values of the Helwig, 2012)—it is the hallmark constituent of Baltic am- leaf material (32.8& to 22.2&, mean: 27.1&) and ber, which itself is a cupressaceous-araucarian resin. The the distribution of isotope values are very similar to those presence of succinate has important implications in the reported for modern resins. In conclusion, we find no con- assessment of the botanical source of fossil resins, and consistent differences between the distributions of d13C values sequently, succinate-bearing resins are marked separately in in modern leaves and modern resins. Table 1. Most of the fossil resins analyzed for this study are of 3.2.2. Fossil resins cupressaceous-araucarian type. Along with fabalean res- Carbon isotope values of fossil resins were found to ins (e.g., Dominican and Mexican ambers), cupressa- range from 28.4& to 18.2&, which means that some ceous-araucarian resins make up the most productive fossil resins are more enriched in 13C compared to their commercial amber deposits (Table 1). The predominance modern counterparts. Data for individual samples are of these resins in the fossil record is due to their terpe- shown in Fig. 6. Similar to modern resins, the fossil resins noid composition because labdane-based resins polymer- at each deposit produce a range of d13C values. These ize rapidly and are highly resistant to chemical ranges are similar to those observed for modern resins breakdown (Fig. 1B). (i.e., 8&) provided that a comparable number of samples 246 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

Single bonds (x-H) Double bonds Single bonds/Skeletal vibrations O-H C-H C=C, C=O C-O, C-C

2848 1160 1460 1448 1050 791 1385 887 succinate region Pinaceous resins

Pseudolarix amabilis -modern-

Ellesmere Island amber (Canada) -Middle Eocene-

Cupressaceous- araucarian resins

Wollemia nobilis -modern-

l Tiger Mountain aber (USA)

c -Middle Eocene-

s Genesee amber (Canada)

f -Early Paleocene-

(

c Fabalean n resins a Hymenaea courbaril

b r -modern-

A) Mexican amber -Early Miocene-

Dominican amber -Late Oligocene-Early Miocene-

Dipterocarpaceous resins

Shorea rubriflora -modern-

Malaysian amber -Middle Miocene-

3600 3400 3200 3000 2800 2600 2400 2200 2000 1800 1600 1400 1200 1000 800 Wavenumbers (cm-1 )

Fig. 3. Micro-FTIR spectra of different resin types distinguished in this study (selected modern and fossil examples). Dashed lines mark spectral features discussed in the text. Yellow backgrounds highlight the most distinctive spectral regions. Characteristic spectral regions for molecular vibrations relevant to resins are marked in blue. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.) were analyzed. For some of the amber deposits, only a To quantify the isotopic shift through time, and to mini- small number of samples were analyzed due to sample mize the effect of sample size, we used the mean d13C values 13 resin availability. Consequently, the range of isotope values for of the fossil resins (d Cmean) as a representative measure for these deposits tends to be narrower. Despite differences in each deposit (Table 2). 13 resin the number of samples for different deposits, a considerable The d Cmean of subrecent Colombian copal (27.5&)is shift in d13C between deposits is observed (Figs. 6 and 7). nearly identical to modern resins, but with increasing age, R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 247

Fractionation efficiency high low Growth conditions (interpreted) optimal typical poor

A Modern resins mean: -26.9‰ = 1.39 n = 160

Frequency

B Bulk C3 leaf mean = -27.1‰ = 1.67 n = 891

Frequency

13C (‰ vs PDB)

Fig. 4. Distribution of d13C values in (A) modern resins (data from this study with additional data of Hymenaea resins from Nissenbaum et al., 2005), and (B) bulk leaf matter from a wide range of C3 plant taxa (data from Ko¨rner et al., 1991; Diefendorf et al., 2010; Kohn, 2010). resins become continuously more enriched in 13C and con- The early Eocene is characterized by a dramatic shift to- 13 resin sequently less negative in their d Cmean values. This trend wards more depleted resin isotope compositions, as demon- towards more enriched isotopic compositions in fossil res- strated by amber from the Tadkeshwar coal ﬁeld of India 13 resin ins continues through the Neogene and into the middle Eo- (d Cmean = 25.8&), and amber preserved in the Panda 13 resin cene. Some of the most productive amber deposits, kimberlite pipe in Canada (d Cmean = 24.8&). The earli- 13 resin including the Dominican (d Cmean = 24.8&) and the Bal- est Eocene amber from the Chickaloon Formation in Alas- 13 resin 13 tic (d Cmean = 23.5&) amber deposits, formed during this ka marks a return to slightly more C-enriched 13 resin time interval. Eocene ambers from Tiger Mountain (Wash- compositions, with a d Cmeanof 23.6&. 13 resin ington, USA) and Ellesmere Island (Nunavut, Canada) are Paleocene amber from Evansburg (d Cmean ¼22:3&) 13 13 resin more than 5.3& enriched in C compared to modern res- and Genesee (d Cmean ¼23:1&) in Alberta, Canada, 13 resin ins, with a d Cmean of 21.7& and 21.6&, respectively. are more enriched than their early Eocene counterparts. The latter are the isotopically most enriched Cenozoic res- However, from the Paleocene into the Late Cretaceous ins in the present sample set. (Campanian), resin compositions become increasingly more 248 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

35 Pinaceae 30 (mean=-26.9‰, =1.12, n=86) Cupressaceae-Araucariaceae (mean=-26.6‰, =1.77, n=50) 25 Fabaceae (angiosperm) (mean=-27.3‰, =1.28, n=24) 20

Frequency 15

0 -31 -30 -29 -28 -27 -26 -25 -24 -23 13C (‰ vs PDB)

Fig. 5. Distribution of d13C values from different types of modern resins, separated according to their source plant families. Data include previously published analyses of Hymenaea (Fabaceae) resins (n = 17) from Nissenbaum et al. (2005). depleted in 13C. The most depleted Cretaceous resins with a of parasites, etc. (Park and Epstein, 1960; Guy et al., 13 resin d Cmean of 23.9& were found in the Horseshoe Canyon 1980; O’Leary, 1981; Ko¨rner et al., 1991; Arens and Jahren, Formation (late Campanian) of southern and central Al- 2000; Edwards et al., 2000; Dawson et al., 2002; Gro¨cke, berta. From the late Campanian and into the Early Creta- 2002; McKellar et al., 2011)(Fig. 8). Each of these factors 13 resin ceous, this trend reverses and the d Cmean shifts to more has a direct influence on the efficiency of fractionation in 13C-enriched compositions, as shown by the coeval Grassy plants during photosynthesis, and each can cause an isoto- 13 resin 13 Lake and Cedar Lake amber (d Cmean ¼23:6&), and the pic shift on the order of several permil in the d Cofa 13 resin New Jersey Raritan amber (d Cmean ¼22:1&)(Fig. 7). plant. Resins that formed between the Cenomanian and Barre- Every plant community will produce a range of d13C val- 13 resin mian are characterized by even more enriched d Cmean ues because the environmental conditions (i.e., growth con- compositions. During this time interval, Burmese amber ditions) for each plant are unique. However, each plant 13 resin 13 resin (d Cmean ¼21:3&), Spanish amber (d Cmean ¼ taxon has its own range of growth conditions under which 13 resin 13 21:4&), and Lebanese amber (d Cmean ¼21:1&) were C fractionation is maximized. The observation that the formed. The oldest Cretaceous amber samples were recov- d13C of modern resins follows a Gaussian distribution ered from the Wealden Group of England (UK). Although (Fig. 4), suggests that the majority of the resin-producing 13 resin the d Cmean (22.8&) of the Wealden amber is more de- plants grew under conditions that were intermediate be- pleted than that of the Burmese, Spanish, and Lebanese tween optimal (i.e., most efficient fractionation) and poor 13 resin amber, its d Cmean is not well constrained because only (i.e., least efficient fractionation). Therefore, the relative two samples were available for analysis. d13C values of plant matter within a local plant community Pre-Cretaceous ambers were limited to samples from the can be viewed primarily as a measure of the growth Triassic (Carnian) Heiligkreuzkofel Formation of the Dol- conditions. 13 resin omites, Italy. With a d Cmean of 20.9&, these ambers are For most plants, including the copious resin producers, 13 the most C-enriched in the current sample set. In accor- it can be assumed that the metabolized CO2 was sourced dance with previously published d13C values of amber from from isotopically undisturbed air, which had a d13C compo- 13 resin the same locality (d Cmean ¼20:12&; DalCorso et al., sition approximating a global atmospheric average. How- 13 2011), our results confirm that these Triassic resins ever, local isotopic effects can alter the d C of the CO2 in are > 6& more 13C-enriched relative to modern resins. the ambient air. Limited air circulation, for example, can lead to a reprocessing of respired air, which is depleted in 4. DISCUSSION 13C(i.e., canopy effect). Metabolizing this depleted air causes the d13C of plants to shift towards more negative val- 4.1. Influences on the d13C of modern plants ues. In extreme cases, it can result in a depletion in 13Cof plant matter of >6& (Medina and Minchin, 1980). The 4.1.1. Local environmental factors canopy effect primarily affect plants in the understory of Experimental studies, including growth chamber experi- dense tropical rainforests. Among the plant families consid- ments, have shown that plant d13C is influenced by a wide ered in this study, however, only the Dipterocarpaceae range of local environmental factors, such as water avail- grow in such dense rainforest environments, and only the ability, water composition (e.g., salinity, nutrients), light Miocene amber from Borneo and the Eocene amber from exposure, local temperature, altitude, humidity, presence Tadkeshwar, India, are of dipterocarpacean origin. R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 249

13C (‰vs VPDB) -32 -30 -28 -26 -24 -22 -20 -18 -16

Pinaceae

Cupressaceae-Araucariaceae modern

Fabaceae

Columbian copal

Malaysian amber

Mexican amber

Dominican amber

Baltic amber

Bitterfeld amber

Giraffe kimberlite amber

Tiger Mountain amber

Ellesmere Island amber

Tadkeshwar amber

Panda kimberlite amber

Alaskan amber

Evansburg amber

Genesee amber

Albertan amber

Cedar Lake amber

Grassy Lake amber

New Jersey amber

Burmese amber

Spanish amber

Lebanese amber

Wealden amber

Italian amber

Fig. 6. d13C of modern and fossil resins analyzed for this study. Modern resin data are separated according to their botanical source, i.e., resin type. Data from fossil resins (grey) are shown for individual deposits. The data are arranged according to the relative age of the fossil resins (see Table 1). Black rectangles represent one standard deviation. Black squares mark mean values and vertical bars median values.

4.1.2. Regional and global environmental factors (MAP > 2000 mm/year), in which fractionation is in- The results of recent studies on the global distribution of creased. For plants that grew in environments with an inter- d13C values in bulk leaf matter indicate that the fraction- mediate MAP, average D13C values only show a weak ation of 13C in plants is not only influenced by local envi- correlation with MAP. Since most of the resin-producing ronmental factors, but also by regional precipitation plant families investigated here grow typically in environ- patterns (Diefendorf et al., 2010; Kohn, 2010; Fig. 8). It ments with intermediate MAP, the effect of MAP on their was found that the average d13C of bulk leaf matter and D13C can be considered minor. This notion is supported 13 13 resin the calculated fractionation values (D C) show some corre- by the observation that the d Cmean values of resins from lation with the mean annual precipitation (MAP) at a given different modern plant families investigated in this study location. The 13C fractionation thereby increases with are very similar. However, the Dipterocarpaceae represent increasing MAP (Diefendorf et al., 2010; Kohn, 2010). an exception, because they commonly grow in areas with However, the correlation is strongly influenced by plants high MAP. Therefore, they are predisposed towards in- that were sampled from very dry environments creased 13C fractionation. This ecological preference may (MAP < 500 mm/year), in which 13C fractionation is re- explain the low d13C values of dipterocarpaceous resins ob- duced, and those sampled from tropical rainforests served by Murray et al. (1998). 250 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

Pinaceae Modern Resins Cupressaceae/Araucariaceae Fabaceae Age (Ma) Pleistocene Columbian copal Pliocene

10 e g Malysian amber o Mioocene

N Mexican amber 20 Dominican amber

Oligocene 30 Baltic/Bitterfeld ambers

e Giraffe kimberlite amber 40 n

o Eocene

l Tiger Mountain amber a Ellesmere Island amber 50 P Tadkeshwar amber Panda kimberlite amber Alaskan amber

60 Paleocene Evansburg amber

Genesee amber Maastrichtian 70 Albertan amber

u Grassy Lake/Cedar Lake ambers

o Campanian

80 c

r Santonian

C Coniacian

90 t

a Turonian New Jersey amber

L Cenomanian Burmese amber 100

Albian Spanish amber

110

a Aptian 120 t

C Lebanese amber

r Barremian

130 a

E Hauterivian

Valanginian 140 Wealden amber Berriasian

c 220 i Italian amber

s Carnian

T 230 Ladinian

-28 -26 -24 -22 -20 13C (‰vs PDB)

Fig. 7. Temporal variations in the mean d13C of fossil resins (amber). Horizontal bars mark one standard deviation (see Fig. 6). Modern resin values are shown for comparison. Diﬀerent resin types are distinguished by color. Pinaceous: red; cupressaceous-araucarian: blue; fabalean: yellow; dipterocarpaceous: green. (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article.)

13 13 resin The d C of atmospheric CO2 (d Cmean) has a direct im- alline carbonates (Swart et al., 2010) indicate that the global 13 13 resin pact on the d C of plant organic matter, and analyses of average d Cmean has evolved by at least 1& (from around gas inclusions in ice cores (Francey et al., 1999) and of cor- 6.5& to <7.5&) over the last 200 years. This rapid R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 251

Table 2 4.1.3. Compositional factors Mean d13C values of fossil resins from individual deposits, In addition to the isotopic variability that is caused by including standard deviations, and number of samples analyzed environmental factors, some plant compounds are isotopi- (arranged by age). cally distinct from bulk plant matter. Cellulose, for exam- No Sample d13Cmean (&) r (&) n ple, is slightly more enriched in 13C(2–3&), whereas 1 Columbian copal 27.5 0.6 6 lignin is slightly more depleted (3–4&) compared to the 2 Malaysian amber 26.1 0.3 12 bulk plant (Park and Epstein, 1960; Benner et al., 1987; 3 Mexican amber 25.5 0.9 15 Schleser et al., 1999; Brugnoli and Farquhar, 2000) 4 Dominican amber 24.8 1.6 32 (Fig. 9). The fractionation that causes the isotopic deviation 5 Baltic amber 23.5 1.0 37 of these plant compounds, consequently, must occur after 6 Bitterfeld amber 23.7 1.0 15 the initial photosynthetic fixation of CO2 by ribulose-1,5- 7 Giraffe kimberlite amber 22.5 1.9 23 bisphosphate carboxylase oxygenase (Rubisco). Unlike cel- 8 Tiger Mountain amber 21.7 0.4 12 lulose and lignin, the d13C values of plant resins overlap 9 Ellesmere Island amber 21.6 1.5 12 with those of primary plant metabolites, such as leaf sugars 10 Tadkeshwar amber 25.8 0.8 10 11 Panda kimberlite amber 24.8 0.2 5 and bulk leaf tissues (Fig. 9), which indicates that fraction- 12 Alaskan amber 23.6 1.2 15 ation during resin biosynthesis (i.e., after the initial CO2 fix- 13 Evansburg amber 22.3 0.8 14 ation) is minor, typically < 1& (Diefendorf et al., 2012). 14 Genesee amber 23.1 2.0 14 This conclusion is important because it suggests that 13 resin 13 bulk plant 13 resin 15 Albertan amber 23.9 1.8 39 d Cmean d Cmean , and shifts in d Cmean reflect shifts 16 Grassy Lake amber 23.7 1.4 34 13 bulk plant in the d Cmean of C3 plants through time. 17 Cedar Lake amber 23.2 1.4 12 The observation that the d13C of individual plant com- 18 New Jersey amber 22.1 1.1 36 pounds can deviate from the d13C of the bulk plant, needs 19 Burmese amber 21.3 1.0 6 to be considered whenever non-specific plants remains are 20 Spanish amber 21.4 1.5 21 used in chemostratigraphic studies, since differential preser- 21 Lebanese amber 21.1 1.2 21 22 Wealden amber 22.8 0.5 2 vation of the individual compounds can result in consider- 13 13 23 Italian amber 20.9 0.7 19 able shifts in d C. For example, mean d C values of bulk organic matter (i.e., coal, coalified tissues, and cuticles) from the Cretaceous and earliest Cenozoic (Strauss and Pe- 13 resin ters-Kottig, 2003) are 2.5& lower compared to d Cmean

13 values from the same time interval. This difference can be Global C effect readily explained by the preferential preservation of 13C-de- 13 13 increase pleted lignin and the breakdown of C-enriched cellulose Cair shift 13 resin p increase in fossil wood, from which the d C is unaffected. CO2 shift mean p increase O2 shift 4.2. Causes for the shift in 13Cresin through time Regional mean mean annual precipitation increase shift The different types of modern resins investigated in this 13 Local study show a very similar distribution of d C values (Fig. 5), which indicates that the structure of their terpenoid local 13C increase air shift 13 (e.g., canopy effect) constituents has little or no effect on their d C. This conclusion, however, contradicts claims that systematic differ- water availability ences in the fractionation behavior between gymnosperms water composition (salinity etc.) humidity and angiosperms exist, with plant matter from angio- 13 light exposure combined effect: sperms, including resins, generally being enriched in C local temperature ~7-8‰variability & nutrient availability by 2–3 (Diefendorf et al., 2012). Although it is conceiv- altitude able that certain plant taxa are prone to isotopic biases, due presence of parasites to their ecological preferences (e.g., dipterocarpaceous trees etc. growing in areas with high MAP and closed canopy condi- Fig. 8. Overview of global, regional, and local factors and their tions), there is no conclusive evidence that the fractionation effect on the d13C of plant organic matter. behavior of C3 plants is affected by their phylogenetic position. 13 resin change in d Cmean is generally linked to the increasing in- The tendency of fossil resins to be isotopically more en- 13 flux of anthropogenic CO2 from fossil fuel combustion, riched in C compared to modern resins could be inter- which is depleted in 13C (Suess effect). Although the post- preted as diagenetic overprinting. After exudation, resins 13 resin industrial decrease in d Cmean has been detected in the generally lose a fraction of their terpenoid constituents— d13C of bulk plant material (e.g., tree rings; Loader et al., mainly volatile mono- and sesquiterpenes—in addition to 2003), it was not possible to identify a similar shift in water, whereas the non-volatile fractions polymerize rap- 13 resin d Cmean between modern and pre-industrial resins (i.e., idly. Once initial polymerization has occurred, any Colombian copal), possibly due to the relatively small num- additional maturation of the resins during fossilization ber of copal samples analyzed. causes only minor structural changes (Lambert et al., 252 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

2008; Tappert et al., 2011). These changes are unlikely to (DeConto et al., 2012), and the rapid burning of Paleocene influence the d13C of the resins significantly. In fact, if an terrestrial carbon (Kurtz et al., 2003). Since the effects of isotopic exchange had taken place during diagenesis, fossil these processes are considered to be short-lived, they are resins from different localities should show an erratic distri- unlikely to be responsible for the more gradual decline in 13 resin 13 bution of the d Cmean arising from different diagenetic con- the d C between the late Paleocene and early Eocene. ditions at each locality, but this is not observed. Therefore, other, more fundamental processes, such as Furthermore, if resins displayed a tendency to become changes in the rate of oxidation of organic matter, must unstable and prone to carbon isotopic exchange through be invoked as driving forces for this d13C decline. time, they should become continuously more enriched in Despite the exceptional d13C excursions in the early 13C with increasing age, which is also not the case. Diage- Paleogene, the shifts in d13C of benthic foraminifera netic resetting should also diminish the natural variability through the Cenozoic are only moderate (2&) even at in d13C of resins, and this would result in samples from their extremes. In fact, between the Middle Eocene and individual localities converging on similar d13C values in- the Miocene, d13C values of marine carbonates remain rel- stead of retaining their original range of d13C values. atively constant, with only minor shifts of 61&. The shifts 13 resin As previously mentioned, an important factor that influ- in d Cmean over the same time interval are approximately ences the d13C of plant organic matter, including resins, is 6&, and they record a long-term trend towards more neg- the carbon isotopic composition of the CO in the atmo- ative values (Fig. 11). In summary, changes in d13Catm may 2 CO2 sphere (d13Catm ). A shift in d13Catm , in fact, would result influence the d13C of plant organic matter to some extent, CO2 CO2 in a comparable shift in the d13C of plant organic matter. but their influence is insufficient to account for the magni- Although changes in d13Catm have been invoked to explain tude of the observed shifts in d13Cresin . CO2 mean 13 13 resin the d C variations observed in fossil plant organic matter Further comparison of the d Cmean with the marine sta- (Arens et al., 2000; Jahren, 2002; Strauss and Peters-Kottig, ble isotope record shows that for much of the Cenozoic, 13 resin 2003; Hasegawa et al., 2003; Jahren et al., 2008), subse- negative shifts in d Cmean broadly correlate with positive quent studies have questioned whether changes in d13Catm shifts of a similar magnitude in the d18O of marine carbon- CO2 alone can be responsible the observed d13C variability ates (Fig. 11). The only pronounced exception is the late (e.g., Gro¨cke, 2002; Beerling and Royer, 2002; Lomax Paleocene to early Eocene isotopic decline, which resulted 13 resin 18 et al., 2012; Schubert and Jahren, 2012). If changes in in a shift in both d Cmean and d O to more negative values. d13Catm were to be the primary cause for the d13C variabil- The marine d18O record is primarily a temperature record; CO2 13 resin 18 ity of fossil plant organic matter, shifts in d Cmean should with less negative d O values reflecting higher average 13 resin correlate with independent proxies for d Cmean, including ocean surface temperatures and consequently higher global those that are based on the d13C of marine carbonates temperatures (Emiliani, 1955; Bemis et al., 1998). The cor- 18 13 resin (e.g., calcareous tests of benthic foraminifera) (Zachos relation between the marine d O and the d Cmean record, et al., 2001; Tipple et al., 2010). However, the comparison therefore, indicates that resins that formed during intervals of the marine d13C record with the resin record shows very of high global temperatures tend to be more enriched in little similarity (Fig. 11). A notable exception is the pro- 13C. Since high global temperatures are generally linked 13 nounced decline in d C values, which started in the late to high atmospheric pCO2 (Royer, 2006), variations in Paleocene and lasted well into the early Eocene (57– atmospheric pCO2 seem to be an obvious alternative to ex- 52 Ma, Fig. 11). This decline, which coincides with times plain shifts in the d13C of plant matter, including resins. of peak Cenozoic temperatures (i.e., Early Eocene climatic The results of recent growth experiments under controlled optimum, EECO), is discernible in both the marine and the pCO2 and water/moisture availability suggest a hyperbolic in- 13 13 resin d C record and it is only interrupted by short-lived crease of plant fractionation (D C) with increasing pCO2 negative excursions in d13C and d18O(e.g., Paleocene Eo- (Schubert and Jahren, 2012). These results are consistent with cene Thermal Maximum, PETM) (Zachos et al., 2008). most previous studies on 13C fractionation in plants, which 13 The cause for the rapid change in atmosphere composi- also reported positive correlations between D CandpCO2 tion during the PETM and other similar short-lived (Park and Epstein, 1960; Beerling and Woodward, 1993; Tre- (<20 ka) events remains a matter of debate (McInerney ydte et al., 2009). However, some studies found negative cor- and Wing, 2011), but the observation that these events relations or no correlations (Beerling, 1996; Jahren et al., are associated with negative excursions in d13C and d18O 2008). The discrepancy between these studies is most likely 13 indicates that a considerable amount of C-depleted CO2 the result of differences in the control of the growth conditions was added to the atmosphere during these intervals, causing (i.e., water availability, light exposure, etc.) because variations peak warming (hyperthermals). Proposed mechanisms to in the growth conditions can mask the isotopic effect caused 13 create such a rapid influx of C-depleted CO2 into the by increasing pCO2. atmosphere include the dissociation and oxidation of meth- Irrespective of the extent of the pCO2 impact, the obser- ane hydrates in marine sediments (Sloan et al., 1992; Dick- vation that 13C fractionation in plants increases with 13 ens et al., 1995, 1997), the release of thermogenic CH4 from increasing pCO2 cannot account for the d C shifts in the sediments through magmatism and its subsequent oxida- resin record (Fig. 7). In the resin record, the least depleted 13 resin tion (Svensen et al., 2004; Westerhold et al., 2009), an in- d Cmean compositions occur precisely during intervals with crease in oxidation of organic matter caused by the predicted high atmospheric pCO2. We conclude, once drying of epicontinental seas (Higgins and Schrag, 2006), again, that some other process must be responsible for 13 resin the release and oxidation of CH4 stored in polar permafrost the shifts in d Cmean. R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 253

Pectin Asparatate Aminoacids Hemicellulose Cellulose Starch Sap sugars Leaf sugars Proteins -carotene Isoprene Lignin Fatty acids Total lipids Bulk leaf Resin (this study) -15 -20 -25 -30 -35 13C (vs PDB)

Fig. 9. Ranges in d13C of diﬀerent plant compounds, modiﬁed from Brugnoli and Farquhar (2000). The d13C range for resin is based on the analyses of modern resins in this study; the range for bulk leaf matter is based on the compilations from Ko¨rner et al. (1991), Diefendorf et al. (2010), and Kohn (2010) (Fig. 4).

13 C=pCO2 0 Stomatal density GEOCARB III Pedogenic carb.

)

(

p Atmospheric O2 (%)

13 Fig. 10. Comparison of calculated pO2 values using diﬀerent models to quantify the impact of pCO2 on plant fractionation (dD CpCO2 ). 13 13 Values for dD CpCO2 were calculated based on the relationship between pCO2 and D C proposed by Schubert and Jahren (2012),in combination with pCO2 estimates from (a) the density of stomata on fossil leaf cuticles (Retallack, 2001), (b) mass-balance modeling 13 (GEOCARB III; Berner and Kothavala, 2001), and (c) the d C of pedogenic carbonates (Ekart et al., 1999). pO2 values calculated for 13 13 dD CpCO2 ¼ 0 (no eﬀect of pCO2 on D C) are shown for comparison.

Another important factor of the atmosphere that influ- mass that is isotopically more depleted in 13C(Berner ences the isotopic fractionation in plants is the oxygen con- et al., 2000). However, the opposite effect (i.e., a decrease tent. Preliminary experimental studies on C3 plants indicate in fractionation) has been observed under lower than that fractionation of carbon during photosynthesis will in- modern pO2 in ambient air (Beerling et al., 2002). This crease if the pO2 in the ambient air is significantly higher means that plants growing under low O2 conditions will 13 than modern values (pO2 = 21%), resulting in plant bio- be more enriched in C than those growing under high 254 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

O2 conditions. The strong influence of O2 on the carbon factor, in this case, is not necessary. A direct relationship isotope fractionation in C3 plants is consistent with theoret- may not be extended to extremely low values ical considerations of leaf-gas exchange processes and the (pO2 < 10%), but we believe a direct relationship is a rea- fixation of CO2 by plants. During photosynthesis, atmo- sonable assumption for moderate pO2 levels as long as spheric CO2 is converted in a carboxylation reaction physiological adaptations of plants (e.g., changes in metab- through the enzyme Rubisco to energy-rich metabolites, olism) do not come into play. In our model, paleo-pO2 at such as glucose. At the same time, atmospheric O2 com- the time of resin formation can be estimated as: petes with CO for reaction sites on Rubisco. The reaction 2 D13C dD13C of Rubisco with O (oxygenation) partially offsets photo- pO ¼ resin pCO2 pOmodern ð2Þ 2 2 D13C 2 synthesis and results in the release of CO2 in the process 13 13 of photorespiration. With higher atmospheric pO2, more where D Cresin denotes the C fractionation of resin-bear- oxygen is used through photorespiration, which leads to ing plants at the time of resin formation, and D13C repre- an increase of CO2 within the plant cell and hence an in- sents the carbon isotope fractionation of plants in the crease in 13C fractionation. The relationship between 13C- current atmosphere 13 modern modern fractionation in plants (D C) and pCO2 has been described (pO2 ¼ 21%; pCO2 ¼ 0:039%), which is approxi- 13 in general terms by Farquhar et al. (1982) as: mately 21&. Based on Farquhar et al. (1982), D Cresin

13 can be calculated as: D C ¼ a þðb aÞci=ca ð1Þ ðd13Catm d13CmeanÞ1000 13 13 CO2 resin where ci denotes the intercellular pCO2, and ca the pCO2 in D C ¼ ð3Þ resin 1000 d13Cmean the ambient atmosphere. The factors a and b represent þ resin fractionations that occur during diffusion through the sto- In this equation, d13Catm denotes the isotopic composi- CO2 mata (a = 4.4&) and during fixation of CO2 by Rubisco tion of CO2 in the ambient air. For modern plants, (b =27&), respectively. d13Catm can be directly measured. For samples from the CO2 13 13 atm The observed effect of atmospheric pO2 on the d Cof geological past, values for d C were calculated from 13 CO2 plant tissues means that the d C of plant organic matter the d13C and d18O of benthic foraminiferal tests, following 18 should be explored as a proxy for pO2 in ancient atmo- the approach of Tipple et al. (2010). The d O composition 13 spheres (paleo-oxybarometer). Given that the d C distribu- is thereby used to correct for the temperature-dependency tions of modern plant organic matter and resins are very of the fractionation between dissolved inorganic carbon 13 18 similar (Fig. 4), we conclude that amber has all the prere- and CO2. For the Cenozoic, the d C and d O data to cal- quisite qualities to be used as a pO -proxy. 13 atm 2 culate d CCO were taken from Zachos et al. (2001). For the 13 2 13 18 To estimate pO2 from plant d C it is necessary to quan- late and mid-Cretaceous, d C and d O values from ben- 13 tify the C fractionation in the plant at varying atmo- thic foraminifera were taken from Li and Keller (1999) (late spheric pO2. This information is typically derived from Campanian) and Huber et al. (1995) (Campanian–Albian). experimental work on plants in growth chambers under Due to the lack of high-quality stable isotope data from controlled pO2 conditions. However, only a few studies to benthic foraminifera for the Early Cretaceous and the Tri- date have attempted to systematically determine the influ- assic, we used average d13C and d18O values from compila- ence of pO2 on plant fractionation, and the results of these tions of Weissert and Erba (2004) and Korte et al. (2005), studies show large discrepancies (e.g., compare Smith et al., which are based on bulk carbonate rocks and carbonaceous 1976; Berry et al., 1972 and Beerling et al., 2002). These dis- macrofossils. Although this approach introduces some crepancies may also be the result of differences in the con- additional uncertainties, most of the error on the d13Catm CO2 trol of the experimental growth conditions, similar to calculations is due to uncertainties in the exact age place- those observed in the CO2 growth experiments, since even ment of the fossil resin samples. small variations in the growth conditions (water and nutri- 13 The factor dD CpCO2 in Eq. (2) denotes the change of 13 ent availability, light exposure etc.) can mask the effects of D C at varying pCO2 relative to the fractionation at mod- varying pO2. Irrespective of the cause of the discrepancies, ern pCO2. As previously mentioned, the influence of pCO2 none of the proposed relationships that are based on exper- on plant fractionation has recently been investigated in imental data (e.g., Berner et al., 2000; Beerling et al., 2002) growth experiments under systematically controlled growth 13 resin can account for the magnitude of the shifts in d Cmean.In conditions by Schubert and Jahren (2012). These authors fact, calculated pO2 values for most of the Cretaceous observed a hyperbolic relationship between the fraction- and early Paleocene would be close to 0% if the relationship ation in C3 plants and pCO2, and formulated a best-fit 13 between pO2 and D C proposed by Berner et al. (2000) and function, using their own and previously published data. Beerling et al. (2002) was applied to fossil resins. More real- Based on this best-fit function proposed by Schubert and istic pO2 values would require much larger empirical scaling 13 Jahren (2012), we calculated dD CpCO2 as: factors than those proposed by these authors. 28:26 0:21 pCO 25 Due to the inconsistencies in the experimental determi- 13 ð Þð Þð 2 þ Þ 13 dD CpCO2 ¼ D C ð4Þ nation of the 13C fractionation factor, we used a direct rela- 28:26 þð0:21ÞðpCO2 þ 25Þ 13 tionship between resin-derived D C and atmospheric pO2 whereby pCO2 is given in ppm. Although there is a general as a first order approximation. Under this assumption, consensus that the pCO2 for most of the Cenozoic and the fractionation effect from photorespiration is propor- Mesozoic was higher than at present, considerable dispute tional to the pO2 in the ambient air. The use of a scaling remains about the absolute pCO2 values in the geological R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 255

Igneous Climatic events Marine stable isotope record Fossil resins (mean value) Peercentage of O in atmospher 2 activity and long-term 13 18 13 C (‰) O (‰) C (‰) 10 15 20 (LIPs) T-trends -1 0 1 2 0 1 2 3 4 5 -28 -26 -24 -22 -20 Pleistocene Arctic ice sheets Pliocene forming 10 Mioocene Neogene 20 cooling

Oligocene 30 Antarctic ice sheets 40 forming Eocene

50 Paleogene EECO PETM North Atlantic

60 Paleocene Deccan Pg warming Maastrichtian K 70 12 8 4 0 Age (Ma) Campanian Ice-free temperature 80

Santonian cooling Coniacian 90 Turonian Late Cretaceous Cenomanian 100

Albian Mid-Cretaceous 11 0 greenhouse

120 Aptian Ontong Java

Barremian Parana-Etendeka warming 130 Early Cretaceous Hauterivian

Valanginian 140 Berriasian

220 Carnian Triassic 230 Ladinian 10 15 20 -28-2 6 -24 -22

13 resin Fig. 11. Variations in predicted atmospheric pO2 (red) and d Cmean (black) through time. The pO2 estimates represent mean values for 13 resin diﬀerent pCO2 models, with the outline indicating the extreme values (Fig. 10). Error bars on d Cmean mark one standard deviation. The marine stable isotope record for the Cenozoic is adopted from Zachos et al. (2001), and illustrates variations in d13C and d18O of benthic foraminifera. Major climatic events, long-term temperature trends, and periods of large-scale igneous activity are shown for comparison. (For interpretation of the references to color in this ﬁgure legend, the reader is referred to the web version of this article.)

13 resin past. To address the uncertainties in the paleo-pCO2 esti- tionation (i.e., higher d Cmean values) of resin-producing mates and their impact on the pO2 results, we calculated plants in the geological past. The particularly high 13 13 resin dD CpCO2 using pCO2 values derived by three independent d Cmean values of early to mid-Cretaceous amber samples approaches: (1) mass balance calculations (GEOCARB III, suggest that the pO2 at the time of their formation was as Berner and Kothavala, 2001), (2) the abundance of stomata low as 10–11% (Fig. 11). Such low pO2 levels are regarded on leaf surfaces (Retallack, 2001), and (3) the d13C of ped- to be below the limit of combustion and seem to contradict ogenic carbonates (Ekart et al., 1999). Despite considerable the presence of charcoal in Cretaceous sedimentary se- differences in the proposed pCO2 levels between these three quences (Jones and Chaloner, 1991; Belcher and McElwain, models, the calculated pO2 values show only minor devia- 2008). However, the presence of charcoal in sedimentary tions (<±1.5%) for most of the Cenozoic and Mesozoic. rocks itself does not preclude low atmospheric pO2, since Larger discrepancies (<±3%) are restricted to the Paleocene the charring process (i.e., pyrolysis) only occurs under and early Eocene (Fig. 10). This indicates that most of the low-O2 conditions. Furthermore, experiments to establish uncertainty in the calculated pO2 values rests on the accu- the lower limit of atmospheric pO2 required for combustion 13 racy of the relationship between D C and pCO2 as deter- did not consider certain factors that influence the combus- mined by Schubert and Jahren (2012). tion in nature, such as the effect of wind, which may lower the required pO2 considerably. 4.2.1. Mesozoic and Cenozoic pO2 evolution Despite some uncertainties about absolute pO2 values, A compilation of the calculated variations in pO2 our low-pO2 estimates for the mid-Cretaceous are consis- 13 resin through time is compared to d Cmean, the marine stable iso- tent with pO2 estimates based on the abundance and mass tope record, and important environmental events in Fig. 11. balance of carbon and pyritic sulfur in sedimentary rocks It is apparent from this compilation that for most of the during that same time interval (Berner and Canfield, Cenozoic and Mesozoic pO2 was considerably lower com- 1989; Berner, 2006, Fig. 12). Our results, on the other hand, 13 pared to today. This is consistent with a reduced C frac- question recently proposed paleo-pO2 estimates that are 256 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 based on the abundance of charcoal in the fossil record, zoic, as indicated by the exceptionally light d18O composi- which predict a higher-than-modern pO2 for the Cretaceous tion of marine carbonates deposited during that time and the Cenozoic (Glasspool and Scott, 2010; Brown et al., (Fig. 11). As previously mentioned, the time interval just 2012, Fig. 12). prior to the EECO (i.e., late Paleocene to early Eocene) 13 resin 13 As previously mentioned, shifts in the Cenozoic d Cmean was marked by a steady decline in the d C of marine car- 18 13 resin broadly resemble those in the marine d O record, with the bonates and in d Cmean, indicating an increased influx of 13 resin 13 exception of the negative d Cmean excursion in the early C-depleted CO2 into the atmosphere and an associated 13 atm Paleogene (Fig. 11). This indicates that changes in plant decrease in d CCO that lasted for 5 million years. In view 2 13 fractionation, i.e., changes that are linked to varying pO2, of our pO2 estimates, this continuous influx of C-depleted may also be linked to changes in global temperatures, and CO2 was most likely the result of an increase in the rate of thus pCO2. In fact, our pO2 estimates for the Cretaceous oxidation (oxidative weathering), caused by the rising glo- and Cenozoic show moderate to strong negative correla- bal temperatures and by elevated pO2 levels during that tions with pCO2 estimates derived by independent methods time interval. The increased rate of oxidative weathering (Fig. 13). These correlations suggest that times of low pO2, would have dramatically accelerated the oxidation of or- on a broad scale, were characterized by high pCO2 and high ganic matter and let to a further increase in pCO2 and global temperatures, and vice versa. greenhouse forcing. At the same time, it would have re- For the Early Cretaceous, it appears that an early phase sulted in a continuous depletion of O2 in the atmosphere, of low pO2 (12–13% O2) was followed by an interval of even which would explain the low pO2 following the EECO. lower pO2 levels (10–11% O2) during the Aptian to Turo- After the height of the early Eocene greenhouse phase, nian. This phase of extremely-low pO2 coincides with the atmospheric pO2 went into a long-term phase of continuous mid-Cretaceous greenhouse—a time interval characterized recovery until it reached modern levels. The recovery in pO2 by high atmospheric pCO2, high global temperatures, and was most likely the result of a decreased rate of oxidative extensive ocean anoxic events (Jenkyns, 1980; Huber weathering and an associated increase in carbon burial. et al., 1995; Wilson and Norris, 2001; Leckie et al., 2002). This phase was also characterized by a continuous decrease From the Turonian to the end of the Cretaceous, atmo- in pCO2 and global temperatures, which culminated in the spheric pO2 continuously rose to >15%, whereas pCO2 glacial inception in the southern and northern hemispheres and global temperatures declined. during the terminal Paleogene. The early Paleogene appears to have been characterized by moderate pO2 levels in the atmosphere (14–18% O2), 4.2.2. Volcanism and its effects on pO2 which declined to a Cenozoic minimum of 12% O2 imme- It is noteworthy that the intervals with lowest resin- diately after the EECO (Fig. 11). The EECO marks the time inferred pO2 during the Cretaceous and the Cenozoic of highest global temperatures and pCO2 during the Ceno- were preceded by episodes of intense volcanism and the

Fig. 12. Comparison of predicted atmospheric pO2 from this study with previously proposed models that are based on mass balance calculations (GEOCARBSULF; Berner, 2006) and the abundance of charcoal in the rock record (Glasspool and Scott, 2010). R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 257

Stomatal density GEOCARB III Pedogenic carbonates

(

p Atmospheric O2 (%)

Fig. 13. Plot of predicted atmospheric pO2 against pCO2 estimates based on the density of stomata on fossil leaf cuticles (Retallack, 2001), mass-balance modeling (GEOCARB III; Berner and Kothavala, 2001), and the d13C of pedogenic carbonates (Ekart et al., 1999). emplacement of some of the largest known igneous prov- prolonged greenhouse episodes that followed the LIP volca- inces (LIPs). These include the Ontong-Java Plateau and nism. In this scenario, the elevated pO2 that prevailed be- the Parana-Etendeka Province (Early Cretaceous), as well fore the onset of the LIP volcanism (i.e., in the Early and as the Deccan and North Atlantic Provinces (Late Creta- Late Cretaceous) would have been an important factor that ceous to Paleocene) (Fig. 11, Larson and Erba, 1999). determined the extent of the subsequent greenhouse phases. The emplacement of these LIPs was associated with the After the cessation of LIP volcanism and the end of the expulsion of vast amounts of mantle-derived volcanic gases, greenhouse periods in the Cretaceous and the Cenozoic, which are typically dominated by H2O and CO2, but also atmospheric pO2 was at its minimum, but slowly increased contain sulfur dioxide (SO2) and other trace gases (Sy- over the following tens of millions of years, as the result of a monds et al., 1994). The expulsion of such large quantities decrease in oxidative weathering (associated with an in- of volcanic gases, particularly of CO2, provides a valid crease in carbon burial) and the absence of the CO2-influx mechanism to cause a rapid increase in atmospheric pCO2 from LIP magmatism. After the mid-Cretaceous, the cycle and an associated increase of greenhouse forcing and global of increasing pO2 and decreasing pCO2 lasted for 40 mil- temperatures during these time intervals. Although there is lion years, but was interrupted by renewed LIP volcanism still some dispute over the long-term effect of volcanogenic at the end of the Cretaceous. The Cenozoic decline of CO2 emissions (Self et al., 2005), the hypothesis that LIP pCO2 and increase in pO2, on the other hand, continued magmatism contributed to changes in atmospheric compo- uninterrupted for 50 million years from the Middle Eo- sition and global climate is supported by the marine d13C cene into modern times. The prolonged absence of large- record, which indicates that the late Paleocene decline of scale volcanism during that time period may have been a 13 13 d C was preceded by a notable rise in d C values that prerequisite for pCO2 and global temperatures to decrease started in the mid-Paleocene. This rise in d13C coincides sufficiently to allow sea ice to form on the polar oceans, with phases of intense volcanism in the Deccan area of In- ultimately triggering the onset of glaciation in both hemi- dia, and in the North Atlantic Igneous Province (Fig. 11). spheres in the latest Paleogene. Marine carbonates deposited at that time reached d13C values of around +2& (Fig. 11; Zachos et al., 2001). These 4.2.3. Effects of pO2 on animal evolution unusually positive values are consistent with a strong influx Although it has been proposed that atmospheric pO2 has of mantle-derived CO2 into the atmosphere because mantle a profound effect on the physiology and evolution of ani- 13 CO2 is enriched in C relative to CO2 derived from organic mals, and that gigantism in animals is linked to high pO2 matter and has a rather restricted d13C range (8to3&, (Graham et al., 1995; Clapham and Karr, 2012), our data mean: 5&)(Exley et al., 1986; Javoy and Pineau, 1991). preclude such a link for the evolution of giant theropod Although the episodes of LIP volcanism in the Early and sauropod dinosaurs. Considering the low pO2 that pre- Cretaceous and the Late Cretaceous/Paleocene may have vailed during the Cretaceous and, at least, part of the Tri- not been sufficient to sustain long-term greenhouse condi- assic, it is more reasonable to assume that a high primary tions, they may have served as a trigger for the subsequent photosynthetic productivity (i.e., food availability) caused increase of oxidative weathering. This, in turn, would have by high pCO2 and high global temperatures was the key provided a feedback for the further rise in pCO2 and the factor in the evolution of dinosaurs and their tendency to 258 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 develop large body sizes. A similar conclusion can be drawn ACKNOWLEDGMENTS for the Eocene radiation of large placental mammals, which We are grateful to Jim Basinger, Madelaine Bo¨hme, Xavier has previously been linked to high pO2 levels as well (Fal- kowski et al., 2005). Based on our data, however, this time Delclo`s, Michael Engel, George Mustoe, Philip Perkins, Guido interval was also characterized by low pO . Roghi, Alexander Schmidt, and Chris Williams for generously pro- 2 viding samples used in this study. Thomas Stachel is thanked for providing access to the FTIR spectrometer housed in his labora- 5. CONCLUSIONS AND FUTURE DIRECTIONS tory. We also thank Stacey Gibb and Gabriela Gonzalez for assis- tance at various stages, and we would like to thank the three Due to their ability to polymerize into highly resistant anonymous reviewers, whose constructive comments have im- organic macromolecules that are demonstrably representa- proved this publication considerably. Much of this research was tive of whole-plant carbon fractionation, resins can pre- supported by the Natural Science and Engineering Research Coun- cil of Canada (Discovery awards to A.P. Wolfe and K. Muehlenb- serve pristine isotopic signatures over geological time. achs, Postdoctoral fellowship to R.C. McKellar). J. Ortega was This makes them ideal natural biomarkers for chemostrati- supported by the Ministerio de Economıá y Competitividad, Ful- graphic studies. As with all plant organic constituents, the bright Espanã, and FECYT. d13C of C3 plant resins is influenced by a range of local environmental and ecological factors, which result in considerable isotopic variability. The variability seen in resins REFERENCES is directly comparable to that of primary plant biomass, which suggests that the d13C of resins can be used as a mea- Alonso J., Arillo A., Barron E., Corral J. C., Grimalt J., Lopez J. sure of the d13C of bulk plant organic matter. To minimize F., Lopez R., Martinez-Delclos X., Ortuno V., Penalver E. and statistical biases caused by variations in the growth condi- Trincao P. R. 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