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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 (pO 2 ) in the geological past has been challenging because of the lack of reliable proxies. Here we develop a technique to estimate paleo-pO 2 using the stable carbon isotope composition (d 13 C) of plant resins—including amber, copal, and resinite—from a wide range of localities and ages (Triassic to modern). Plant resins are par- ticularly suitable as proxies because their highly cross-linked terpenoid structures allow the preservation of pristine d 13 C signatures over geological timescales. The distribution of d 13 C values of modern resins (n = 126) indicates that (a) resin-producing plant fam- ilies 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 d 13 C of around 8& (d 13 C range: 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 d 13 C of resins, we used mean d 13 C values (d 13 C resin mean ) for each geological resin deposit. Fossil resins (n = 412) are generally enriched in 13 C compared to their modern counterparts, with shifts in d 13 C resin mean 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 samples, with a d 13 C resin mean between 22& and 21&, formed during the Triassic, the mid-Cretaceous, and the early Eocene. Exper- imental evidence and theoretical considerations suggest that neither change in pCO 2 nor in the d 13 C of atmospheric CO 2 can account for the observed shifts in d 13 C resin mean . The fractionation of 13 C in resin-producing plants (D 13 C), instead, is primarily influ- enced by atmospheric pO 2 , with more fractionation occurring at higher pO 2 . The enriched d 13 C resin mean values suggest that atmospheric pO 2 during most of the Mesozoic and Cenozoic was considerably lower (pO 2 = 10–20%) than today (pO 2 = 21%). In addition, a correlation between the d 13 C resin mean and the marine d 18 O record implies that pO 2 , pCO 2 , and global temperatures were inversely linked, which suggests that intervals of low pO 2 were generally accompanied by high pCO 2 and elevated global temperatures. Inter- vals with the lowest inferred pO 2 , including the mid-Cretaceous and the early Eocene, were preceded by large-scale volcanism dur- ing the emplacement of large igneous provinces (LIPs). This suggests that the influx of mantle-derived volcanic CO 2 triggered an initial phase of warming, which led to an increase in oxidative weathering, thereby further increasing greenhouse forcing. This pro- cess resulted in the rapid decline of atmospheric pO 2 during the mid-Cretaceous and the early Eocene greenhouse periods. After the cessation in LIP volcanism and the decrease in oxidative weathering rates, atmospheric pO 2 levels continuously increased over tens of millions of years, whereas CO 2 levels and temperatures continuously declined. These findings suggest that atmospheric pO 2 had a considerable impact on the evolution of the climate on Earth, and that the d 13 C 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. 0016-7037/$ - see front matter Ó 2013 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.gca.2013.07.011 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). www.elsevier.com/locate/gca Available online at www.sciencedirect.com ScienceDirect Geochimica et Cosmochimica Acta 121 (2013) 240–262

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Available online at www.sciencedirect.com

www.elsevier.com/locate/gca

ScienceDirect

Geochimica et Cosmochimica Acta 121 (2013) 240–262

Stable carbon isotopes of C3 plant resins and ambersrecord 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, Canadab Institute of Mineralogy and Petrography, University of Innsbruck, Innsbruck, Austria

c Division of Entomology (Paleoentomology), Natural History Museum, University of Kansas, Lawrence, KS, USAd Departament d’Estratigrafia, Paleontologia i Geociencies 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 ofreliable proxies. Here we develop a technique to estimate paleo-pO2 using the stable carbon isotope composition (d13C) of plantresins—including amber, copal, and resinite—from a wide range of localities and ages (Triassic to modern). Plant resins are par-ticularly suitable as proxies because their highly cross-linked terpenoid structures allow the preservation of pristine d13C signaturesover geological timescales. The distribution of d13C values of modern resins (n = 126) indicates that (a) resin-producing plant fam-ilies generally have a similar fractionation behavior during resin biosynthesis, and (b) the fractionation observed in resins is similarto that of bulk plant matter. Resins exhibit a natural variability in d13C of around 8& (d13C range: �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 d13C ofresins, we used mean d13C values (d13Cresin

mean) for each geological resin deposit. Fossil resins (n = 412) are generally enriched in 13Ccompared to their modern counterparts, with shifts in d13Cresin

mean of up to 6&. These isotopic shifts follow distinctive trends throughtime, which are unrelated to post-depositional processes including polymerization and diagenesis. The most enriched fossil resinsamples, with a d13Cresin

mean between �22& and�21&, formed during the Triassic, the mid-Cretaceous, and the early Eocene. Exper-imental evidence and theoretical considerations suggest that neither change in pCO2 nor in the d13C of atmospheric CO2 canaccount for the observed shifts in d13Cresin

mean. The fractionation of 13C in resin-producing plants (D13C), instead, is primarily influ-enced by atmospheric pO2, with more fractionation occurring at higher pO2. The enriched d13Cresin

mean values suggest that atmosphericpO2 during most of the Mesozoic and Cenozoic was considerably lower (pO2 = 10–20%) than today (pO2 = 21%). In addition, acorrelation between the d13Cresin

mean and the marine d18O record implies that pO2, pCO2, and global temperatures were inverselylinked, 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 dur-ing the emplacement of large igneous provinces (LIPs). This suggests that the influx of mantle-derived volcanic CO2 triggered aninitial phase of warming, which led to an increase in oxidative weathering, thereby further increasing greenhouse forcing. This pro-cess resulted in the rapid decline of atmospheric pO2 during the mid-Cretaceous and the early Eocene greenhouse periods. After thecessation in LIP volcanism and the decrease in oxidative weathering rates, atmospheric pO2 levels continuously increased over tensof millions of years, whereas CO2 levels and temperatures continuously declined. These findings suggest that atmospheric pO2 hada considerable impact on the evolution of the climate on Earth, and that the d13C of fossil resins can be used as a novel tool toassess the changes of atmospheric compositions since the emergence of resin-producing plants in the Paleozoic.� 2013 Elsevier Ltd. All rights reserved.

0016-7037/$ - see front matter � 2013 Elsevier Ltd. All rights reserved.

http://dx.doi.org/10.1016/j.gca.2013.07.011

⇑ 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).

R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 241

1. INTRODUCTION

Our knowledge of the composition of the atmospherethrough time and of its impact on climate change andthe evolution of life is based on evidence preserved inthe fossil record. For the Pleistocene and Holocene, a di-rect and continuous record of the partial pressures andisotopic compositions of atmospheric gases is preservedin the form of air vesicles entrapped in polar ice sheets(e.g., Petit et al., 1999). With increasing age, the recordof atmospheric compositions becomes less directly accessi-ble, and inferences about the composition of the atmo-sphere in the geological past rely on indirect measuresor proxies. The quality of such information depends di-rectly on the ability of proxies to reliably capture the com-position of ancient atmospheres.

Attempts have been made to reconstruct atmosphericpCO2 and pO2 using mass balance calculations and fluxmodeling (Berner, 1999; Berner and Kothavala, 2001; Ber-ner et al., 2007). Other approaches to estimate the paleo-pCO2 have utilized the density of stomata on fossil leaf cuti-cles (Retallack, 2001, 2002; Beerling and Royer, 2002) andthe carbon stable isotopic composition of pedogenic car-bonates (Cerling and Hay, 1986; Ekart et al., 1999; Bowenand Beerling, 2004). One of the richest sources of informa-tion on ancient atmospheric compositions comes from themarine sedimentary record, in particular from the stableisotope composition (d13C, d18O, d11B) of biogenic carbon-ates (e.g., Clarke and Jenkyns, 1999; Pearson and Palmer,2000; Royer et al., 2001; Zachos et al., 2001). This informa-tion has been used to infer changes in the isotopic compo-sition of atmospheric CO2 through time and to derivepaleo-temperature estimates. More recently, the d13C ofmarine phytoplankton biomarkers have been used to esti-mate the pCO2 of Cenozoic atmospheres (Pagani et al.,2005). Additional constraints on the composition of ancientatmospheres were obtained from the marine record usingthe biochemical cycles of sulfur (Berner, 2006; Halevyet al., 2012).

For terrestrial environments, attempts to estimate paleo-pCO2 have frequently exploited the d13C of fossil organicmatter, primarily plant organic matter (Strauss and Pe-ters-Kottig, 2003; Jahren et al., 2008). A range of plant con-stituents have been targeted for d13C analysis, includingwood, foliage, and cellulose extracts, as well as the coalsto which they contribute (Stuiver and Braziunas, 1987;Grocke, 1998, 1999, 2002; Lucke et al., 1999; Peters-Kottiget al., 2006; Poole and Van Bergen, 2006; Poole et al., 2006;Bechtel et al., 2008). However, three key issues need to beaddressed when plant organic matter is used for stable car-bon isotopic analyses in geochemical and chemostrati-graphic studies:

(1) Plants are composed of a complex range of organiccompounds, including polysaccharides (e.g., cellu-lose), lignin, lipids, cuticular waxes, and photosynth-ates. Due to differential fractionation of carbonisotopes during biosynthesis, the isotopic composi-tion of individual plant fractions typically differ fromeach other and from that of the bulk plant (Park and

Epstein, 1960; Brugnoli and Farquhar, 2000). Inaddition, some plant groups use distinctive metabolicpathways that fundamentally affect their isotopiccomposition. For example, C4 and CAM plants aregenerally more enriched in 13C compared to C3plants (O’Leary, 1988; Farquhar et al., 1989). There-fore, if stable isotopes are used in chemostratigraphicstudies, it is crucial to determine the organic compo-sition of the material under investigation and itsbotanical source.

(2) Environmental and ecological factors can play animportant role in the isotopic fractionation of plants,and individual isotope values of fossil or modernplant material may reflect local instead of large-scaleenvironmental conditions.

(3) Plant constituents can undergo substantial composi-tional changes during and after burial and diagenesis.These changes have the potential to alter the isotopiccomposition of the fossil plant material, to the extentof preventing a meaningful paleoenvironmentalinterpretation.

To determine environmentally mediated changes in thed13C of fossil plant organic matter through time, we ana-lyzed modern and fossil resins from a wide range of locali-ties and ages. Compared to other terrestrial plantconstituents, resins have chemical properties that makethem particularly suitable as proxies of environmentalchanges over geological time. Resins can polymerize intohighly cross-linked hydrocarbons that have considerablepreservation potential in the geological record (Fig. 1). Thismeans that a reliable terrestrial stable isotope history canpotentially be reconstructed well into the Paleozoic, consid-ering that the earliest known fossil resins are from the Car-boniferous (White, 1914; Bray and Anderson, 2009).

Unlike many other fossil plant organic constituents,resins can often be assigned to a specific source plant fam-ily, and only a few plant families produce significantamounts of preservable resins, all of which have C3metabolism. Furthermore, resins are chemically conserva-tive, which means that the composition of the resins didnot change significantly as plants evolved. These factorslimit not only the chemical variability of fossil resins,but also any potential family-specific biases, which couldinfluence the isotopic composition of fossil resins inchemostratigraphic studies.

Although large samples of gem-quality fossil resins, oramber, are found in fewer than �50 locations worldwide,small pieces (<5 mm) of non-precious resin, which aresometimes referred to as resinites, are found frequently,and are often associated with coal deposits (Stach et al.,1982). Previous studies of the carbon isotopic compositionof resins were limited because they were conducted on onlya few localities, and in most cases, using only a few samples(Nissenbaum and Yakir, 1995; Murray et al., 1998; Nissen-baum et al., 2005; McKellar et al., 2008; DalCorso et al.,2011). Our study is the first attempt to integrate the d13Cdata of resins from a wide range of localities and ages,and to determine their potential as proxies for changes inatmosphere composition through geological time.

BA

C

ED

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 CanyonFormation (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 cupressaceousconifer) 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.

242 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

2. SAMPLES AND METHODS

In total, 538 d13C measurements were conducted onmodern (n = 126) and fossil resins (n = 412) from variouslocations worldwide, ranging in age from recent to Triassic(Fig. 2, Table 1). Analytical work was performed on hard-ened resins that were visibly free of inclusions and vesicles.Modern resins were collected primarily from the trunks ofnative trees growing under a range of climatic conditions(i.e., tropical to subarctic), although specimens from culti-vars were also analyzed in some cases. We restricted thesampling of modern resins to representatives of the familiesPinaceae (Pseudotsuga, Picea, Pinus), Cupressaceae (Meta-

sequoia, Thuja), Araucariaceae (Agathis), and Fabaceae(Hymenaea), as these families represent the most importantresin producers in modern environments and the fossil re-cord (Langenheim, 2003).

The fossil resin samples include many small, non-pre-cious varieties, in addition to genuine ambers. To simplifythe terminology, all fossil resins are referred to herein asamber, with the exception of sub-recent resins (>200 years

old), which are referred to as copal. The sample set includesspecimens from many well-known deposits (e.g., Baltic am-ber, Bitterfeld amber, Dominican amber, Myanmar or Bur-mese amber, and Lebanese amber), but also material fromnewly discovered occurrences (Table 1).

For many amber deposits, the age of amber formationcan be tightly constrained (±2 Ma or better) by indepen-dent stratigraphic controls or radiometric dating (Table 1).However, for some of the deposits, the age constraints aremore provisional, especially for localities that show evi-dence of re-working or re-deposition (e.g., Baltic and Bitter-feld ambers). For these deposits, the age estimates haverelatively large uncertainties (�±4 Ma).

To examine the chemical composition of the resins, rep-resentative samples of modern and fossil resins from eachlocality were analyzed using micro-Fourier transform infra-red (FTIR) spectroscopy. The resulting absorption spectrawere used to (1) classify the resins into distinct composi-tional groups (see below), and (2) determine the botanicalaffinity of the fossil resins. Particular emphasis was placedon identifying the botanical source of fossil resins from

pinaceouscupressaceous-araucarianfabaceousdipterocarpaceous

Resin TypeNeogenePaleogeneLate CretaceousEarly Cretaceous

Age

Triassic

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

R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 243

previously undescribed localities. Absorption spectra werecollected between 4000 and 650 cm�1 (wavenumbers) usinga Thermo Nicolet Nexus 470 FTIR spectrometer equippedwith a Nicolet Continuum IR microscope, with a spectralresolution of 4 cm�1. A total of 200 interferograms werecollected and co-added for each spectrum. Spectra were col-lected from inclusion-free, freshly broken resin chips placedon an infrared-transparent NaCl disc. To avoid the effectsof oversaturation of the spectra, sample thickness was keptbelow 5 lm. Depending on the quality of the sample, thespot size was set to values between 50 � 50 and100 � 100 lm (square aperture). Further details concerningthe FTIR analytical procedures are presented in Tappertet al. (2011).

The stable carbon isotope compositions of the resins weredetermined using a Finnigan MAT 252 dual inlet mass spec-trometer at the University of Alberta. Samples of untreatedresin (�1–5 mg) were combusted together with �1 g CuOas an oxygen source in a sealed and evacuated quartz tubeat 800 �C for �12 h. The d13C analyses were normalized toNBS-18 and NBS-19, and the results are given with respectto V-PDB (Coplen et al., 1983). Instrumental precision andaccuracy was on the order of ±0.02&. Repeated analysesof individual samples were within ±0.1&.

3. RESULTS

3.1. Infrared spectroscopy

The comparison of infrared spectra from modern andfossil resins (Fig. 3) shows that resins, irrespective of theirage and botanical origin, produce spectra that are very sim-ilar. This similarity is due to the dominance of diterpenoidsin the composition of the resins. However, some variabilityin the spectra exists and differences in the position andheight of specific absorption features can be used to distin-guish different resin types. These subtle spectroscopic differ-ences primarily reflect variations in the structure of the

dominant terpenoid structures, but they can also be influ-enced by the presence of additional organic componentsand the degree of polymerization.

Four compositional types of resins are distinguished inthis study. These resin types were produced by specific plantfamilies and are, accordingly, classified into the followingfour categories:

(1) Pinaceous resins are exclusively produced by conifersbelonging to the family Pinaceae. The infrared spec-tra of pinaceous resins are characterized by distinc-tive absorption peaks at 1460 and 1385 cm�1

(Fig. 3). The amplitude of the peak at 1460 cm�1 isgenerally lower than the peak at 1385 cm�1. Anothertrait is the absence or the weak development of apeak at 2848 cm�1. This peak is typically well devel-oped in cupressaceous-araucarian resins (see below).Further details about the spectral characteristics ofpinaceous resins can be found in Tappert et al.(2011). Based on results of gas chromatography–mass spectrometry (GC–MS), pinaceous resins consistprimarily of diterpenoids that are based on pimaraneand abietane structures (Langenheim, 2003).

(2) Cupressaceous-araucarian resins are mainly pro-duced by conifers of the families Cupressaceae andAraucariaceae, but also by Sciadopitys verticillata,the sole extant member of the family Sciadopytaceae(Wolfe et al., 2009). The infrared spectra of cupressa-ceous-araucarian resins are characterized by distinc-tive absorption peaks at 1448, 887, and 791 cm�1

(Fig. 3). In fossil resins, the 887 cm�1 peak is oftenreduced or absent as a result of the fossilization pro-cess. The amplitude of the absorption peak at1385 cm�1 is typically lower than that of the adjacentpeak at 1448 cm�1. Further details about the spectralcharacteristics of cupressaceous-araucarian resins andtheir distinction from pinaceous resins can be foundin Tappert et al. (2011). Cupressaceous-araucarian

Table 1Overview of the fossil resins analyzed in this study, including their age and botanical source.

No.Sample Location Stratigraphicsetting

Age/Epochof formation

AbsoluteAge[Ma]

CompositionalType (based onFTIRspectroscopy)

ClassaBotanicalsource

References to ageand/or botanicalsource

1 Colombiancopal

Colombia Undetermined Pliocene–Holocene

0.0002–2.5

Fabalean Ic Hymenaea sp. Ragazzi et al. (2003),Lambert et al. (1995)

2 Malaysianamber

Borneo Merit-Pila coalfield

MiddleMiocene

12 DipterocarpaceousII DipterocarpaceaeLangenheim and Beck(1965),Schlee and Chan (1992)

3 Mexicanamber

Chiapas La Quinta Fm.–Balumtun Fm.

Early Miocene 13–19 Fabalean Ic Hymenaea sp. Langenheim and Beck(1965),Cunningham et al., 1983,Solorzano Kraemer (2007)

4 Dominicanamber

DominicanRep.

La Toca Fm.–Yanigua Fm.

Late Oligocene–Early Miocene

15–20 Fabalean Ic Hymenaea sp. Langenheim and Beck(1965),Iturralde-Vinent andMacPhee (1996)

5 Balticamber

Balticcoast

Blaue Erde Fm.(redeposited)

Late Eocene–Early Oligocene

33–40 Cupressaceous-araucariansuc

Ia Sciadopitys sp. Langenheim (2003),Wolfe et al. (2009)

6 Bitterfeldamber

Germany Bitterfeld coalfield (redeposited)

Late Eocene–Early Oligocene

33–40 Cupressaceous-araucariansuc

Ia Sciadopitys sp. Fuhrmann (2005),Wolfe et al. (2009)

7 Giraffekimberliteamber

Canada(NT)

Peat sequence inkimberlite crater

Eocene 39–38 Cupressaceous-araucarian

Ib Metasequoia sp. Tappert et al. (2011)

8 TigerMountainamber

USA (WA) Tiger MountainFm.

Middle Eocene 47–48 Cupressaceous-araucarian

Ib Metasequoia sp. Mustoe (1985)

9 EllesmereIsland amber

Canada(NU)

Eureka SoundFm.

Middle Eocene 50 PinaceousSUC V Pseudolarix sp. Francis (1988), this study

10 Tadkeshwaramber

India Cambay shale Eocene 50–52 DipterocarpaceousII DipterocarpaceaeRust et al. (2010)

11 Pandakimberliteamber

Canada(NT)

Kimberlite craterinfill

Early Eocene 53 Cupressaceous-araucarian

Ib Metasequoia sp. Wolfe et al. (2012)

12 Alaskanamber

USA (AK) Chickaloon Fm. Paleocene–EarlyEocene

53–55 Cupressaceous-araucarian

Ib Metasequoia sp. Triplehorn et al. (1984),this study

13 Evansburgamber

Canada(AB)

Upper ScollardFm.

Paleocene 59–62 Cupressaceous-araucarian

Ib Metasequoia sp. This study

14 Geneseeamber

Canada(AB)

Scollard Fm. Early Paleocene 65 Cupressaceous-araucarian

Ib Metasequoia sp. This study

15 Albertanamber

Canada(AB)

HorseshoeCanyon Fm.

Campanian–Maastrichtian

70–73 Cupressaceous-araucarian

Ib Parataxodium sp.McKellar and Wolfe(2010)

16 Grassy Lakeamber

Canada(AB)

Foremost Fm. Campanian 76–78 Cupressaceous-araucarian

Ib Parataxodium sp.McKellar and Wolfe(2010)

17 Cedar Lakeamber

Canada(MT)

Foremost Fm.(redeposited)

Campanian 76–78 Cupressaceous-araucarian

Ib Parataxodium sp.McKellar and Wolfe(2010)

18 New Jerseyamber

USA (NJ) Raritan Fm. Turonian 90–94 Cupressaceous-araucarian

Ib Cupressaceae Grimaldi et al. (1989),Anderson (2006)

19 Burmeseamber

Myanmar Hukawng Basin Late Albian–EarlyCenomanian

96–102 Cupressaceous-araucarian

Ib Cupressaceae Grimaldi et al. (2002),Cruickshank and Ko(2003)

20 Spanishamber

Spain,various

Eschucha Fm. Albian 100–110 Cupressaceous-araucarian

Ib CheirolepidiaceaeAlonso et al. (2000)

21 Lebaneseamber

Lebanon Gres de Base Fm.Late Barremian–Early Aptian

123–127 Cupressaceous-araucarian

Ib Cupressaceae Nissenbaum and Horowitz(1992), Azar et al. (2010)

22 Wealdenamber

UK (Isleof Wight)

Hastings beds(Wealden Group)

Berriasian–Valanginian

138–142 Cupressaceous-araucarian

Ib Cupressaceae Nicholas et al. (1993)

23 Italianamber

Italy(Dolomites)

HeiligkreuzkofelFm.

Carnian 220–225 Cupressaceous-araucarian

Ib CheirolepidiaceaeRoghi et al. (2006),Ragazzi et al. (2003)

SUC: contains succinic acid esters (succinate).a Classification based on Anderson et al. (1992) and Anderson and Botto (1993).

244 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 245

resins consist primarily of diterpenoids that are basedon labdane skeletal structures, such as communicacid polymers (Thomas, 1970; Anderson et al., 1992).

(3) Fabalean resins are produced by the tropical angio-sperm Hymenaea (Fabaceae, Fabales). The infraredspectra of these resins superficially resemble spectrafrom pinaceous and cupressaceous-araucarian repre-sentatives (Fig. 3). Similar to pinaceous resins, faba-lean resins produce an absorption peak at 1460 cm�1,but the amplitude of the 1385 cm�1 peak is lower orequivalent to the 1460 cm�1 peak. A distinctive fea-ture of fabalean resin spectra is the lack of strongabsorption features between 970 and 1050 cm�1.Fabalean resins also produce an absorption peak at887 cm�1. Like cupressaceous-araucarian resins,fabalean resins are primarily based on labdane struc-tured diterpenoids, with zanzibaric and ozic acidpolymers being typical components (Cunninghamet al., 1983).

(4) Dipterocarpaceous resins are produced by membersof the angiosperm family Dipterocarpaceae, whichis a family of diverse and widespread tropical rainfor-est trees. The infrared spectra of dipterocarpaceousresins are quite distinct from other resin spectra(Fig. 3). The most characteristic spectroscopic fea-tures are the distinctive absorption triplet between1360 and 1400 cm�1, and an additional absorptionpeak at 1050 cm�1. Dipterocarpaceous resins areprimarily based on sesquiterpenoids, such as cadin-ene, but triterpenoids can also be present (Ander-son et al., 1992). The resin produced bydipterocarpaceous trees is colloquially referred to asDammar.

Some of the resins, including Baltic amber and resinsfrom Ellesmere Island, were found to produce an additionaland distinctive spectral feature at 1160 cm�1, which is asso-ciated with a broad shoulder or a broad peak between 1200and 1300 cm�1 (Fig. 3). These spectral features generally re-late to the presence of esterified succinic acid (succinate) asan additional component in the resin (Beck et al., 1965;Beck, 1986). Although succinate has been identified in somepinaceous resins—namely in resins from Pseudolarix amabi-

lis and in some fossil resins of Pseudolarix-affinity from theCanadian Arctic (Anderson and LePage, 1995; Poulin andHelwig, 2012)—it is the hallmark constituent of Baltic am-ber, which itself is a cupressaceous-araucarian resin. Thepresence of succinate has important implications in theassessment of the botanical source of fossil resins, and con-sequently, succinate-bearing resins are marked separately inTable 1.

Most of the fossil resins analyzed for this study are ofcupressaceous-araucarian type. Along with fabalean res-ins (e.g., Dominican and Mexican ambers), cupressa-ceous-araucarian resins make up the most productivecommercial amber deposits (Table 1). The predominanceof these resins in the fossil record is due to their terpe-noid composition because labdane-based resins polymer-ize rapidly and are highly resistant to chemicalbreakdown (Fig. 1B).

Although pinaceous resins are produced in extremelyhigh abundance in northern-hemisphere forests, their pres-ervation potential in the fossil record is limited. The mainconstituent of pinaceous resins—diterpenes with pimaraneand abietane skeletal structures—do not possess the func-tional groups required to form stable polymers. Therefore,they are prone to decomposition. Nevertheless, underfavorable conditions, even pinaceous resins can persist formillions of years (Wolfe et al., 2009).

In accordance with previous studies (Langenheim andBeck, 1965; Broughton, 1974; Tappert et al., 2011), all res-ins of a given type produce very similar FTIR absorptionspectra, irrespective of whether the samples are modernor fossil, which indicates that chemical transformationsduring burial and maturation are minor. This observationis important because it supports the premise that resinscan resist isotopic exchange over geological timescales.

3.2. Carbon stable isotopes

3.2.1. Modern resins

The d13C values of modern resins analyzed in this studyrange from �31.2& to �23.6&. These values are consistentwith previously published d13C analyses of modern resinsfrom comparable plant taxa (Nissenbaum and Yakir,1995; Murray et al., 1998; Nissenbaum et al., 2005). Thed13C values of the modern resins follow a near Gaussiandistribution, with a mean value of �26.9& (Fig. 4). It isnotable that resins from individual plant families have sim-ilar d13C distributions and similar mean d13C values(Fig. 5). This indicates that the isotopic fractionation thatoccurs during resin biosynthesis is nearly identical for theplant families considered in this study, despite differencesin the terpenoid profile of individual resin types.

Fig. 4 also shows the carbon isotopic compositions ofbulk leaf material from modern C3-plants using data com-piled by Korner et al. (1991), Diefendorf et al. (2010), andKohn (2010). The leaf material was sampled from a widerange of plant taxa growing under diverse climatic condi-tions, from alpine and polar to tropical. Therefore, the iso-topic composition should approximate the averagecomposition of modern C3 plant organic matter. However,samples from extremely dry environments and samples thatwere potentially influenced by the canopy effect (see below)were excluded. It is notable that the isotope values of theleaf material (�32.8& to �22.2&, mean: �27.1&) andthe distribution of isotope values are very similar to thosereported for modern resins. In conclusion, we find no con-sistent differences between the distributions of d13C valuesin modern leaves and modern resins.

3.2.2. Fossil resins

Carbon isotope values of fossil resins were found torange from �28.4& to �18.2&, which means that somefossil resins are more enriched in 13C compared to theirmodern counterparts. Data for individual samples areshown in Fig. 6. Similar to modern resins, the fossil resinsat each deposit produce a range of d13C values. Theseranges are similar to those observed for modern resins(i.e., �8&) provided that a comparable number of samples

Wollemia nobilis-modern-

Tiger Mountain aber (USA)-Middle Eocene-

Ellesmere Island amber (Canada)-Middle Eocene-

Genesee amber (Canada)-Early Paleocene-

Pseudolarix amabilis-modern-

Dominican amber-Late Oligocene-Early Miocene-

Hymenaea courbaril-modern-

Shorea rubriflora-modern-

Malaysian amber-Middle Miocene-

Pinaceous resins

Cupressaceous-araucarian resins

Fabaleanresins

Dipterocarpaceousresins

Wavenumbers (cm )-180010001200140016001800200022002400260028003000320034003600

887

1460

13857911050

succinateregion

11601448

2848fo(

ecnabrosbA)ytiralcroftsef

Mexican amber-Early Miocene-

Single bonds (x-H)

O-H C-H

Double bonds Single bonds/Skeletal vibrations

C=C, C=O C-O, C-C

Fig. 3. Micro-FTIR spectra of different resin types distinguished in this study (selected modern and fossil examples). Dashed lines markspectral features discussed in the text. Yellow backgrounds highlight the most distinctive spectral regions. Characteristic spectral regions formolecular vibrations relevant to resins are marked in blue. (For interpretation of the references to color in this figure legend, the reader isreferred to the web version of this article.)

246 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

were analyzed. For some of the amber deposits, only asmall number of samples were analyzed due to sampleavailability. Consequently, the range of isotope values forthese deposits tends to be narrower. Despite differences inthe number of samples for different deposits, a considerableshift in d13C between deposits is observed (Figs. 6 and 7).

To quantify the isotopic shift through time, and to mini-mize the effect of sample size, we used the mean d13C valuesof the fossil resins (d13Cresin

mean) as a representative measure foreach deposit (Table 2).

The d13Cresinmean of subrecent Colombian copal (�27.5&) is

nearly identical to modern resins, but with increasing age,

mean: -26.9= 1.39

n = 160

‰Fr

eque

ncy

Modern resins

mean = -27.1= 1.67

n = 891

Freq

uenc

y

Bulk C3 leaf

Growth conditions (interpreted)optimal typical poor

Fractionation efficiencywolhgih

A

B

13C (‰ vs PDB)

Fig. 4. Distribution of d13C values in (A) modern resins (data from this study with additional data of Hymenaea resins from Nissenbaumet al., 2005), and (B) bulk leaf matter from a wide range of C3 plant taxa (data from Korner et al., 1991; Diefendorf et al., 2010; Kohn, 2010).

R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 247

resins become continuously more enriched in 13C and con-sequently less negative in their d13Cresin

mean values. This trendtowards more enriched isotopic compositions in fossil res-ins continues through the Neogene and into the middle Eo-cene. Some of the most productive amber deposits,including the Dominican (d13Cresin

mean = �24.8&) and the Bal-tic (d13Cresin

mean = �23.5&) amber deposits, formed during thistime interval. Eocene ambers from Tiger Mountain (Wash-ington, USA) and Ellesmere Island (Nunavut, Canada) aremore than 5.3& enriched in 13C compared to modern res-ins, with a d13Cresin

mean of �21.7& and �21.6&, respectively.The latter are the isotopically most enriched Cenozoic res-ins in the present sample set.

The early Eocene is characterized by a dramatic shift to-wards more depleted resin isotope compositions, as demon-strated by amber from the Tadkeshwar coal field of India(d13Cresin

mean = �25.8&), and amber preserved in the Pandakimberlite pipe in Canada (d13Cresin

mean = �24.8&). The earli-est Eocene amber from the Chickaloon Formation in Alas-ka marks a return to slightly more 13C-enrichedcompositions, with a d13Cresin

meanof �23.6&.Paleocene amber from Evansburg (d13Cresin

mean ¼ �22:3&)and Genesee (d13Cresin

mean ¼ �23:1&) in Alberta, Canada,are more enriched than their early Eocene counterparts.However, from the Paleocene into the Late Cretaceous(Campanian), resin compositions become increasingly more

-31 -30 -29 -28 -27 -26 -25 -24 -230

5

10

15

20

25

30

35

Freq

uenc

y

Pinaceae(mean=-26.9‰, =1.12, n=86)

Cupressaceae-Araucariaceae(mean=-26.6‰, =1.77, n=50)

Fabaceae (angiosperm)(mean=-27.3‰, =1.28, n=24)

13C (‰ vs PDB)

Fig. 5. Distribution of d13C values from different types of modern resins, separated according to their source plant families. Data includepreviously published analyses of Hymenaea (Fabaceae) resins (n = 17) from Nissenbaum et al. (2005).

248 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

depleted in 13C. The most depleted Cretaceous resins with ad13Cresin

mean of �23.9& were found in the Horseshoe CanyonFormation (late Campanian) of southern and central Al-berta. From the late Campanian and into the Early Creta-ceous, this trend reverses and the d13Cresin

mean shifts to more13C-enriched compositions, as shown by the coeval GrassyLake and Cedar Lake amber (d13Cresin

mean ¼ �23:6&), and theNew Jersey Raritan amber (d13Cresin

mean ¼ �22:1&) (Fig. 7).Resins that formed between the Cenomanian and Barre-

mian are characterized by even more enriched d13Cresinmean

compositions. During this time interval, Burmese amber(d13Cresin

mean ¼ �21:3&), Spanish amber (d13Cresinmean ¼

�21:4&), and Lebanese amber (d13Cresinmean ¼ �21:1&) were

formed. The oldest Cretaceous amber samples were recov-ered from the Wealden Group of England (UK). Althoughthe d13Cresin

mean (�22.8&) of the Wealden amber is more de-pleted than that of the Burmese, Spanish, and Lebaneseamber, its d13Cresin

mean is not well constrained because onlytwo samples were available for analysis.

Pre-Cretaceous ambers were limited to samples from theTriassic (Carnian) Heiligkreuzkofel Formation of the Dol-omites, Italy. With a d13Cresin

mean of �20.9&, these ambers arethe most 13C-enriched in the current sample set. In accor-dance with previously published d13C values of amber fromthe same locality (d13Cresin

mean ¼ �20:12&; DalCorso et al.,2011), our results confirm that these Triassic resinsare > 6& more 13C-enriched relative to modern resins.

4. DISCUSSION

4.1. Influences on the d13C of modern plants

4.1.1. Local environmental factors

Experimental studies, including growth chamber experi-ments, have shown that plant d13C is influenced by a widerange of local environmental factors, such as water avail-ability, water composition (e.g., salinity, nutrients), lightexposure, local temperature, altitude, humidity, presence

of parasites, etc. (Park and Epstein, 1960; Guy et al.,1980; O’Leary, 1981; Korner et al., 1991; Arens and Jahren,2000; Edwards et al., 2000; Dawson et al., 2002; Grocke,2002; McKellar et al., 2011) (Fig. 8). Each of these factorshas a direct influence on the efficiency of fractionation inplants during photosynthesis, and each can cause an isoto-pic shift on the order of several permil in the d13C of aplant.

Every plant community will produce a range of d13C val-ues because the environmental conditions (i.e., growth con-ditions) for each plant are unique. However, each planttaxon has its own range of growth conditions under which13C fractionation is maximized. The observation that thed13C of modern resins follows a Gaussian distribution(Fig. 4), suggests that the majority of the resin-producingplants grew under conditions that were intermediate be-tween optimal (i.e., most efficient fractionation) and poor(i.e., least efficient fractionation). Therefore, the relatived13C values of plant matter within a local plant communitycan be viewed primarily as a measure of the growthconditions.

For most plants, including the copious resin producers,it can be assumed that the metabolized CO2 was sourcedfrom isotopically undisturbed air, which had a d13C compo-sition approximating a global atmospheric average. How-ever, local isotopic effects can alter the d13C of the CO2 inthe ambient air. Limited air circulation, for example, canlead to a reprocessing of respired air, which is depleted in13C (i.e., canopy effect). Metabolizing this depleted aircauses the d13C of plants to shift towards more negative val-ues. In extreme cases, it can result in a depletion in 13C ofplant matter of >6& (Medina and Minchin, 1980). Thecanopy effect primarily affect plants in the understory ofdense tropical rainforests. Among the plant families consid-ered in this study, however, only the Dipterocarpaceaegrow in such dense rainforest environments, and only theMiocene amber from Borneo and the Eocene amber fromTadkeshwar, India, are of dipterocarpacean origin.

-32 -30 -28 -26 -24 -22 -20 -18 -16

Pinaceae

Cupressaceae-Araucariaceae

Fabaceae

Columbian copal

Malaysian amber

Mexican amber

Dominican amber

Baltic amber

Giraffe kimberlite amber

Tiger Mountain amber

Ellesmere Island amber

Panda kimberlite amber

Alaskan amber

Evansburg amber

Genesee amber

Albertan amber

Grassy Lake amber

New Jersey amber

Burmese amber

Spanish amber

Lebanese amber

Wealden amber

Italian amber

13C (‰vs VPDB)

Cedar Lake amber

Bitterfeld amber

modern

Tadkeshwar 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., resintype. 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.

R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 249

4.1.2. Regional and global environmental factors

The results of recent studies on the global distribution ofd13C values in bulk leaf matter indicate that the fraction-ation of 13C in plants is not only influenced by local envi-ronmental factors, but also by regional precipitationpatterns (Diefendorf et al., 2010; Kohn, 2010; Fig. 8). Itwas found that the average d13C of bulk leaf matter andthe calculated fractionation values (D13C) show some corre-lation with the mean annual precipitation (MAP) at a givenlocation. The 13C fractionation thereby increases withincreasing MAP (Diefendorf et al., 2010; Kohn, 2010).However, the correlation is strongly influenced by plantsthat were sampled from very dry environments(MAP < 500 mm/year), in which 13C fractionation is re-duced, and those sampled from tropical rainforests

(MAP > 2000 mm/year), in which fractionation is in-creased. For plants that grew in environments with an inter-mediate MAP, average D13C values only show a weakcorrelation with MAP. Since most of the resin-producingplant families investigated here grow typically in environ-ments with intermediate MAP, the effect of MAP on theirD13C can be considered minor. This notion is supportedby the observation that the d13Cresin

mean values of resins fromdifferent modern plant families investigated in this studyare very similar. However, the Dipterocarpaceae representan exception, because they commonly grow in areas withhigh MAP. Therefore, they are predisposed towards in-creased 13C fractionation. This ecological preference mayexplain the low d13C values of dipterocarpaceous resins ob-served by Murray et al. (1998).

PleistocenePliocene

Mioocene

Paleocene

Eocene

Oligocene

Maastrichtian

Campanian

SantonianConiacian

Turonian

Cenomanian

Albian

Aptian

Barremian

Hauterivian

Valanginian

Berriasian

enegoeN

enegoe laPsuoeca ter

CetaL

suo ecaterC

y lra E

13C (‰vs PDB)-28 -26 -24 -22 -20

Modern Resins

10

20

30

40

50

60

70

80

90

100

110

120

130

140

Age (Ma)

Albertan amber

New Jersey amber

Grassy Lake/Cedar Lake ambers

Giraffe kimberlite amber

Baltic/Bitterfeld ambers

Dominican amber

Mexican amber

Spanish amber

Lebanese amber

Wealden amber

Columbian copal

Pinaceae

Fabaceae

Cupressaceae/Araucariaceae

Tiger Mountain amber

Alaskan amber

Ellesmere Island amber

Burmese amber

Panda kimberlite amber

Evansburg amber

Genesee amber

cissairT

Carnian Italian amber220

Ladinian230

Malysian amber

Tadkeshwar amber

Fig. 7. Temporal variations in the mean d13C of fossil resins (amber). Horizontal bars mark one standard deviation (see Fig. 6). Modern resinvalues 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 ofthis article.)

250 R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262

The d13C of atmospheric CO2 (d13Cresinmean) has a direct im-

pact on the d13C of plant organic matter, and analyses ofgas inclusions in ice cores (Francey et al., 1999) and of cor-

alline carbonates (Swart et al., 2010) indicate that the globalaverage d13Cresin

mean has evolved by at least 1& (from around�6.5& to <�7.5&) over the last 200 years. This rapid

Table 2Mean d13C values of fossil resins from individual deposits,including standard deviations, and number of samples analyzed(arranged by age).

No Sample d13Cmean (&) r (&) n

1 Columbian copal �27.5 0.6 62 Malaysian amber �26.1 0.3 123 Mexican amber �25.5 0.9 154 Dominican amber �24.8 1.6 325 Baltic amber �23.5 1.0 376 Bitterfeld amber �23.7 1.0 157 Giraffe kimberlite amber �22.5 1.9 238 Tiger Mountain amber �21.7 0.4 129 Ellesmere Island amber �21.6 1.5 12

10 Tadkeshwar amber �25.8 0.8 1011 Panda kimberlite amber �24.8 0.2 512 Alaskan amber �23.6 1.2 1513 Evansburg amber �22.3 0.8 1414 Genesee amber �23.1 2.0 1415 Albertan amber �23.9 1.8 3916 Grassy Lake amber �23.7 1.4 3417 Cedar Lake amber �23.2 1.4 1218 New Jersey amber �22.1 1.1 3619 Burmese amber �21.3 1.0 620 Spanish amber �21.4 1.5 2121 Lebanese amber �21.1 1.2 2122 Wealden amber �22.8 0.5 223 Italian amber �20.9 0.7 19

water availability

light exposurelocal temperature

altitude

water composition (salinity etc.)

nutrient availability

local 13Cair(e.g., canopy effect)

humidity

presence of parasites

Global13Cair

pCO2

pO2

Regionalmean annual precipitation

Local

~7-8‰variability

13C effectshiftshiftshift

shift

shift

etc.

combined effect:

increase

increase

increase

increase

increase

Fig. 8. Overview of global, regional, and local factors and theireffect on the d13C of plant organic matter.

R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 251

change in d13Cresinmean is generally linked to the increasing in-

flux of anthropogenic CO2 from fossil fuel combustion,which is depleted in 13C (Suess effect). Although the post-industrial decrease in d13Cresin

mean has been detected in thed13C of bulk plant material (e.g., tree rings; Loader et al.,2003), it was not possible to identify a similar shift ind13Cresin

mean between modern and pre-industrial resins (i.e.,Colombian copal), possibly due to the relatively small num-ber of copal samples analyzed.

4.1.3. Compositional factors

In addition to the isotopic variability that is caused byenvironmental factors, some plant compounds are isotopi-cally distinct from bulk plant matter. Cellulose, for exam-ple, is slightly more enriched in 13C (�2–3&), whereaslignin is slightly more depleted (�3–4&) compared to thebulk plant (Park and Epstein, 1960; Benner et al., 1987;Schleser et al., 1999; Brugnoli and Farquhar, 2000)(Fig. 9). The fractionation that causes the isotopic deviationof these plant compounds, consequently, must occur afterthe initial photosynthetic fixation of CO2 by ribulose-1,5-bisphosphate carboxylase oxygenase (Rubisco). Unlike cel-lulose and lignin, the d13C values of plant resins overlapwith those of primary plant metabolites, such as leaf sugarsand bulk leaf tissues (Fig. 9), which indicates that fraction-ation during resin biosynthesis (i.e., after the initial CO2 fix-ation) is minor, typically < 1& (Diefendorf et al., 2012).This conclusion is important because it suggests thatd13Cresin

mean � d13Cbulk plantmean , and shifts in d13Cresin

mean reflect shiftsin the d13Cbulk plant

mean of C3 plants through time.The observation that the d13C of individual plant com-

pounds can deviate from the d13C of the bulk plant, needsto be considered whenever non-specific plants remains areused in chemostratigraphic studies, since differential preser-vation of the individual compounds can result in consider-able shifts in d13C. For example, mean d13C values of bulkorganic matter (i.e., coal, coalified tissues, and cuticles)from the Cretaceous and earliest Cenozoic (Strauss and Pe-ters-Kottig, 2003) are �2.5& lower compared to d13Cresin

mean

values from the same time interval. This difference can bereadily explained by the preferential preservation of 13C-de-pleted lignin and the breakdown of 13C-enriched cellulosein fossil wood, from which the d13Cresin

mean is unaffected.

4.2. Causes for the shift in 13Cresinmean through time

The different types of modern resins investigated in thisstudy show a very similar distribution of d13C values(Fig. 5), which indicates that the structure of their terpenoidconstituents has little or no effect on their d13C. This con-clusion, however, contradicts claims that systematic differ-ences in the fractionation behavior between gymnospermsand angiosperms exist, with plant matter from angio-sperms, including resins, generally being enriched in 13Cby �2–3& (Diefendorf et al., 2012). Although it is conceiv-able that certain plant taxa are prone to isotopic biases, dueto their ecological preferences (e.g., dipterocarpaceous treesgrowing in areas with high MAP and closed canopy condi-tions), there is no conclusive evidence that the fractionationbehavior of C3 plants is affected by their phylogeneticposition.

The tendency of fossil resins to be isotopically more en-riched in 13C compared to modern resins could be inter-preted as diagenetic overprinting. After exudation, resinsgenerally lose a fraction of their terpenoid constituents—mainly volatile mono- and sesquiterpenes—in addition towater, whereas the non-volatile fractions polymerize rap-idly. Once initial polymerization has occurred, anyadditional maturation of the resins during fossilizationcauses 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 toinfluence the d13C of the resins significantly. In fact, if anisotopic exchange had taken place during diagenesis, fossilresins from different localities should show an erratic distri-bution of the d13Cresin

mean arising from different diagenetic con-ditions at each locality, but this is not observed.Furthermore, if resins displayed a tendency to becomeunstable and prone to carbon isotopic exchange throughtime, they should become continuously more enriched in13C with increasing age, which is also not the case. Diage-netic resetting should also diminish the natural variabilityin d13C of resins, and this would result in samples fromindividual localities converging on similar d13C values in-stead of retaining their original range of d13C values.

As previously mentioned, an important factor that influ-ences the d13C of plant organic matter, including resins, isthe carbon isotopic composition of the CO2 in the atmo-sphere (d13Catm

CO2). A shift in d13Catm

CO2, in fact, would result

in a comparable shift in the d13C of plant organic matter.Although changes in d13Catm

CO2have been invoked to explain

the d13C variations observed in fossil plant organic matter(Arens et al., 2000; Jahren, 2002; Strauss and Peters-Kottig,2003; Hasegawa et al., 2003; Jahren et al., 2008), subse-quent studies have questioned whether changes in d13Catm

CO2

alone can be responsible the observed d13C variability(e.g., Grocke, 2002; Beerling and Royer, 2002; Lomaxet al., 2012; Schubert and Jahren, 2012). If changes ind13Catm

CO2were to be the primary cause for the d13C variabil-

ity of fossil plant organic matter, shifts in d13Cresinmean should

correlate with independent proxies for d13Cresinmean, including

those that are based on the d13C of marine carbonates(e.g., calcareous tests of benthic foraminifera) (Zachoset al., 2001; Tipple et al., 2010). However, the comparisonof the marine d13C record with the resin record shows verylittle similarity (Fig. 11). A notable exception is the pro-nounced decline in d13C values, which started in the latePaleocene and lasted well into the early Eocene (�57–52 Ma, Fig. 11). This decline, which coincides with timesof peak Cenozoic temperatures (i.e., Early Eocene climaticoptimum, EECO), is discernible in both the marine and theresin d13C record and it is only interrupted by short-livednegative excursions in d13C and d18O (e.g., Paleocene Eo-cene Thermal Maximum, PETM) (Zachos et al., 2008).

The cause for the rapid change in atmosphere composi-tion during the PETM and other similar short-lived(<20 ka) events remains a matter of debate (McInerneyand Wing, 2011), but the observation that these eventsare associated with negative excursions in d13C and d18Oindicates that a considerable amount of 13C-depleted CO2

was added to the atmosphere during these intervals, causingpeak warming (hyperthermals). Proposed mechanisms tocreate such a rapid influx of 13C-depleted CO2 into theatmosphere include the dissociation and oxidation of meth-ane hydrates in marine sediments (Sloan et al., 1992; Dick-ens et al., 1995, 1997), the release of thermogenic CH4 fromsediments through magmatism and its subsequent oxida-tion (Svensen et al., 2004; Westerhold et al., 2009), an in-crease in oxidation of organic matter caused by thedrying of epicontinental seas (Higgins and Schrag, 2006),the release and oxidation of CH4 stored in polar permafrost

(DeConto et al., 2012), and the rapid burning of Paleoceneterrestrial carbon (Kurtz et al., 2003). Since the effects ofthese processes are considered to be short-lived, they areunlikely to be responsible for the more gradual decline inthe d13C between the late Paleocene and early Eocene.Therefore, other, more fundamental processes, such aschanges in the rate of oxidation of organic matter, mustbe invoked as driving forces for this d13C decline.

Despite the exceptional d13C excursions in the earlyPaleogene, the shifts in d13C of benthic foraminiferathrough the Cenozoic are only moderate (�2&) even attheir extremes. In fact, between the Middle Eocene andthe Miocene, d13C values of marine carbonates remain rel-atively constant, with only minor shifts of 61&. The shiftsin d13Cresin

mean over the same time interval are approximately6&, and they record a long-term trend towards more neg-ative values (Fig. 11). In summary, changes in d13Catm

CO2may

influence the d13C of plant organic matter to some extent,but their influence is insufficient to account for the magni-tude of the observed shifts in d13Cresin

mean.Further comparison of the d13Cresin

mean with the marine sta-ble isotope record shows that for much of the Cenozoic,negative shifts in d13Cresin

mean broadly correlate with positiveshifts of a similar magnitude in the d18O of marine carbon-ates (Fig. 11). The only pronounced exception is the latePaleocene to early Eocene isotopic decline, which resultedin a shift in both d13Cresin

mean and d18O to more negative values.The marine d18O record is primarily a temperature record;with less negative d18O values reflecting higher averageocean surface temperatures and consequently higher globaltemperatures (Emiliani, 1955; Bemis et al., 1998). The cor-relation between the marine d18O and the d13Cresin

mean record,therefore, indicates that resins that formed during intervalsof high global temperatures tend to be more enriched in13C. Since high global temperatures are generally linkedto high atmospheric pCO2 (Royer, 2006), variations inatmospheric pCO2 seem to be an obvious alternative to ex-plain shifts in the d13C of plant matter, including resins.

The results of recent growth experiments under controlledpCO2 and water/moisture availability suggest a hyperbolic in-crease of plant fractionation (D13C) with increasing pCO2

(Schubert and Jahren, 2012). These results are consistent withmost previous studies on 13C fractionation in plants, whichalso reported positive correlations between D13C and pCO2

(Park and Epstein, 1960; Beerling and Woodward, 1993; Tre-ydte et al., 2009). However, some studies found negative cor-relations or no correlations (Beerling, 1996; Jahren et al.,2008). The discrepancy between these studies is most likelythe result of differences in the control of the growth conditions(i.e., water availability, light exposure, etc.) because variationsin the growth conditions can mask the isotopic effect causedby increasing pCO2.

Irrespective of the extent of the pCO2 impact, the obser-vation that 13C fractionation in plants increases withincreasing pCO2 cannot account for the d13C shifts in theresin record (Fig. 7). In the resin record, the least depletedd13Cresin

mean compositions occur precisely during intervals withpredicted high atmospheric pCO2. We conclude, onceagain, that some other process must be responsible forthe shifts in d13Cresin

mean.

-25-20 -30 -35-15

PectinAsparatate

HemicelluloseAminoacids

Sap sugars

CelluloseStarch

Leaf sugars

Proteins-carotene

IsopreneLigninFatty acids

Total lipids

Bulk leaf

Resin(this study)

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 theanalyses of modern resins in this study; the range for bulk leaf matter is based on the compilations from Korner et al. (1991), Diefendorf et al.(2010), and Kohn (2010) (Fig. 4).

)aM (

egA

13C = 0pCO2

Stomatal densityGEOCARB IIIPedogenic carb.

Atmospheric O (%)p 2

Fig. 10. Comparison of calculated pO2 values using different models to quantify the impact of pCO2 on plant fractionation (dD13CpCO2).

Values for dD13CpCO2were calculated based on the relationship between pCO2 and D13C 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(GEOCARB III; Berner and Kothavala, 2001), and (c) the d13C of pedogenic carbonates (Ekart et al., 1999). pO2 values calculated fordD13CpCO2

¼ 0 (no effect of pCO2 on D13C) are shown for comparison.

R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 253

Another important factor of the atmosphere that influ-ences the isotopic fractionation in plants is the oxygen con-tent. Preliminary experimental studies on C3 plants indicatethat fractionation of carbon during photosynthesis will in-crease if the pO2 in the ambient air is significantly higherthan modern values (pO2 = 21%), resulting in plant bio-

mass that is isotopically more depleted in 13C (Berneret al., 2000). However, the opposite effect (i.e., a decreasein fractionation) has been observed under lower thanmodern pO2 in ambient air (Beerling et al., 2002). Thismeans that plants growing under low O2 conditions willbe more enriched in 13C 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 carbonisotope fractionation in C3 plants is consistent with theoret-ical considerations of leaf-gas exchange processes and thefixation of CO2 by plants. During photosynthesis, atmo-spheric CO2 is converted in a carboxylation reactionthrough the enzyme Rubisco to energy-rich metabolites,such as glucose. At the same time, atmospheric O2 com-petes with CO2 for reaction sites on Rubisco. The reactionof Rubisco with O2 (oxygenation) partially offsets photo-synthesis and results in the release of CO2 in the processof photorespiration. With higher atmospheric pO2, moreoxygen is used through photorespiration, which leads toan increase of CO2 within the plant cell and hence an in-crease in 13C fractionation. The relationship between 13C-fractionation in plants (D13C) and pCO2 has been describedin general terms by Farquhar et al. (1982) as:

D13C ¼ aþ ðb� aÞci=ca ð1Þ

where ci denotes the intercellular pCO2, and ca the pCO2 inthe ambient atmosphere. The factors a and b representfractionations that occur during diffusion through the sto-mata (a = 4.4&) and during fixation of CO2 by Rubisco(b = 27&), respectively.

The observed effect of atmospheric pO2 on the d13C ofplant tissues means that the d13C of plant organic mattershould be explored as a proxy for pO2 in ancient atmo-spheres (paleo-oxybarometer). Given that the d13C distribu-tions of modern plant organic matter and resins are verysimilar (Fig. 4), we conclude that amber has all the prere-quisite qualities to be used as a pO2-proxy.

To estimate pO2 from plant d13C it is necessary to quan-tify the 13C fractionation in the plant at varying atmo-spheric pO2. This information is typically derived fromexperimental work on plants in growth chambers undercontrolled pO2 conditions. However, only a few studies todate have attempted to systematically determine the influ-ence of pO2 on plant fractionation, and the results of thesestudies show large discrepancies (e.g., compare Smith et al.,1976; Berry et al., 1972 and Beerling et al., 2002). These dis-crepancies may also be the result of differences in the con-trol of the experimental growth conditions, similar tothose observed in the CO2 growth experiments, since evensmall variations in the growth conditions (water and nutri-ent availability, light exposure etc.) can mask the effects ofvarying pO2. Irrespective of the cause of the discrepancies,none of the proposed relationships that are based on exper-imental data (e.g., Berner et al., 2000; Beerling et al., 2002)can account for the magnitude of the shifts in d13Cresin

mean. Infact, calculated pO2 values for most of the Cretaceousand early Paleocene would be close to 0% if the relationshipbetween pO2 and D13C proposed by Berner et al. (2000) andBeerling et al. (2002) was applied to fossil resins. More real-istic pO2 values would require much larger empirical scalingfactors than those proposed by these authors.

Due to the inconsistencies in the experimental determi-nation of the 13C fractionation factor, we used a direct rela-tionship between resin-derived D13C and atmospheric pO2

as a first order approximation. Under this assumption,the fractionation effect from photorespiration is propor-tional to the pO2 in the ambient air. The use of a scaling

factor, in this case, is not necessary. A direct relationshipmay not be extended to extremely low values(pO2 < 10%), but we believe a direct relationship is a rea-sonable assumption for moderate pO2 levels as long asphysiological adaptations of plants (e.g., changes in metab-olism) do not come into play. In our model, paleo-pO2 atthe time of resin formation can be estimated as:

pO2 ¼D13Cresin � dD13CpCO2

D13C� pOmodern

2 ð2Þ

where D13Cresin denotes the 13C fractionation of resin-bear-ing plants at the time of resin formation, and D13C repre-sents the carbon isotope fractionation of plants in thecurrent atmosphere(pOmodern

2 ¼ 21%; pCOmodern2 ¼ 0:039%), which is approxi-

mately 21&. Based on Farquhar et al. (1982), D13Cresin

can be calculated as:

D13C13resin ¼

ðd13CatmCO2� d13Cmean

resin Þ � 1000

1000þ d13Cmeanresin

ð3Þ

In this equation, d13CatmCO2

denotes the isotopic composi-tion of CO2 in the ambient air. For modern plants,d13Catm

CO2can be directly measured. For samples from the

geological past, values for d13CatmCO2

were calculated fromthe d13C and d18O of benthic foraminiferal tests, followingthe approach of Tipple et al. (2010). The d18O compositionis thereby used to correct for the temperature-dependencyof the fractionation between dissolved inorganic carbonand CO2. For the Cenozoic, the d13C and d18O data to cal-culate d13Catm

CO2were taken from Zachos et al. (2001). For the

late and mid-Cretaceous, d13C and d18O values from ben-thic foraminifera were taken from Li and Keller (1999) (lateCampanian) and Huber et al. (1995) (Campanian–Albian).Due to the lack of high-quality stable isotope data frombenthic foraminifera for the Early Cretaceous and the Tri-assic, we used average d13C and d18O values from compila-tions of Weissert and Erba (2004) and Korte et al. (2005),which are based on bulk carbonate rocks and carbonaceousmacrofossils. Although this approach introduces someadditional uncertainties, most of the error on the d13Catm

CO2

calculations is due to uncertainties in the exact age place-ment of the fossil resin samples.

The factor dD13CpCO2in Eq. (2) denotes the change of

D13C at varying pCO2 relative to the fractionation at mod-ern pCO2. As previously mentioned, the influence of pCO2

on plant fractionation has recently been investigated ingrowth experiments under systematically controlled growthconditions by Schubert and Jahren (2012). These authorsobserved a hyperbolic relationship between the fraction-ation in C3 plants and pCO2, and formulated a best-fitfunction, using their own and previously published data.Based on this best-fit function proposed by Schubert andJahren (2012), we calculated dD13CpCO2

as:

dD13CpCO2¼ ð28:26Þð0:21ÞðpCO2 þ 25Þ

28:26þ ð0:21ÞðpCO2 þ 25Þ � D13C ð4Þ

whereby pCO2 is given in ppm. Although there is a generalconsensus that the pCO2 for most of the Cenozoic andMesozoic was higher than at present, considerable disputeremains about the absolute pCO2 values in the geological

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Fig. 11. Variations in predicted atmospheric pO2 (red) and d13Cresinmean (black) through time. The pO2 estimates represent mean values for

different pCO2 models, with the outline indicating the extreme values (Fig. 10). Error bars on d13Cresinmean 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 benthicforaminifera. Major climatic events, long-term temperature trends, and periods of large-scale igneous activity are shown for comparison. (Forinterpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 255

past. To address the uncertainties in the paleo-pCO2 esti-mates and their impact on the pO2 results, we calculateddD13CpCO2

using pCO2 values derived by three independentapproaches: (1) mass balance calculations (GEOCARB III,Berner and Kothavala, 2001), (2) the abundance of stomataon leaf surfaces (Retallack, 2001), and (3) the d13C of ped-ogenic carbonates (Ekart et al., 1999). Despite considerabledifferences in the proposed pCO2 levels between these threemodels, the calculated pO2 values show only minor devia-tions (<±1.5%) for most of the Cenozoic and Mesozoic.Larger discrepancies (<±3%) are restricted to the Paleoceneand early Eocene (Fig. 10). This indicates that most of theuncertainty in the calculated pO2 values rests on the accu-racy of the relationship between D13C and pCO2 as deter-mined by Schubert and Jahren (2012).

4.2.1. Mesozoic and Cenozoic pO2 evolution

A compilation of the calculated variations in pO2

through time is compared to d13Cresinmean, the marine stable iso-

tope record, and important environmental events in Fig. 11.It is apparent from this compilation that for most of theCenozoic and Mesozoic pO2 was considerably lower com-pared to today. This is consistent with a reduced 13C frac-

tionation (i.e., higher d13Cresinmean values) of resin-producing

plants in the geological past. The particularly highd13Cresin

mean values of early to mid-Cretaceous amber samplessuggest that the pO2 at the time of their formation was aslow as 10–11% (Fig. 11). Such low pO2 levels are regardedto be below the limit of combustion and seem to contradictthe presence of charcoal in Cretaceous sedimentary se-quences (Jones and Chaloner, 1991; Belcher and McElwain,2008). However, the presence of charcoal in sedimentaryrocks itself does not preclude low atmospheric pO2, sincethe charring process (i.e., pyrolysis) only occurs underlow-O2 conditions. Furthermore, experiments to establishthe lower limit of atmospheric pO2 required for combustiondid not consider certain factors that influence the combus-tion in nature, such as the effect of wind, which may lowerthe required pO2 considerably.

Despite some uncertainties about absolute pO2 values,our low-pO2 estimates for the mid-Cretaceous are consis-tent with pO2 estimates based on the abundance and massbalance of carbon and pyritic sulfur in sedimentary rocksduring that same time interval (Berner and Canfield,1989; Berner, 2006, Fig. 12). Our results, on the other hand,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,which predict a higher-than-modern pO2 for the Cretaceousand the Cenozoic (Glasspool and Scott, 2010; Brown et al.,2012, Fig. 12).

As previously mentioned, shifts in the Cenozoic d13Cresinmean

broadly resemble those in the marine d18O record, with theexception of the negative d13Cresin

mean excursion in the earlyPaleogene (Fig. 11). This indicates that changes in plantfractionation, i.e., changes that are linked to varying pO2,may also be linked to changes in global temperatures, andthus pCO2. In fact, our pO2 estimates for the Cretaceousand Cenozoic show moderate to strong negative correla-tions with pCO2 estimates derived by independent methods(Fig. 13). These correlations suggest that times of low pO2,on a broad scale, were characterized by high pCO2 and highglobal temperatures, and vice versa.

For the Early Cretaceous, it appears that an early phaseof low pO2 (12–13% O2) was followed by an interval of evenlower pO2 levels (10–11% O2) during the Aptian to Turo-nian. This phase of extremely-low pO2 coincides with themid-Cretaceous greenhouse—a time interval characterizedby high atmospheric pCO2, high global temperatures, andextensive ocean anoxic events (Jenkyns, 1980; Huberet al., 1995; Wilson and Norris, 2001; Leckie et al., 2002).From the Turonian to the end of the Cretaceous, atmo-spheric pO2 continuously rose to >15%, whereas pCO2

and global temperatures declined.The early Paleogene appears to have been characterized

by moderate pO2 levels in the atmosphere (14–18% O2),which declined to a Cenozoic minimum of �12% O2 imme-diately after the EECO (Fig. 11). The EECO marks the timeof highest global temperatures and pCO2 during the Ceno-

Fig. 12. Comparison of predicted atmospheric pO2 from this study wcalculations (GEOCARBSULF; Berner, 2006) and the abundance of cha

zoic, as indicated by the exceptionally light d18O composi-tion of marine carbonates deposited during that time(Fig. 11). As previously mentioned, the time interval justprior to the EECO (i.e., late Paleocene to early Eocene)was marked by a steady decline in the d13C of marine car-bonates and in d13Cresin

mean, indicating an increased influx of13C-depleted CO2 into the atmosphere and an associateddecrease in d13Catm

CO2that lasted for �5 million years. In view

of our pO2 estimates, this continuous influx of 13C-depletedCO2 was most likely the result of an increase in the rate ofoxidation (oxidative weathering), caused by the rising glo-bal temperatures and by elevated pO2 levels during thattime interval. The increased rate of oxidative weatheringwould have dramatically accelerated the oxidation of or-ganic matter and let to a further increase in pCO2 andgreenhouse forcing. At the same time, it would have re-sulted in a continuous depletion of O2 in the atmosphere,which would explain the low pO2 following the EECO.

After the height of the early Eocene greenhouse phase,atmospheric pO2 went into a long-term phase of continuousrecovery until it reached modern levels. The recovery in pO2

was most likely the result of a decreased rate of oxidativeweathering and an associated increase in carbon burial.This phase was also characterized by a continuous decreasein pCO2 and global temperatures, which culminated in theglacial inception in the southern and northern hemispheresduring the terminal Paleogene.

4.2.2. Volcanism and its effects on pO2

It is noteworthy that the intervals with lowest resin-inferred pO2 during the Cretaceous and the Cenozoicwere preceded by episodes of intense volcanism and the

ith previously proposed models that are based on mass balancercoal in the rock record (Glasspool and Scott, 2010).

Atmospheric O (%)p 2

cirehpsomtA

)vmpp(

OCp

2

Stomatal densityGEOCARB IIIPedogenic carbonates

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).

R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262 257

emplacement of some of the largest known igneous prov-inces (LIPs). These include the Ontong-Java Plateau andthe Parana-Etendeka Province (Early Cretaceous), as wellas the Deccan and North Atlantic Provinces (Late Creta-ceous to Paleocene) (Fig. 11, Larson and Erba, 1999).The emplacement of these LIPs was associated with theexpulsion of vast amounts of mantle-derived volcanic gases,which are typically dominated by H2O and CO2, but alsocontain sulfur dioxide (SO2) and other trace gases (Sy-monds et al., 1994). The expulsion of such large quantitiesof volcanic gases, particularly of CO2, provides a validmechanism to cause a rapid increase in atmospheric pCO2

and an associated increase of greenhouse forcing and globaltemperatures during these time intervals. Although there isstill some dispute over the long-term effect of volcanogenicCO2 emissions (Self et al., 2005), the hypothesis that LIPmagmatism contributed to changes in atmospheric compo-sition and global climate is supported by the marine d13Crecord, which indicates that the late Paleocene decline ofd13C was preceded by a notable rise in d13C values thatstarted in the mid-Paleocene. This rise in d13C coincideswith phases of intense volcanism in the Deccan area of In-dia, and in the North Atlantic Igneous Province (Fig. 11).Marine carbonates deposited at that time reached d13C val-ues of around +2& (Fig. 11; Zachos et al., 2001). Theseunusually positive values are consistent with a strong influxof mantle-derived CO2 into the atmosphere because mantleCO2 is enriched in 13C relative to CO2 derived from organicmatter and has a rather restricted d13C range (�8 to �3&,mean: ��5&) (Exley et al., 1986; Javoy and Pineau, 1991).

Although the episodes of LIP volcanism in the EarlyCretaceous and the Late Cretaceous/Paleocene may havenot been sufficient to sustain long-term greenhouse condi-tions, they may have served as a trigger for the subsequentincrease of oxidative weathering. This, in turn, would haveprovided a feedback for the further rise in pCO2 and the

prolonged greenhouse episodes that followed the LIP volca-nism. In this scenario, the elevated pO2 that prevailed be-fore the onset of the LIP volcanism (i.e., in the Early andLate Cretaceous) would have been an important factor thatdetermined the extent of the subsequent greenhouse phases.

After the cessation of LIP volcanism and the end of thegreenhouse periods in the Cretaceous and the Cenozoic,atmospheric pO2 was at its minimum, but slowly increasedover the following tens of millions of years, as the result of adecrease in oxidative weathering (associated with an in-crease in carbon burial) and the absence of the CO2-influxfrom LIP magmatism. After the mid-Cretaceous, the cycleof increasing pO2 and decreasing pCO2 lasted for �40 mil-lion years, but was interrupted by renewed LIP volcanismat the end of the Cretaceous. The Cenozoic decline ofpCO2 and increase in pO2, on the other hand, continueduninterrupted for �50 million years from the Middle Eo-cene into modern times. The prolonged absence of large-scale volcanism during that time period may have been aprerequisite for pCO2 and global temperatures to decreasesufficiently to allow sea ice to form on the polar oceans,ultimately triggering the onset of glaciation in both hemi-spheres in the latest Paleogene.

4.2.3. Effects of pO2 on animal evolution

Although it has been proposed that atmospheric pO2 hasa profound effect on the physiology and evolution of ani-mals, and that gigantism in animals is linked to high pO2

(Graham et al., 1995; Clapham and Karr, 2012), our datapreclude such a link for the evolution of giant theropodand sauropod dinosaurs. Considering the low pO2 that pre-vailed during the Cretaceous and, at least, part of the Tri-assic, it is more reasonable to assume that a high primaryphotosynthetic productivity (i.e., food availability) causedby high pCO2 and high global temperatures was the keyfactor 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 drawnfor the Eocene radiation of large placental mammals, whichhas previously been linked to high pO2 levels as well (Fal-kowski et al., 2005). Based on our data, however, this timeinterval was also characterized by low pO2.

5. CONCLUSIONS AND FUTURE DIRECTIONS

Due to their ability to polymerize into highly resistantorganic macromolecules that are demonstrably representa-tive of whole-plant carbon fractionation, resins can pre-serve pristine isotopic signatures over geological time.This makes them ideal natural biomarkers for chemostrati-graphic studies. As with all plant organic constituents, thed13C of C3 plant resins is influenced by a range of localenvironmental and ecological factors, which result in con-siderable isotopic variability. The variability seen in resinsis directly comparable to that of primary plant biomass,which suggests that the d13C of resins can be used as a mea-sure of the d13C of bulk plant organic matter. To minimizestatistical biases caused by variations in the growth condi-tions of individual plants, it is necessary to analyze suffi-ciently large sample populations and to compare theirmean isotopic values (d13Cresin

mean). This is particularly impor-tant for the comparison of resins of different ages. Sincemajor diagenetic influences on the d13C of fossil resins arenegligible, variations in d13Cresin

mean can be related to long-termglobal environmental changes. In fact, our present compila-tion and analysis of d13C values from exhaustively sampledresins, which is the largest data set of its kind, strongly sug-gests that the pO2 of ancient atmospheres is reliably repre-sented in the isotopic composition of these secondarymetabolites.

Resins can be used as proxies of atmospheric composi-tion for much of the Cenozoic and Mesozoic, and perhapseven into the Paleozoic, given that resin-producing plantshave existed since at least the Carboniferous. Resins fromdifferent localities, different botanical sources, and differentages are directly comparable because they consist of similarmacromolecules, and they are resistant to post-depositionalisotopic exchange. These advantages are complemented bythe fact that the analysis of resins does not require extensivesample treatment, which limits biases introduced by labora-tory procedures. Therefore, resins can provide a consistentand highly resolved terrestrial d13C record, particularlywhen samples from stratigraphically and chronologicallywell-characterized sedimentary units are analyzed.

To improve our understanding of the fractionationbehavior of plant organic matter in relation to atmosphericcomposition, it is also necessary to conduct additional, sys-tematically controlled growth chamber experiments on awider range of plants, including resin-producing plant taxa.This information will help to more accurately quantifycompositional changes of the atmosphere that occurredthroughout Earth’s history. Additional constraints on thecomposition of ancient atmospheres may also be gainedby combining isotopic information with careful analysisof biological inclusions within amber—e.g., insects, plants,and fungi—which are frequently preserved with exquisitedetail.

ACKNOWLEDGMENTS

We are grateful to Jim Basinger, Madelaine Bohme, XavierDelclos, Michael Engel, George Mustoe, Philip Perkins, GuidoRoghi, Alexander Schmidt, and Chris Williams for generously pro-viding samples used in this study. Thomas Stachel is thanked forproviding access to the FTIR spectrometer housed in his labora-tory. We also thank Stacey Gibb and Gabriela Gonzalez for assis-tance at various stages, and we would like to thank the threeanonymous reviewers, whose constructive comments have im-proved this publication considerably. Much of this research wassupported by the Natural Science and Engineering Research Coun-cil of Canada (Discovery awards to A.P. Wolfe and K. Muehlenb-achs, Postdoctoral fellowship to R.C. McKellar). J. Ortega wassupported by the Ministerio de Economıa y Competitividad, Ful-bright Espana, and FECYT.

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