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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 Geociències 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
reliable proxies. Here we develop a technique to estimate paleo-pO2 using the stable carbon isotope composition (d13C) 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 d13C signatures
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& (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 of
13
resins, we used mean d13C values (d13 Cresin
C
mean ) for each geological resin deposit. Fossil resins (n = 412) are generally enriched in
compared to their modern counterparts, with shifts in d13 Cresin
of
up
to
6&.
These
isotopic
shifts
follow
distinctive
trends
through
mean
time, which are unrelated to post-depositional processes including polymerization and diagenesis. The most enriched fossil resin
samples, with a d13 Cresin
mean between 22& and 21&, formed during the Triassic, the mid-Cretaceous, and the early Eocene. Experimental evidence and theoretical considerations suggest that neither change in pCO2 nor in the d13C of atmospheric CO2 can
13
account for the observed shifts in d13 Cresin
C in resin-producing plants (D13C), instead, is primarily influmean . The fractionation of
enced by atmospheric pO2, with more fractionation occurring at higher pO2. The enriched d13 Cresin
mean 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
18
correlation between the d13 Cresin
mean 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. Intervals 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).
0016-7037/$ - see front matter Ó 2013 Elsevier Ltd. All rights reserved.
http://dx.doi.org/10.1016/j.gca.2013.07.011
R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262
1. INTRODUCTION
Our knowledge of the composition of the atmosphere
through time and of its impact on climate change and
the evolution of life is based on evidence preserved in
the fossil record. For the Pleistocene and Holocene, a direct and continuous record of the partial pressures and
isotopic compositions of atmospheric gases is preserved
in the form of air vesicles entrapped in polar ice sheets
(e.g., Petit et al., 1999). With increasing age, the record
of atmospheric compositions becomes less directly accessible, and inferences about the composition of the atmosphere in the geological past rely on indirect measures
or proxies. The quality of such information depends directly on the ability of proxies to reliably capture the composition of ancient atmospheres.
Attempts have been made to reconstruct atmospheric
pCO2 and pO2 using mass balance calculations and flux
modeling (Berner, 1999; Berner and Kothavala, 2001; Berner et al., 2007). Other approaches to estimate the paleopCO2 have utilized the density of stomata on fossil leaf cuticles (Retallack, 2001, 2002; Beerling and Royer, 2002) and
the carbon stable isotopic composition of pedogenic carbonates (Cerling and Hay, 1986; Ekart et al., 1999; Bowen
and Beerling, 2004). One of the richest sources of information on ancient atmospheric compositions comes from the
marine sedimentary record, in particular from the stable
isotope composition (d13C, d18O, d11B) of biogenic carbonates (e.g., Clarke and Jenkyns, 1999; Pearson and Palmer,
2000; Royer et al., 2001; Zachos et al., 2001). This information has been used to infer changes in the isotopic composition of atmospheric CO2 through time and to derive
paleo-temperature estimates. More recently, the d13C of
marine phytoplankton biomarkers have been used to estimate the pCO2 of Cenozoic atmospheres (Pagani et al.,
2005). Additional constraints on the composition of ancient
atmospheres were obtained from the marine record using
the biochemical cycles of sulfur (Berner, 2006; Halevy
et al., 2012).
For terrestrial environments, attempts to estimate paleopCO2 have frequently exploited the d13C of fossil organic
matter, primarily plant organic matter (Strauss and Peters-Kottig, 2003; Jahren et al., 2008). A range of plant constituents have been targeted for d13C analysis, including
wood, foliage, and cellulose extracts, as well as the coals
to which they contribute (Stuiver and Braziunas, 1987;
Gröcke, 1998, 1999, 2002; Lücke et al., 1999; Peters-Kottig
et al., 2006; Poole and Van Bergen, 2006; Poole et al., 2006;
Bechtel et al., 2008). However, three key issues need to be
addressed when plant organic matter is used for stable carbon isotopic analyses in geochemical and chemostratigraphic studies:
(1) Plants are composed of a complex range of organic
compounds, including polysaccharides (e.g., cellulose), lignin, lipids, cuticular waxes, and photosynthates. Due to differential fractionation of carbon
isotopes during biosynthesis, the isotopic composition of individual plant fractions typically differ from
each other and from that of the bulk plant (Park and
241
Epstein, 1960; Brugnoli and Farquhar, 2000). In
addition, some plant groups use distinctive metabolic
pathways that fundamentally affect their isotopic
composition. For example, C4 and CAM plants are
generally more enriched in 13C compared to C3
plants (O’Leary, 1988; Farquhar et al., 1989). Therefore, if stable isotopes are used in chemostratigraphic
studies, it is crucial to determine the organic composition of the material under investigation and its
botanical source.
(2) Environmental and ecological factors can play an
important role in the isotopic fractionation of plants,
and individual isotope values of fossil or modern
plant material may reflect local instead of large-scale
environmental conditions.
(3) Plant constituents can undergo substantial compositional changes during and after burial and diagenesis.
These changes have the potential to alter the isotopic
composition of the fossil plant material, to the extent
of preventing a meaningful paleoenvironmental
interpretation.
To determine environmentally mediated changes in the
d13C of fossil plant organic matter through time, we analyzed modern and fossil resins from a wide range of localities and ages. Compared to other terrestrial plant
constituents, resins have chemical properties that make
them particularly suitable as proxies of environmental
changes over geological time. Resins can polymerize into
highly cross-linked hydrocarbons that have considerable
preservation potential in the geological record (Fig. 1). This
means that a reliable terrestrial stable isotope history can
potentially be reconstructed well into the Paleozoic, considering that the earliest known fossil resins are from the Carboniferous (White, 1914; Bray and Anderson, 2009).
Unlike many other fossil plant organic constituents,
resins can often be assigned to a specific source plant family, and only a few plant families produce significant
amounts of preservable resins, all of which have C3
metabolism. Furthermore, resins are chemically conservative, which means that the composition of the resins did
not change significantly as plants evolved. These factors
limit not only the chemical variability of fossil resins,
but also any potential family-specific biases, which could
influence the isotopic composition of fossil resins in
chemostratigraphic studies.
Although large samples of gem-quality fossil resins, or
amber, are found in fewer than 50 locations worldwide,
small pieces (<5 mm) of non-precious resin, which are
sometimes referred to as resinites, are found frequently,
and are often associated with coal deposits (Stach et al.,
1982). Previous studies of the carbon isotopic composition
of resins were limited because they were conducted on only
a few localities, and in most cases, using only a few samples
(Nissenbaum and Yakir, 1995; Murray et al., 1998; Nissenbaum et al., 2005; McKellar et al., 2008; DalCorso et al.,
2011). Our study is the first attempt to integrate the d13C
data of resins from a wide range of localities and ages,
and to determine their potential as proxies for changes in
atmosphere composition through geological time.
242
R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262
A
B
C
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
In total, 538 d13C measurements were conducted on
modern (n = 126) and fossil resins (n = 412) from various
locations worldwide, ranging in age from recent to Triassic
(Fig. 2, Table 1). Analytical work was performed on hardened resins that were visibly free of inclusions and vesicles.
Modern resins were collected primarily from the trunks of
native trees growing under a range of climatic conditions
(i.e., tropical to subarctic), although specimens from cultivars were also analyzed in some cases. We restricted the
sampling of modern resins to representatives of the families
Pinaceae (Pseudotsuga, Picea, Pinus), Cupressaceae (Metasequoia, Thuja), Araucariaceae (Agathis), and Fabaceae
(Hymenaea), as these families represent the most important
resin producers in modern environments and the fossil record (Langenheim, 2003).
The fossil resin samples include many small, non-precious varieties, in addition to genuine ambers. To simplify
the terminology, all fossil resins are referred to herein as
amber, with the exception of sub-recent resins (>200 years
old), which are referred to as copal. The sample set includes
specimens from many well-known deposits (e.g., Baltic amber, Bitterfeld amber, Dominican amber, Myanmar or Burmese amber, and Lebanese amber), but also material from
newly discovered occurrences (Table 1).
For many amber deposits, the age of amber formation
can be tightly constrained (±2 Ma or better) by independent stratigraphic controls or radiometric dating (Table 1).
However, for some of the deposits, the age constraints are
more provisional, especially for localities that show evidence of re-working or re-deposition (e.g., Baltic and Bitterfeld ambers). For these deposits, the age estimates have
relatively large uncertainties (±4 Ma).
To examine the chemical composition of the resins, representative samples of modern and fossil resins from each
locality were analyzed using micro-Fourier transform infrared (FTIR) spectroscopy. The resulting absorption spectra
were used to (1) classify the resins into distinct compositional groups (see below), and (2) determine the botanical
affinity of the fossil resins. Particular emphasis was placed
on identifying the botanical source of fossil resins from
R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262
Resin Type
pinaceous
cupressaceous-araucarian
fabaceous
dipterocarpaceous
243
Age
Neogene
Paleogene
Late Cretaceous
Early Cretaceous
Triassic
Fig. 2. Locations of amber deposits and types of amber analyzed in this study.
previously undescribed localities. Absorption spectra were
collected between 4000 and 650 cm1 (wavenumbers) using
a Thermo Nicolet Nexus 470 FTIR spectrometer equipped
with a Nicolet Continuum IR microscope, with a spectral
resolution of 4 cm1. A total of 200 interferograms were
collected and co-added for each spectrum. Spectra were collected from inclusion-free, freshly broken resin chips placed
on an infrared-transparent NaCl disc. To avoid the effects
of oversaturation of the spectra, sample thickness was kept
below 5 lm. Depending on the quality of the sample, the
spot size was set to values between 50 50 and
100 100 lm (square aperture). Further details concerning
the FTIR analytical procedures are presented in Tappert
et al. (2011).
The stable carbon isotope compositions of the resins were
determined using a Finnigan MAT 252 dual inlet mass spectrometer at the University of Alberta. Samples of untreated
resin (1–5 mg) were combusted together with 1 g CuO
as an oxygen source in a sealed and evacuated quartz tube
at 800 °C for 12 h. The d13C analyses were normalized to
NBS-18 and NBS-19, and the results are given with respect
to V-PDB (Coplen et al., 1983). Instrumental precision and
accuracy was on the order of ±0.02&. Repeated analyses
of individual samples were within ±0.1&.
3. RESULTS
3.1. Infrared spectroscopy
The comparison of infrared spectra from modern and
fossil resins (Fig. 3) shows that resins, irrespective of their
age and botanical origin, produce spectra that are very similar. This similarity is due to the dominance of diterpenoids
in the composition of the resins. However, some variability
in the spectra exists and differences in the position and
height of specific absorption features can be used to distinguish different resin types. These subtle spectroscopic differences primarily reflect variations in the structure of the
dominant terpenoid structures, but they can also be influenced by the presence of additional organic components
and the degree of polymerization.
Four compositional types of resins are distinguished in
this study. These resin types were produced by specific plant
families and are, accordingly, classified into the following
four categories:
(1) Pinaceous resins are exclusively produced by conifers
belonging to the family Pinaceae. The infrared spectra of pinaceous resins are characterized by distinctive absorption peaks at 1460 and 1385 cm1
(Fig. 3). The amplitude of the peak at 1460 cm1 is
generally lower than the peak at 1385 cm1. Another
trait is the absence or the weak development of a
peak at 2848 cm1. This peak is typically well developed in cupressaceous-araucarian resins (see below).
Further details about the spectral characteristics of
pinaceous resins can be found in Tappert et al.
(2011). Based on results of gas chromatography–
mass spectrometry (GC–MS), pinaceous resins consist
primarily of diterpenoids that are based on pimarane
and abietane structures (Langenheim, 2003).
(2) Cupressaceous-araucarian resins are mainly produced by conifers of the families Cupressaceae and
Araucariaceae, but also by Sciadopitys verticillata,
the sole extant member of the family Sciadopytaceae
(Wolfe et al., 2009). The infrared spectra of cupressaceous-araucarian resins are characterized by distinctive absorption peaks at 1448, 887, and 791 cm1
(Fig. 3). In fossil resins, the 887 cm1 peak is often
reduced or absent as a result of the fossilization process. The amplitude of the absorption peak at
1385 cm1 is typically lower than that of the adjacent
peak at 1448 cm1. Further details about the spectral
characteristics of cupressaceous-araucarian resins and
their distinction from pinaceous resins can be found
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
ClassaBotanical
source
Stratigraphic
setting
Age/Epoch
of formation
AbsoluteCompositional
Age
Type (based on
[Ma]
FTIR
spectroscopy)
1 Colombian Colombia
copal
2 Malaysian Borneo
amber
Undetermined
Pliocene–
Holocene
Middle
Miocene
0.0002– Fabalean
Ic
2.5
12
DipterocarpaceousII
3 Mexican
amber
La Quinta Fm.– Early Miocene 13–19
Balumtun Fm.
Fabalean
Ic
La Toca Fm.–
Yanigua Fm.
Fabalean
Ic
Blaue Erde Fm. Late Eocene– 33–40
(redeposited)
Early Oligocene
Bitterfeld coal
Late Eocene– 33–40
field (redeposited)Early Oligocene
Peat sequence in Eocene
39–38
kimberlite crater
Cupressaceousaraucariansuc
Cupressaceousaraucariansuc
Cupressaceousaraucarian
Ia
Ib
Ragazzi et al. (2003),
Lambert et al. (1995)
DipterocarpaceaeLangenheim and Beck
(1965),
Schlee and Chan (1992)
Hymenaea sp.
Langenheim and Beck
(1965),
Cunningham et al., 1983,
Solórzano Kraemer (2007)
Hymenaea sp.
Langenheim and Beck
(1965),
Iturralde-Vinent and
MacPhee (1996)
Sciadopitys sp. Langenheim (2003),
Wolfe et al. (2009)
Sciadopitys sp. Fuhrmann (2005),
Wolfe et al. (2009)
Metasequoia sp. Tappert et al. (2011)
Tiger Mountain Middle Eocene 47–48
Fm.
Cupressaceousaraucarian
Ib
Metasequoia sp. Mustoe (1985)
Eureka Sound
Fm.
Cambay shale
Middle Eocene 50
PinaceousSUC
V
Pseudolarix sp.
Eocene
50–52
DipterocarpaceousII
DipterocarpaceaeRust et al. (2010)
53
Cupressaceousaraucarian
Ib
Metasequoia sp. Wolfe et al. (2012)
Chickaloon Fm. Paleocene–Early 53–55
Eocene
Upper Scollard Paleocene
59–62
Fm.
Scollard Fm.
Early Paleocene 65
Ib
Ib
Metasequoia sp. Triplehorn et al. (1984),
this study
Metasequoia sp. This study
Ib
Metasequoia sp. This study
Horseshoe
Canyon Fm.
Foremost Fm.
Campanian–
Maastrichtian
Campanian
Ib
Foremost Fm.
(redeposited)
Raritan Fm.
Campanian
Cupressaceousaraucarian
Cupressaceousaraucarian
Cupressaceousaraucarian
70–73 Cupressaceousaraucarian
76–78 Cupressaceousaraucarian
76–78 Cupressaceousaraucarian
90–94 Cupressaceousaraucarian
96–102 Cupressaceousaraucarian
Ib
Parataxodium sp.McKellar and Wolfe
(2010)
Parataxodium sp.McKellar and Wolfe
(2010)
Parataxodium sp.McKellar and Wolfe
(2010)
Cupressaceae
Grimaldi et al. (1989),
Anderson (2006)
Cupressaceae
Grimaldi et al. (2002),
Cruickshank and Ko
(2003)
CheirolepidiaceaeAlonso et al. (2000)
Ib
Cupressaceae
Ib
Cupressaceae
Ib
CheirolepidiaceaeRoghi et al. (2006),
Ragazzi et al. (2003)
Chiapas
4 Dominican Dominican
amber
Rep.
5 Baltic
Baltic
amber
coast
6 Bitterfeld
Germany
amber
7 Giraffe
Canada
kimberlite (NT)
amber
8 Tiger
USA (WA)
Mountain
amber
9 Ellesmere Canada
Island amber(NU)
10 Tadkeshwar India
amber
11 Panda
Canada
kimberlite (NT)
amber
12 Alaskan
USA (AK)
amber
13 Evansburg Canada
amber
(AB)
14 Genesee
Canada
amber
(AB)
15 Albertan
Canada
amber
(AB)
16 Grassy Lake Canada
amber
(AB)
17 Cedar Lake Canada
amber
(MT)
18 New Jersey USA (NJ)
amber
19 Burmese
Myanmar
amber
20 Spanish
amber
21 Lebanese
amber
22 Wealden
amber
23 Italian
amber
SUC
a
Spain,
various
Lebanon
UK (Isle
of Wight)
Italy
(Dolomites)
Merit-Pila coal
field
Late Oligocene– 15–20
Early Miocene
Kimberlite crater Early Eocene
infill
Turonian
Hukawng Basin Late Albian–
Early
Cenomanian
Eschucha Fm. Albian
100–110 Cupressaceousaraucarian
Grès de Base Fm.Late Barremian–123–127 CupressaceousEarly Aptian
araucarian
Hastings beds
Berriasian–
138–142 Cupressaceous(Wealden Group)Valanginian
araucarian
Heiligkreuzkofel Carnian
220–225 CupressaceousFm.
araucarian
: contains succinic acid esters (succinate).
Classification based on Anderson et al. (1992) and Anderson and Botto (1993).
Ia
Ib
Ib
Ib
Ib
References to age
and/or botanical
source
Hymenaea sp.
Francis (1988), this study
Nissenbaum and Horowitz
(1992), Azar et al. (2010)
Nicholas et al. (1993)
R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262
resins consist primarily of diterpenoids that are based
on labdane skeletal structures, such as communic
acid polymers (Thomas, 1970; Anderson et al., 1992).
(3) Fabalean resins are produced by the tropical angiosperm Hymenaea (Fabaceae, Fabales). The infrared
spectra of these resins superficially resemble spectra
from pinaceous and cupressaceous-araucarian representatives (Fig. 3). Similar to pinaceous resins, fabalean resins produce an absorption peak at 1460 cm1,
but the amplitude of the 1385 cm1 peak is lower or
equivalent to the 1460 cm1 peak. A distinctive feature of fabalean resin spectra is the lack of strong
absorption features between 970 and 1050 cm1.
Fabalean resins also produce an absorption peak at
887 cm1. Like cupressaceous-araucarian resins,
fabalean resins are primarily based on labdane structured diterpenoids, with zanzibaric and ozic acid
polymers being typical components (Cunningham
et al., 1983).
(4) Dipterocarpaceous resins are produced by members
of the angiosperm family Dipterocarpaceae, which
is a family of diverse and widespread tropical rainforest trees. The infrared spectra of dipterocarpaceous
resins are quite distinct from other resin spectra
(Fig. 3). The most characteristic spectroscopic features are the distinctive absorption triplet between
1360 and 1400 cm1, and an additional absorption
peak at 1050 cm1. Dipterocarpaceous resins are
primarily based on sesquiterpenoids, such as cadinene, but triterpenoids can also be present (Anderson et al., 1992). The resin produced by
dipterocarpaceous trees is colloquially referred to as
Dammar.
Some of the resins, including Baltic amber and resins
from Ellesmere Island, were found to produce an additional
and distinctive spectral feature at 1160 cm1, which is associated with a broad shoulder or a broad peak between 1200
and 1300 cm1 (Fig. 3). These spectral features generally relate to the presence of esterified succinic acid (succinate) as
an additional component in the resin (Beck et al., 1965;
Beck, 1986). Although succinate has been identified in some
pinaceous resins—namely in resins from Pseudolarix amabilis and in some fossil resins of Pseudolarix-affinity from the
Canadian Arctic (Anderson and LePage, 1995; Poulin and
Helwig, 2012)—it is the hallmark constituent of Baltic amber, which itself is a cupressaceous-araucarian resin. The
presence of succinate has important implications in the
assessment of the botanical source of fossil resins, and consequently, succinate-bearing resins are marked separately in
Table 1.
Most of the fossil resins analyzed for this study are of
cupressaceous-araucarian type. Along with fabalean resins (e.g., Dominican and Mexican ambers), cupressaceous-araucarian resins make up the most productive
commercial amber deposits (Table 1). The predominance
of these resins in the fossil record is due to their terpenoid composition because labdane-based resins polymerize rapidly and are highly resistant to chemical
breakdown (Fig. 1B).
245
Although pinaceous resins are produced in extremely
high abundance in northern-hemisphere forests, their preservation potential in the fossil record is limited. The main
constituent of pinaceous resins—diterpenes with pimarane
and abietane skeletal structures—do not possess the functional groups required to form stable polymers. Therefore,
they are prone to decomposition. Nevertheless, under
favorable conditions, even pinaceous resins can persist for
millions of years (Wolfe et al., 2009).
In accordance with previous studies (Langenheim and
Beck, 1965; Broughton, 1974; Tappert et al., 2011), all resins of a given type produce very similar FTIR absorption
spectra, irrespective of whether the samples are modern
or fossil, which indicates that chemical transformations
during burial and maturation are minor. This observation
is important because it supports the premise that resins
can 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 study
range from 31.2& to 23.6&. These values are consistent
with previously published d13C analyses of modern resins
from comparable plant taxa (Nissenbaum and Yakir,
1995; Murray et al., 1998; Nissenbaum et al., 2005). The
d13C values of the modern resins follow a near Gaussian
distribution, with a mean value of 26.9& (Fig. 4). It is
notable that resins from individual plant families have similar d13C distributions and similar mean d13C values
(Fig. 5). This indicates that the isotopic fractionation that
occurs during resin biosynthesis is nearly identical for the
plant families considered in this study, despite differences
in the terpenoid profile of individual resin types.
Fig. 4 also shows the carbon isotopic compositions of
bulk leaf material from modern C3-plants using data compiled by Körner et al. (1991), Diefendorf et al. (2010), and
Kohn (2010). The leaf material was sampled from a wide
range of plant taxa growing under diverse climatic conditions, from alpine and polar to tropical. Therefore, the isotopic composition should approximate the average
composition of modern C3 plant organic matter. However,
samples from extremely dry environments and samples that
were potentially influenced by the canopy effect (see below)
were excluded. It is notable that the isotope values of the
leaf material (32.8& to 22.2&, mean: 27.1&) and
the distribution of isotope values are very similar to those
reported for modern resins. In conclusion, we find no consistent differences between the distributions of d13C values
in modern leaves and modern resins.
3.2.2. Fossil resins
Carbon isotope values of fossil resins were found to
range from 28.4& to 18.2&, which means that some
fossil resins are more enriched in 13C compared to their
modern counterparts. Data for individual samples are
shown in Fig. 6. Similar to modern resins, the fossil resins
at each deposit produce a range of d13C values. These
ranges are similar to those observed for modern resins
(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
C=C, C=O
C-O, C-C
C-H
O-H
2848
1460 1448
1160
791
1050
887
1385
succinate
region
Pinaceous resins
Pseudolarix amabilis
-modern-
Ellesmere Island amber (Canada)
-Middle Eocene-
Absorbance (offest for clarity)
Cupressaceousaraucarian resins
Wollemia nobilis
-modern-
Tiger Mountain aber (USA)
-Middle Eocene-
Genesee amber (Canada)
-Early Paleocene-
Fabalean
resins
Hymenaea courbaril
-modern-
Mexican amber
-Early Miocene-
Dominican amber
-Late Oligocene-Early Miocene-
Dipterocarpaceous
resins
Shorea rubriflora
-modernMalaysian amber
-Middle Miocene-
3600
3400
3200
3000
2800
2600
2400
2200
2000
1800
1600
1400
1200
1000
800
-1
Wavenumbers (cm )
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
small number of samples were analyzed due to sample
availability. Consequently, the range of isotope values for
these deposits tends to be narrower. Despite differences in
the number of samples for different deposits, a considerable
shift in d13C between deposits is observed (Figs. 6 and 7).
To quantify the isotopic shift through time, and to minimize the effect of sample size, we used the mean d13C values
of the fossil resins (d13 Cresin
mean ) as a representative measure for
each deposit (Table 2).
The d13 Cresin
mean of subrecent Colombian copal (27.5&) is
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
poor
Modern resins
mean: -26.9‰
= 1.39
n = 160
Bulk C3 leaf
mean = -27.1‰
= 1.67
n = 891
Frequency
A
typical
Frequency
B
13
C (‰ 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 Körner et al., 1991; Diefendorf et al., 2010; Kohn, 2010).
resins become continuously more enriched in 13C and consequently less negative in their d13 Cresin
mean values. This trend
towards more enriched isotopic compositions in fossil resins continues through the Neogene and into the middle Eocene. Some of the most productive amber deposits,
including the Dominican (d13 Cresin
mean = 24.8&) and the Baltic (d13 Cresin
=
23.5&)
amber
deposits, formed during this
mean
time interval. Eocene ambers from Tiger Mountain (Washington, USA) and Ellesmere Island (Nunavut, Canada) are
more than 5.3& enriched in 13C compared to modern resins, with a d13 Cresin
mean of 21.7& and 21.6&, respectively.
The latter are the isotopically most enriched Cenozoic resins in the present sample set.
The early Eocene is characterized by a dramatic shift towards more depleted resin isotope compositions, as demonstrated by amber from the Tadkeshwar coal field of India
(d13 Cresin
mean = 25.8&), and amber preserved in the Panda
kimberlite pipe in Canada (d13 Cresin
mean = 24.8&). The earliest Eocene amber from the Chickaloon Formation in Alaska marks a return to slightly more 13C-enriched
compositions, with a d13 Cresin
mean of 23.6&.
Paleocene amber from Evansburg (d13 Cresin
mean ¼ 22:3&)
and Genesee (d13 Cresin
¼
23:1&)
in
Alberta,
Canada,
mean
are more enriched than their early Eocene counterparts.
However, from the Paleocene into the Late Cretaceous
(Campanian), resin compositions become increasingly more
248
R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262
35
30
Frequency
25
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)
20
15
10
5
0
-31
-30
-29
-28
13
-27
-26
-25
-24
-23
C (‰ 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
d13 Cresin
mean of 23.9& were found in the Horseshoe Canyon
Formation (late Campanian) of southern and central Alberta. From the late Campanian and into the Early Cretaceous, this trend reverses and the d13 Cresin
mean shifts to more
13
C-enriched compositions, as shown by the coeval Grassy
Lake and Cedar Lake amber (d13 Cresin
mean ¼ 23:6&), and the
New Jersey Raritan amber (d13 Cresin
mean ¼ 22:1&) (Fig. 7).
Resins that formed between the Cenomanian and Barremian are characterized by even more enriched d13 Cresin
mean
compositions. During this time interval, Burmese amber
(d13 Cresin
Spanish
amber
(d13 Cresin
mean ¼ 21:3&),
mean ¼
13 resin
21:4&), and Lebanese amber (d Cmean ¼ 21:1&) were
formed. The oldest Cretaceous amber samples were recovered from the Wealden Group of England (UK). Although
the d13 Cresin
mean (22.8&) of the Wealden amber is more depleted than that of the Burmese, Spanish, and Lebanese
amber, its d13 Cresin
mean is not well constrained because only
two samples were available for analysis.
Pre-Cretaceous ambers were limited to samples from the
Triassic (Carnian) Heiligkreuzkofel Formation of the Dolomites, Italy. With a d13 Cresin
mean of 20.9&, these ambers are
the most 13C-enriched in the current sample set. In accordance with previously published d13C values of amber from
the same locality (d13 Cresin
mean ¼ 20:12&; DalCorso et al.,
2011), our results confirm that these Triassic resins
are > 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 experiments, have shown that plant d13C is influenced by a wide
range of local environmental factors, such as water availability, water composition (e.g., salinity, nutrients), light
exposure, local temperature, altitude, humidity, presence
of parasites, etc. (Park and Epstein, 1960; Guy et al.,
1980; O’Leary, 1981; Körner et al., 1991; Arens and Jahren,
2000; Edwards et al., 2000; Dawson et al., 2002; Gröcke,
2002; McKellar et al., 2011) (Fig. 8). Each of these factors
has a direct influence on the efficiency of fractionation in
plants during photosynthesis, and each can cause an isotopic shift on the order of several permil in the d13C of a
plant.
Every plant community will produce a range of d13C values because the environmental conditions (i.e., growth conditions) for each plant are unique. However, each plant
taxon has its own range of growth conditions under which
13
C fractionation is maximized. The observation that the
d13C of modern resins follows a Gaussian distribution
(Fig. 4), suggests that the majority of the resin-producing
plants grew under conditions that were intermediate between optimal (i.e., most efficient fractionation) and poor
(i.e., least efficient fractionation). Therefore, the relative
d13C values of plant matter within a local plant community
can be viewed primarily as a measure of the growth
conditions.
For most plants, including the copious resin producers,
it can be assumed that the metabolized CO2 was sourced
from isotopically undisturbed air, which had a d13C composition approximating a global atmospheric average. However, local isotopic effects can alter the d13C of the CO2 in
the ambient air. Limited air circulation, for example, can
lead to a reprocessing of respired air, which is depleted in
13
C (i.e., canopy effect). Metabolizing this depleted air
causes the d13C of plants to shift towards more negative values. In extreme cases, it can result in a depletion in 13C of
plant matter of >6& (Medina and Minchin, 1980). The
canopy effect primarily affect plants in the understory of
dense tropical rainforests. Among the plant families considered in this study, however, only the Dipterocarpaceae
grow in such dense rainforest environments, and only the
Miocene amber from Borneo and the Eocene amber from
Tadkeshwar, India, are of dipterocarpacean origin.
R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262
13
-32
-30
-28
-26
249
C (‰vs VPDB)
-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
The results of recent studies on the global distribution of
d13C values in bulk leaf matter indicate that the fractionation of 13C in plants is not only influenced by local environmental factors, but also by regional precipitation
patterns (Diefendorf et al., 2010; Kohn, 2010; Fig. 8). It
was found that the average d13C of bulk leaf matter and
the calculated fractionation values (D13C) show some correlation with the mean annual precipitation (MAP) at a given
location. The 13C fractionation thereby increases with
increasing MAP (Diefendorf et al., 2010; Kohn, 2010).
However, the correlation is strongly influenced by plants
that were sampled from very dry environments
(MAP < 500 mm/year), in which 13C fractionation is reduced, and those sampled from tropical rainforests
(MAP > 2000 mm/year), in which fractionation is increased. For plants that grew in environments with an intermediate MAP, average D13C values only show a weak
correlation with MAP. Since most of the resin-producing
plant families investigated here grow typically in environments with intermediate MAP, the effect of MAP on their
D13C can be considered minor. This notion is supported
by the observation that the d13 Cresin
mean values of resins from
different modern plant families investigated in this study
are very similar. However, the Dipterocarpaceae represent
an exception, because they commonly grow in areas with
high MAP. Therefore, they are predisposed towards increased 13C fractionation. This ecological preference may
explain the low d13C values of dipterocarpaceous resins observed 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)
Neogene
10
Columbian copal
Pleistocene
Pliocene
Malysian amber
Mioocene
Mexican amber
20
Dominican amber
Oligocene
30
40
50
Paleogene
Baltic/Bitterfeld ambers
60
Giraffe kimberlite amber
Eocene
Tiger Mountain amber
Ellesmere Island amber
Tadkeshwar amber
Panda kimberlite amber
Alaskan amber
Evansburg amber
Paleocene
Genesee amber
Maastrichtian
70
80
90
Late Cretaceous
Albertan amber
Grassy Lake/Cedar Lake ambers
Campanian
Santonian
Coniacian
Turonian
New Jersey amber
Cenomanian
Burmese amber
100
Spanish amber
120
130
Early Cretaceous
Albian
110
Aptian
Lebanese amber
Barremian
Hauterivian
Valanginian
140
Wealden amber
220
230
Triassic
Berriasian
Italian amber
Carnian
Ladinian
-28
-26
-24
-22
-20
13
C (‰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.)
The d13C of atmospheric CO2 (d13 Cresin
mean ) has a direct impact on the d13C of plant organic matter, and analyses of
gas inclusions in ice cores (Francey et al., 1999) and of cor-
alline carbonates (Swart et al., 2010) indicate that the global
average d13 Cresin
mean has evolved by at least 1& (from around
6.5& to <7.5&) over the last 200 years. This rapid
R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262
Table 2
Mean 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
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23
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
Grassy Lake amber
Cedar Lake amber
New Jersey amber
Burmese amber
Spanish amber
Lebanese amber
Wealden amber
Italian amber
27.5
26.1
25.5
24.8
23.5
23.7
22.5
21.7
21.6
25.8
24.8
23.6
22.3
23.1
23.9
23.7
23.2
22.1
21.3
21.4
21.1
22.8
20.9
0.6
0.3
0.9
1.6
1.0
1.0
1.9
0.4
1.5
0.8
0.2
1.2
0.8
2.0
1.8
1.4
1.4
1.1
1.0
1.5
1.2
0.5
0.7
6
12
15
32
37
15
23
12
12
10
5
15
14
14
39
34
12
36
6
21
21
2
19
Global
13
C effect
Cair
increase
shift
pCO2
increase
shift
pO2
increase
shift
13
mean annual precipitation
increase
shift
Local
local
13
Cair
(e.g., canopy effect)
water availability
water composition (salinity etc.)
humidity
light exposure
local temperature
nutrient availability
altitude
presence of parasites
etc.
4.1.3. Compositional factors
In addition to the isotopic variability that is caused by
environmental factors, some plant compounds are isotopically distinct from bulk plant matter. Cellulose, for example, is slightly more enriched in 13C (2–3&), whereas
lignin is slightly more depleted (3–4&) compared to the
bulk 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 deviation
of these plant compounds, consequently, must occur after
the initial photosynthetic fixation of CO2 by ribulose-1,5bisphosphate carboxylase oxygenase (Rubisco). Unlike cellulose and lignin, the d13C values of plant resins overlap
with those of primary plant metabolites, such as leaf sugars
and bulk leaf tissues (Fig. 9), which indicates that fractionation during resin biosynthesis (i.e., after the initial CO2 fixation) is minor, typically < 1& (Diefendorf et al., 2012).
This conclusion is important because it suggests that
13 bulk plant
, and shifts in d13 Cresin
d13 Cresin
mean d Cmean
mean reflect shifts
13 bulk plant
in the d Cmean
of C3 plants through time.
The observation that the d13C of individual plant compounds can deviate from the d13C of the bulk plant, needs
to be considered whenever non-specific plants remains are
used in chemostratigraphic studies, since differential preservation of the individual compounds can result in considerable shifts in d13C. For example, mean d13C values of bulk
organic matter (i.e., coal, coalified tissues, and cuticles)
from the Cretaceous and earliest Cenozoic (Strauss and Peters-Kottig, 2003) are 2.5& lower compared to d13 Cresin
mean
values from the same time interval. This difference can be
readily explained by the preferential preservation of 13C-depleted lignin and the breakdown of 13C-enriched cellulose
in fossil wood, from which the d13 Cresin
mean is unaffected.
4.2. Causes for the shift in
Regional
increase
shift
combined effect:
~7-8‰variability
Fig. 8. Overview of global, regional, and local factors and their
effect on the d13C of plant organic matter.
change in d13 Cresin
mean is generally linked to the increasing influx of anthropogenic CO2 from fossil fuel combustion,
which is depleted in 13C (Suess effect). Although the postindustrial decrease in d13 Cresin
mean has been detected in the
d13C of bulk plant material (e.g., tree rings; Loader et al.,
2003), it was not possible to identify a similar shift in
d13 Cresin
mean between modern and pre-industrial resins (i.e.,
Colombian copal), possibly due to the relatively small number of copal samples analyzed.
251
13
Cresin
mean through time
The different types of modern resins investigated in this
study show a very similar distribution of d13C values
(Fig. 5), which indicates that the structure of their terpenoid
constituents has little or no effect on their d13C. This conclusion, however, contradicts claims that systematic differences in the fractionation behavior between gymnosperms
and angiosperms exist, with plant matter from angiosperms, including resins, generally being enriched in 13C
by 2–3& (Diefendorf et al., 2012). Although it is conceivable that certain plant taxa are prone to isotopic biases, due
to their ecological preferences (e.g., dipterocarpaceous trees
growing in areas with high MAP and closed canopy conditions), there is no conclusive evidence that the fractionation
behavior of C3 plants is affected by their phylogenetic
position.
The tendency of fossil resins to be isotopically more enriched in 13C compared to modern resins could be interpreted as diagenetic overprinting. After exudation, resins
generally lose a fraction of their terpenoid constituents—
mainly volatile mono- and sesquiterpenes—in addition to
water, whereas the non-volatile fractions polymerize rapidly. Once initial polymerization has occurred, any
additional maturation of the resins during fossilization
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
influence the d13C of the resins significantly. In fact, if an
isotopic exchange had taken place during diagenesis, fossil
resins from different localities should show an erratic distribution of the d13 Cresin
mean arising from different diagenetic conditions at each locality, but this is not observed.
Furthermore, if resins displayed a tendency to become
unstable and prone to carbon isotopic exchange through
time, they should become continuously more enriched in
13
C with increasing age, which is also not the case. Diagenetic resetting should also diminish the natural variability
in d13C of resins, and this would result in samples from
individual localities converging on similar d13C values instead of retaining their original range of d13C values.
As previously mentioned, an important factor that influences the d13C of plant organic matter, including resins, is
the carbon isotopic composition of the CO2 in the atmo13 atm
sphere (d13 Catm
CO2 ). A shift in d CCO2 , in fact, would result
in a comparable shift in the d13C of plant organic matter.
Although changes in d13 Catm
CO2 have 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), subsequent studies have questioned whether changes in d13 Catm
CO2
alone can be responsible the observed d13C variability
(e.g., Gröcke, 2002; Beerling and Royer, 2002; Lomax
et al., 2012; Schubert and Jahren, 2012). If changes in
13
d13 Catm
CO2 were to be the primary cause for the d C variability of fossil plant organic matter, shifts in d13 Cresin
mean should
correlate with independent proxies for d13 Cresin
mean , including
those that are based on the d13C of marine carbonates
(e.g., calcareous tests of benthic foraminifera) (Zachos
et al., 2001; Tipple et al., 2010). However, the comparison
of the marine d13C record with the resin record shows very
little similarity (Fig. 11). A notable exception is the pronounced decline in d13C values, which started in the late
Paleocene and lasted well into the early Eocene (57–
52 Ma, Fig. 11). This decline, which coincides with times
of peak Cenozoic temperatures (i.e., Early Eocene climatic
optimum, EECO), is discernible in both the marine and the
resin d13C record and it is only interrupted by short-lived
negative excursions in d13C and d18O (e.g., Paleocene Eocene Thermal Maximum, PETM) (Zachos et al., 2008).
The cause for the rapid change in atmosphere composition during the PETM and other similar short-lived
(<20 ka) events remains a matter of debate (McInerney
and Wing, 2011), but the observation that these events
are associated with negative excursions in d13C and d18O
indicates that a considerable amount of 13C-depleted CO2
was added to the atmosphere during these intervals, causing
peak warming (hyperthermals). Proposed mechanisms to
create such a rapid influx of 13C-depleted CO2 into the
atmosphere include the dissociation and oxidation of methane hydrates in marine sediments (Sloan et al., 1992; Dickens et al., 1995, 1997), the release of thermogenic CH4 from
sediments through magmatism and its subsequent oxidation (Svensen et al., 2004; Westerhold et al., 2009), an increase in oxidation of organic matter caused by the
drying 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 Paleocene
terrestrial carbon (Kurtz et al., 2003). Since the effects of
these processes are considered to be short-lived, they are
unlikely to be responsible for the more gradual decline in
the d13C between the late Paleocene and early Eocene.
Therefore, other, more fundamental processes, such as
changes in the rate of oxidation of organic matter, must
be invoked as driving forces for this d13C decline.
Despite the exceptional d13C excursions in the early
Paleogene, the shifts in d13C of benthic foraminifera
through the Cenozoic are only moderate (2&) even at
their extremes. In fact, between the Middle Eocene and
the Miocene, d13C values of marine carbonates remain relatively constant, with only minor shifts of 61&. The shifts
in d13 Cresin
mean over the same time interval are approximately
6&, and they record a long-term trend towards more negative values (Fig. 11). In summary, changes in d13 Catm
CO2 may
influence the d13C of plant organic matter to some extent,
but their influence is insufficient to account for the magnitude of the observed shifts in d13 Cresin
mean .
Further comparison of the d13 Cresin
mean with the marine stable isotope record shows that for much of the Cenozoic,
negative shifts in d13 Cresin
mean broadly correlate with positive
shifts of a similar magnitude in the d18O of marine carbonates (Fig. 11). The only pronounced exception is the late
Paleocene to early Eocene isotopic decline, which resulted
18
in a shift in both d13 Cresin
mean and d O to more negative values.
18
The marine d O record is primarily a temperature record;
with less negative d18O values reflecting higher average
ocean surface temperatures and consequently higher global
temperatures (Emiliani, 1955; Bemis et al., 1998). The correlation between the marine d18O and the d13 Cresin
mean record,
therefore, indicates that resins that formed during intervals
of high global temperatures tend to be more enriched in
13
C. Since high global temperatures are generally linked
to high atmospheric pCO2 (Royer, 2006), variations in
atmospheric pCO2 seem to be an obvious alternative to explain shifts in the d13C of plant matter, including resins.
The results of recent growth experiments under controlled
pCO2 and water/moisture availability suggest a hyperbolic increase of plant fractionation (D13C) with increasing pCO2
(Schubert and Jahren, 2012). These results are consistent with
most previous studies on 13C fractionation in plants, which
also reported positive correlations between D13C and pCO2
(Park and Epstein, 1960; Beerling and Woodward, 1993; Treydte et al., 2009). However, some studies found negative correlations or no correlations (Beerling, 1996; Jahren et al.,
2008). The discrepancy between these studies is most likely
the result of differences in the control of the growth conditions
(i.e., water availability, light exposure, etc.) because variations
in the growth conditions can mask the isotopic effect caused
by increasing pCO2.
Irrespective of the extent of the pCO2 impact, the observation that 13C fractionation in plants increases with
increasing pCO2 cannot account for the d13C shifts in the
resin record (Fig. 7). In the resin record, the least depleted
d13 Cresin
mean compositions occur precisely during intervals with
predicted high atmospheric pCO2. We conclude, once
again, that some other process must be responsible for
the shifts in d13 Cresin
mean .
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
13
C (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 Körner et al. (1991), Diefendorf et al.
(2010), and Kohn (2010) (Fig. 4).
Age (Ma)
13
CpCO2= 0
Stomatal density
GEOCARB III
Pedogenic carb.
Atmospheric pO2 (%)
Fig. 10. Comparison of calculated pO2 values using different models to quantify the impact of pCO2 on plant fractionation (dD13 CpCO2 ).
Values for dD13 CpCO2 were 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 for
dD13 CpCO2 ¼ 0 (no effect of pCO2 on D13C) are shown for comparison.
Another important factor of the atmosphere that influences the isotopic fractionation in plants is the oxygen content. Preliminary experimental studies on C3 plants indicate
that fractionation of carbon during photosynthesis will increase if the pO2 in the ambient air is significantly higher
than modern values (pO2 = 21%), resulting in plant bio-
mass that is isotopically more depleted in 13C (Berner
et al., 2000). However, the opposite effect (i.e., a decrease
in fractionation) has been observed under lower than
modern pO2 in ambient air (Beerling et al., 2002). This
means that plants growing under low O2 conditions will
be 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 carbon
isotope fractionation in C3 plants is consistent with theoretical considerations of leaf-gas exchange processes and the
fixation of CO2 by plants. During photosynthesis, atmospheric CO2 is converted in a carboxylation reaction
through the enzyme Rubisco to energy-rich metabolites,
such as glucose. At the same time, atmospheric O2 competes with CO2 for reaction sites on Rubisco. The reaction
of Rubisco with O2 (oxygenation) partially offsets photosynthesis and results in the release of CO2 in the process
of photorespiration. With higher atmospheric pO2, more
oxygen is used through photorespiration, which leads to
an increase of CO2 within the plant cell and hence an increase in 13C fractionation. The relationship between 13Cfractionation in plants (D13C) and pCO2 has been described
in general terms by Farquhar et al. (1982) as:
D13 C ¼ a þ ðb aÞci =ca
ð1Þ
where ci denotes the intercellular pCO2, and ca the pCO2 in
the ambient atmosphere. The factors a and b represent
fractionations that occur during diffusion through the stomata (a = 4.4&) and during fixation of CO2 by Rubisco
(b = 27&), respectively.
The observed effect of atmospheric pO2 on the d13C of
plant tissues means that the d13C of plant organic matter
should be explored as a proxy for pO2 in ancient atmospheres (paleo-oxybarometer). Given that the d13C distributions of modern plant organic matter and resins are very
similar (Fig. 4), we conclude that amber has all the prerequisite qualities to be used as a pO2-proxy.
To estimate pO2 from plant d13C it is necessary to quantify the 13C fractionation in the plant at varying atmospheric pO2. This information is typically derived from
experimental work on plants in growth chambers under
controlled pO2 conditions. However, only a few studies to
date have attempted to systematically determine the influence of pO2 on plant fractionation, and the results of these
studies show large discrepancies (e.g., compare Smith et al.,
1976; Berry et al., 1972 and Beerling et al., 2002). These discrepancies may also be the result of differences in the control of the experimental growth conditions, similar to
those observed in the CO2 growth experiments, since even
small variations in the growth conditions (water and nutrient availability, light exposure etc.) can mask the effects of
varying pO2. Irrespective of the cause of the discrepancies,
none of the proposed relationships that are based on experimental data (e.g., Berner et al., 2000; Beerling et al., 2002)
can account for the magnitude of the shifts in d13 Cresin
mean . In
fact, calculated pO2 values for most of the Cretaceous
and early Paleocene would be close to 0% if the relationship
between pO2 and D13C proposed by Berner et al. (2000) and
Beerling et al. (2002) was applied to fossil resins. More realistic pO2 values would require much larger empirical scaling
factors than those proposed by these authors.
Due to the inconsistencies in the experimental determination of the 13C fractionation factor, we used a direct relationship between resin-derived D13C and atmospheric pO2
as a first order approximation. Under this assumption,
the fractionation effect from photorespiration is proportional to the pO2 in the ambient air. The use of a scaling
factor, in this case, is not necessary. A direct relationship
may not be extended to extremely low values
(pO2 < 10%), but we believe a direct relationship is a reasonable assumption for moderate pO2 levels as long as
physiological adaptations of plants (e.g., changes in metabolism) do not come into play. In our model, paleo-pO2 at
the time of resin formation can be estimated as:
pO2 ¼
D13 Cresin dD13 CpCO2
pOmodern
2
D13 C
ð2Þ
where D13Cresin denotes the 13C fractionation of resin-bearing plants at the time of resin formation, and D13C represents the carbon isotope fractionation of plants in the
current
atmosphere
(pOmodern
¼ 21%; pCOmodern
¼ 0:039%), which is approxi2
2
mately 21&. Based on Farquhar et al. (1982), D13Cresin
can be calculated as:
D13 C13
resin ¼
13 mean
ðd13 Catm
CO2 d Cresin Þ 1000
1000 þ d13 Cmean
resin
ð3Þ
In this equation, d13 Catm
CO2 denotes the isotopic composition of CO2 in the ambient air. For modern plants,
d13 Catm
CO2 can be directly measured. For samples from the
geological past, values for d13 Catm
CO2 were calculated from
the d13C and d18O of benthic foraminiferal tests, following
the approach of Tipple et al. (2010). The d18O composition
is thereby used to correct for the temperature-dependency
of the fractionation between dissolved inorganic carbon
and CO2. For the Cenozoic, the d13C and d18O data to calculate d13 Catm
CO2 were taken from Zachos et al. (2001). For the
late and mid-Cretaceous, d13C and d18O values from benthic foraminifera were taken from Li and Keller (1999) (late
Campanian) and Huber et al. (1995) (Campanian–Albian).
Due to the lack of high-quality stable isotope data from
benthic foraminifera for the Early Cretaceous and the Triassic, we used average d13C and d18O values from compilations of Weissert and Erba (2004) and Korte et al. (2005),
which are based on bulk carbonate rocks and carbonaceous
macrofossils. Although this approach introduces some
additional uncertainties, most of the error on the d13 Catm
CO2
calculations is due to uncertainties in the exact age placement of the fossil resin samples.
The factor dD13 CpCO2 in Eq. (2) denotes the change of
13
D C at varying pCO2 relative to the fractionation at modern pCO2. As previously mentioned, the influence of pCO2
on plant fractionation has recently been investigated in
growth experiments under systematically controlled growth
conditions by Schubert and Jahren (2012). These authors
observed a hyperbolic relationship between the fractionation in C3 plants and pCO2, and formulated a best-fit
function, using their own and previously published data.
Based on this best-fit function proposed by Schubert and
Jahren (2012), we calculated dD13 CpCO2 as:
dD13 CpCO2 ¼
ð28:26Þð0:21ÞðpCO2 þ 25Þ
D13 C
28:26 þ ð0:21ÞðpCO2 þ 25Þ
ð4Þ
whereby pCO2 is given in ppm. Although there is a general
consensus that the pCO2 for most of the Cenozoic and
Mesozoic was higher than at present, considerable dispute
remains about the absolute pCO2 values in the geological
R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262
Igneous
activity
(LIPs)
Percentage of O2 in atmosphere
10
15
20
10
Neogene
Pleistocene
Pliocene
5
-28
C (‰)
-26
-24
-22
-26
-24
-22
-20
North Atlantic
Deccan
Paleogene
Eocene
Antarctic
ice sheets
forming
EECO
PETM
Paleocene
Pg
K
warming
12
Late Cretaceous
Age (Ma)
13
4
Mioocene
Maastrichtian
70
90
1
2 0
Oligocene
60
80
O (‰)
3
2
cooling
30
50
-1
C (‰)
0 1
Fossil resins (mean value)
18
13
Arctic ice sheets
forming
20
40
Marine stable isotope record
Climatic events
and long-term
T-trends
255
8
4
0
Ice-free temperature
Campanian
cooling
Santonian
Coniacian
Turonian
Cenomanian
Ontong Java
Parana-Etendeka
100
120
130
Early Cretaceous
Albian
11 0
Aptian
Barremian
Mid-Cretaceous
greenhouse
warming
Hauterivian
Valanginian
140
220
230
Triassic
Berriasian
Carnian
Ladinian
10
15
20
-28
13
Cresin
mean
Fig. 11. Variations in predicted atmospheric pO2 (red) and d
(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 d13 Cresin
mean 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.)
past. To address the uncertainties in the paleo-pCO2 estimates and their impact on the pO2 results, we calculated
dD13 CpCO2 using pCO2 values derived by three independent
approaches: (1) mass balance calculations (GEOCARB III,
Berner and Kothavala, 2001), (2) the abundance of stomata
on leaf surfaces (Retallack, 2001), and (3) the d13C of pedogenic carbonates (Ekart et al., 1999). Despite considerable
differences in the proposed pCO2 levels between these three
models, the calculated pO2 values show only minor deviations (<±1.5%) for most of the Cenozoic and Mesozoic.
Larger discrepancies (<±3%) are restricted to the Paleocene
and early Eocene (Fig. 10). This indicates that most of the
uncertainty in the calculated pO2 values rests on the accuracy of the relationship between D13C and pCO2 as determined 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 d13 Cresin
mean , the marine stable isotope record, and important environmental events in Fig. 11.
It is apparent from this compilation that for most of the
Cenozoic and Mesozoic pO2 was considerably lower compared to today. This is consistent with a reduced 13C frac-
tionation (i.e., higher d13 Cresin
mean values) of resin-producing
plants in the geological past. The particularly high
d13 Cresin
mean values of early to mid-Cretaceous amber samples
suggest that the pO2 at the time of their formation was as
low as 10–11% (Fig. 11). Such low pO2 levels are regarded
to be below the limit of combustion and seem to contradict
the presence of charcoal in Cretaceous sedimentary sequences (Jones and Chaloner, 1991; Belcher and McElwain,
2008). However, the presence of charcoal in sedimentary
rocks itself does not preclude low atmospheric pO2, since
the charring process (i.e., pyrolysis) only occurs under
low-O2 conditions. Furthermore, experiments to establish
the lower limit of atmospheric pO2 required for combustion
did not consider certain factors that influence the combustion in nature, such as the effect of wind, which may lower
the required pO2 considerably.
Despite some uncertainties about absolute pO2 values,
our low-pO2 estimates for the mid-Cretaceous are consistent with pO2 estimates based on the abundance and mass
balance of carbon and pyritic sulfur in sedimentary rocks
during 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
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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 Cretaceous
and the Cenozoic (Glasspool and Scott, 2010; Brown et al.,
2012, Fig. 12).
As previously mentioned, shifts in the Cenozoic d13 Cresin
mean
broadly resemble those in the marine d18O record, with the
exception of the negative d13 Cresin
mean excursion in the early
Paleogene (Fig. 11). This indicates that changes in plant
fractionation, i.e., changes that are linked to varying pO2,
may also be linked to changes in global temperatures, and
thus pCO2. In fact, our pO2 estimates for the Cretaceous
and Cenozoic show moderate to strong negative correlations 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 high
global temperatures, and vice versa.
For the Early Cretaceous, it appears that an early phase
of low pO2 (12–13% O2) was followed by an interval of even
lower pO2 levels (10–11% O2) during the Aptian to Turonian. This phase of extremely-low pO2 coincides with the
mid-Cretaceous greenhouse—a time interval characterized
by high atmospheric pCO2, high global temperatures, and
extensive ocean anoxic events (Jenkyns, 1980; Huber
et al., 1995; Wilson and Norris, 2001; Leckie et al., 2002).
From the Turonian to the end of the Cretaceous, atmospheric 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 immediately after the EECO (Fig. 11). The EECO marks the time
of highest global temperatures and pCO2 during the Ceno-
zoic, as indicated by the exceptionally light d18O composition of marine carbonates deposited during that time
(Fig. 11). As previously mentioned, the time interval just
prior to the EECO (i.e., late Paleocene to early Eocene)
was marked by a steady decline in the d13C of marine carbonates and in d13 Cresin
mean , indicating an increased influx of
13
C-depleted CO2 into the atmosphere and an associated
decrease in d13 Catm
CO2 that lasted for 5 million years. In view
of our pO2 estimates, this continuous influx of 13C-depleted
CO2 was most likely the result of an increase in the rate of
oxidation (oxidative weathering), caused by the rising global temperatures and by elevated pO2 levels during that
time interval. The increased rate of oxidative weathering
would have dramatically accelerated the oxidation of organic matter and let to a further increase in pCO2 and
greenhouse forcing. At the same time, it would have resulted 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 continuous
recovery until it reached modern levels. The recovery in pO2
was most likely the result of a decreased rate of oxidative
weathering and an associated increase in carbon burial.
This phase was also characterized by a continuous decrease
in pCO2 and global temperatures, which culminated in the
glacial inception in the southern and northern hemispheres
during the terminal Paleogene.
4.2.2. Volcanism and its effects on pO2
It is noteworthy that the intervals with lowest resininferred pO2 during the Cretaceous and the Cenozoic
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
Atmospheric pCO2 (ppmv)
Stomatal density
GEOCARB III
Pedogenic carbonates
Atmospheric pO2 (%)
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 provinces (LIPs). These include the Ontong-Java Plateau and
the Parana-Etendeka Province (Early Cretaceous), as well
as the Deccan and North Atlantic Provinces (Late Cretaceous to Paleocene) (Fig. 11, Larson and Erba, 1999).
The emplacement of these LIPs was associated with the
expulsion of vast amounts of mantle-derived volcanic gases,
which are typically dominated by H2O and CO2, but also
contain sulfur dioxide (SO2) and other trace gases (Symonds et al., 1994). The expulsion of such large quantities
of volcanic gases, particularly of CO2, provides a valid
mechanism to cause a rapid increase in atmospheric pCO2
and an associated increase of greenhouse forcing and global
temperatures during these time intervals. Although there is
still some dispute over the long-term effect of volcanogenic
CO2 emissions (Self et al., 2005), the hypothesis that LIP
magmatism contributed to changes in atmospheric composition and global climate is supported by the marine d13C
record, which indicates that the late Paleocene decline of
d13C was preceded by a notable rise in d13C values that
started in the mid-Paleocene. This rise in d13C coincides
with phases of intense volcanism in the Deccan area of India, and in the North Atlantic Igneous Province (Fig. 11).
Marine carbonates deposited at that time reached d13C values of around +2& (Fig. 11; Zachos et al., 2001). These
unusually positive values are consistent with a strong influx
of mantle-derived CO2 into the atmosphere because mantle
CO2 is enriched in 13C relative to CO2 derived from organic
matter 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 Early
Cretaceous and the Late Cretaceous/Paleocene may have
not been sufficient to sustain long-term greenhouse conditions, they may have served as a trigger for the subsequent
increase of oxidative weathering. This, in turn, would have
provided a feedback for the further rise in pCO2 and the
prolonged greenhouse episodes that followed the LIP volcanism. In this scenario, the elevated pO2 that prevailed before the onset of the LIP volcanism (i.e., in the Early and
Late Cretaceous) would have been an important factor that
determined the extent of the subsequent greenhouse phases.
After the cessation of LIP volcanism and the end of the
greenhouse periods in the Cretaceous and the Cenozoic,
atmospheric pO2 was at its minimum, but slowly increased
over the following tens of millions of years, as the result of a
decrease in oxidative weathering (associated with an increase in carbon burial) and the absence of the CO2-influx
from LIP magmatism. After the mid-Cretaceous, the cycle
of increasing pO2 and decreasing pCO2 lasted for 40 million years, but was interrupted by renewed LIP volcanism
at the end of the Cretaceous. The Cenozoic decline of
pCO2 and increase in pO2, on the other hand, continued
uninterrupted for 50 million years from the Middle Eocene into modern times. The prolonged absence of largescale volcanism during that time period may have been a
prerequisite for pCO2 and global temperatures to decrease
sufficiently to allow sea ice to form on the polar oceans,
ultimately triggering the onset of glaciation in both hemispheres in the latest Paleogene.
4.2.3. Effects of pO2 on animal evolution
Although it has been proposed that atmospheric pO2 has
a profound effect on the physiology and evolution of animals, and that gigantism in animals is linked to high pO2
(Graham et al., 1995; Clapham and Karr, 2012), our data
preclude such a link for the evolution of giant theropod
and sauropod dinosaurs. Considering the low pO2 that prevailed during the Cretaceous and, at least, part of the Triassic, it is more reasonable to assume that a high primary
photosynthetic productivity (i.e., food availability) caused
by high pCO2 and high global temperatures was the key
factor in the evolution of dinosaurs and their tendency to
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R. Tappert et al. / Geochimica et Cosmochimica Acta 121 (2013) 240–262
develop large body sizes. A similar conclusion can be drawn
for the Eocene radiation of large placental mammals, which
has previously been linked to high pO2 levels as well (Falkowski et al., 2005). Based on our data, however, this time
interval was also characterized by low pO2.
5. CONCLUSIONS AND FUTURE DIRECTIONS
Due to their ability to polymerize into highly resistant
organic macromolecules that are demonstrably representative of whole-plant carbon fractionation, resins can preserve pristine isotopic signatures over geological time.
This makes them ideal natural biomarkers for chemostratigraphic studies. As with all plant organic constituents, the
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
is directly comparable to that of primary plant biomass,
which suggests that the d13C of resins can be used as a measure of the d13C of bulk plant organic matter. To minimize
statistical biases caused by variations in the growth conditions of individual plants, it is necessary to analyze sufficiently large sample populations and to compare their
mean isotopic values (d13 Cresin
mean ). This is particularly important for the comparison of resins of different ages. Since
major diagenetic influences on the d13C of fossil resins are
negligible, variations in d13 Cresin
mean can be related to long-term
global environmental changes. In fact, our present compilation and analysis of d13C values from exhaustively sampled
resins, which is the largest data set of its kind, strongly suggests that the pO2 of ancient atmospheres is reliably represented in the isotopic composition of these secondary
metabolites.
Resins can be used as proxies of atmospheric composition for much of the Cenozoic and Mesozoic, and perhaps
even into the Paleozoic, given that resin-producing plants
have existed since at least the Carboniferous. Resins from
different localities, different botanical sources, and different
ages are directly comparable because they consist of similar
macromolecules, and they are resistant to post-depositional
isotopic exchange. These advantages are complemented by
the fact that the analysis of resins does not require extensive
sample treatment, which limits biases introduced by laboratory procedures. Therefore, resins can provide a consistent
and highly resolved terrestrial d13C record, particularly
when samples from stratigraphically and chronologically
well-characterized sedimentary units are analyzed.
To improve our understanding of the fractionation
behavior of plant organic matter in relation to atmospheric
composition, it is also necessary to conduct additional, systematically controlled growth chamber experiments on a
wider range of plants, including resin-producing plant taxa.
This information will help to more accurately quantify
compositional changes of the atmosphere that occurred
throughout Earth’s history. Additional constraints on the
composition of ancient atmospheres may also be gained
by combining isotopic information with careful analysis
of biological inclusions within amber—e.g., insects, plants,
and fungi—which are frequently preserved with exquisite
detail.
ACKNOWLEDGMENTS
We are grateful to Jim Basinger, Madelaine Böhme, Xavier
Delclòs, Michael Engel, George Mustoe, Philip Perkins, Guido
Roghi, Alexander Schmidt, and Chris Williams for generously providing samples used in this study. Thomas Stachel is thanked for
providing access to the FTIR spectrometer housed in his laboratory. We also thank Stacey Gibb and Gabriela Gonzalez for assistance at various stages, and we would like to thank the three
anonymous reviewers, whose constructive comments have improved this publication considerably. Much of this research was
supported by the Natural Science and Engineering Research Council of Canada (Discovery awards to A.P. Wolfe and K. Muehlenbachs, Postdoctoral fellowship to R.C. McKellar). J. Ortega was
supported by the Ministerio de Economı́a y Competitividad, Fulbright España, and FECYT.
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