The late Palaeocene to the middle Eocene (57.5 to 46.5 Ma)
recorded a total of 39 hyperthermals – periods of rapid global warming
documented by prominent negative carbon isotope excursions (CIEs) as well as
peaks in iron content – have been recognized in marine cores. Documenting
how the Earth system responded to rapid climatic shifts during hyperthermals
provides fundamental information to constrain climatic models. However,
while hyperthermals have been well documented in the marine sedimentary
record, only a few have been recognized and described in continental
deposits, thereby limiting our ability to understand the effect and record
of global warming on terrestrial systems. Hyperthermals in the continental
record could be a powerful correlation tool to help connect marine and
continental deposits, addressing issues of environmental signal propagation
from land to sea. In this study, we generate new stable carbon isotope data
(δ13C values) across the well-exposed and time-constrained
fluvial sedimentary succession of the early Eocene Castissent Formation in
the south central Pyrenees (Spain). The δ13C values of
pedogenic carbonate reveal – similarly to the global records – stepped
CIEs, culminating in a minimum δ13C value that we correlate
with the hyperthermal event “U” at ca. 50 Ma. This general trend towards more
negative values is most probably linked to higher primary productivity
leading to an overall higher respiration of soil organic matter during these
climatic events. The relative enrichment in immobile elements (Zr, Ti, Al)
and higher estimates of mean annual precipitation together with the
occurrence of small iron oxide and iron hydroxide nodules during the CIEs suggest
intensification of chemical weathering and/or longer exposure of soils in a
highly seasonal climate. The results show that even relatively small-scale
hyperthermals compared with their prominent counterparts, such as PETM, ETM2,
and ETM3, can leave a recognizable signature in the terrestrial stratigraphic
record, providing insights into the dynamics of the carbon cycle in
continental environments during these events.
Introduction
From the end of the Palaeocene, a period of global warming reached its
climax during the Early Eocene Climatic Optimum (EECO)
(Westerhold
and Röhl, 2009; Hyland and Sheldon, 2013). The EECO started ca. 53 Ma ago
and lasted until ca. 49 Ma ago
(Westerhold et al., 2018),
after which the climate began to cool (∼ Eocene–Oligocene
transition; Zachos et al., 2001, 2008).
Superimposed on, and coeval to, this globally warm epoch, brief periods of
pronounced global warming known as “hyperthermals” standout as anomalies
outside of background climate variability
(Kirtland-Turner et
al., 2014; Dunkley Jones et al., 2018). The Palaeocene–Eocene Thermal
Maximum (PETM; ∼56 Ma) was the first of these events to be
identified globally because of its exceptional magnitude and preservation in
both marine and continental deposits (Koch et
al., 1992). To date, for the late Palaeocene – early Eocene period, a total
of 39 hyperthermal events of lesser magnitude have been identified from
marine cores
(Lourens
et al., 2005; Sexton et al., 2011; Kirtland-Turner et al., 2014; Lauretano
et al., 2015, 2016; and Westerhold et al., 2018), among which the most prominent
and studied are the Eocene Thermal Maximum (ETM) 2 and 3, H2, and I1 and I2
events
(Cramer et al., 2003; Lourens et al., 2005; Nicolo et al., 2007; Lunt et al., 2011;
Deconto et al., 2012; Kirtland-Turner et al., 2014; Lauretano et al., 2016; and
Westerhold et al., 2017) (Fig. 1). In the marine stratigraphic record, these
events are primarily characterized by paired negative excursion in carbon
and oxygen isotope data exceeding background variability
(Cramer
et al., 2003; Nicolo et al., 2007; Zachos et al., 2008; Sluijs and Dickens,
2012; and Lauretano et al., 2016), i.e. typically with amplitude greater than
the standard deviation (SD) of pre-hyperthermal background values.
Late Palaeocene and early Eocene benthic carbon isotope record from
Sites 1209, 1258, 1262, and 1263. Top of Chron C22r and top of T. orthostylus
zone from site 1263 from
Westerhold et al. (2017). Hyperthermal nomenclature from
Cramer et al. (2003),
Lauretano
et al. (2016), and Westerhold et al. (2017). Castissent Fm. extension in
green.
In deep marine settings, the carbon isotope excursions (CIE) are typically
paired with an increase in iron concentration and decrease in carbonate
content, indicating ocean acidification potentially linked with high
atmospheric CO2 concentrations
(Nicolo
et al., 2007; Slotnick et al., 2012; and Westerhold et al., 2018). In coastal
marine sections, early Eocene hyperthermal events are generally associated
with an enhanced flux of terrigenous material, interpreted as linked to
accelerated hydrological cycle and higher seasonality
(Schmitz
et al., 2001; Bowen et al., 2004; Nicolo et al., 2007; Slotnick et al.,
2012; Payros et al., 2015; and Dunkley Jones et al., 2018), although several
studies document a spatially heterogeneous hydrological climatic response
during the PETM
(Bolle
and Adatte, 2001; Kraus and Riggins, 2007; Giusberti et al., 2016; and
Carmichael et al., 2017). In fluvial systems, the abrupt warming of the PETM
was found to be associated with expansion and coarsening of alluvial facies
combined with an increase in the magnitude of flood discharge
(Foreman
et al., 2012; Pujalte et al., 2015; and Chen et al., 2018), as well as enhanced
pedogenesis (Abels et al., 2012). Yet,
how continental systems reacted to the other, smaller-magnitude
hyperthermals of the early Eocene remains to be documented. In particular,
because of the subaerial nature and lateral preservation dynamics of
alluvial systems (e.g.
Foreman and Straub, 2017; Straub and Foreman, 2018), the extent to which
fluvial successions can provide complete and faithful archives of past
climatic events, especially those with the smallest magnitudes, is still
largely unknown
(Foreman
and Straub, 2017; Trampush et al., 2017; and Straub and Foreman, 2018).
Addressing this question is particularly critical for studies focussing on
environmental signal propagation in source-to-sink systems
(e.g.
Castelltort and Van Den Driessche, 2003; Duller et al., 2019; Romans et al.,
2016; and Schlunegger and Castelltort, 2016), which require high-resolution
continental-marine correlations such as those provided by the PETM
(e.g. Duller et al., 2019) or
by other hyperthermals of the early Eocene.
Simplified situation and geological map of the study area with
main depositional palaeo-environments
(e.g. Nijman, 1998). The
Castissent Fm. is a prominent fluvial unit particularly well exposed in the
Noguera Ribagorçana and Isáabena river valleys. (1) Chiriveta section (2) Mas de Faro (3) La Roca section. Main palaeoflow directions indicated in
orange (from Nijman and Puigdefabregas, 1978). Regional
map after Teixell (1998).
To address these issues, we explored the geochemical signature (carbon and
oxygen stable isotopes; major and trace elements) and the sedimentology of
the fluvial deposits of the Ypresian Age Castissent Fm. (south central
Pyrenees, Spain, Fig. 2). First, we generated a new carbon isotope profile
from a palaeosol succession rich in carbonate nodules across the Castissent
Fm. in order to compare these results with a global δ13C
record. The data suggest that this fluvial succession preserves a record of
hyperthermal “U” event at ca. 50 Ma, providing important constraints to its
depositional the age. Second, we used the major and trace element
composition of bulk floodplain material in order to explore the climatic
impact of such a hyperthermal, including empirical reconstructions of mean
annual precipitation, allowing us to discuss soil dynamics during global
warming. This study identifies for the first time in a continental
succession an event so far only recorded in marine sediments, thereby
demonstrating the global breadth of these climatic events and the
complementarity of oceanographic and terrestrial archives.
(a) Time constraints on the Castissent Fm. MP zone from the
continental section from Checa Soler (2004)
and Payros et al. (2009). SBZ and NP in the Campo section from
(Schaub, 1966, 1981; Kapellos and Schaub,
1973; and Tosquella, 1995), magnetostratigraphy from Bentham
and Burbank (1996). SBZ in El Pueyo section from
Payros et al. (2009).
Magnetostratigraphy in El Pueyo from Poyatos-Moré (2014). (b) Extended map of the study area. For map legend and references,
see Fig. 2.
Geological setting
The Castissent Formation comprises fluvial deposit of Ypresian age cropping
out in the Tremp-Graus Basin (South Pyrenean foreland basin; Marzo et al., 1988,
Fig. 2). The Castissent Fm. is defined by its prominent overall sand-rich
character and is composed by three coarse-grained channel complexes
(labelled as Members A, B, and C) separated by four marine incursions (M0 to
M3) inferred from the observation of marginal coastal bioclast-rich horizons
developed up into the upper deltaic plain and correlative with finer
dark-grey mudstones and calcretes in the fluvial segment of the Castissent
(Marzo et al., 1988).
This major fluvial progradation is correlated westwards with deep-water
turbidite sequences of the Arro and Fosado Formations in the Ainsa Basin
(Fig. 3, Mutti et al., 1988; Nijman
and Nio, 1975; Nijman and Puigdefabregas, 1978; and Pickering and Bayliss,
2009). In the upstream, eastern counterparts of the Castissent Fm., the
channel complexes are intercalated with yellow to red coloured palaeosols.
Sub-spherical to slightly elongated carbonate nodules with diameters
ranging from 1 mm to 4 cm are omnipresent in the palaeosols (Fig. S1 in the Supplement).
Studies of the Castissent Fm. tentatively attributed their occurrence to an
important pulse of exhumation and thrust activity in the hinterland at ca. 50 Ma, in possible combination with a late Ypresian sea-level fall
(Puigdefabregas
et al., 1986; Marzo et al., 1988; Whitchurch et al., 2011; and Castelltort et
al., 2017), both resulting in reduced available accommodation space
enhancing progradation and amalgamation
(Chanvry et al., 2018).
The Chiriveta section, encompassing the Castissent Fm., is situated in a
continental palaeogeographic position prone to pedogenesis and slightly
off-axis from the more “in-axis” amalgamated sand-rich-type section of
Mas de Faro (Fig. 2); for palaeo-position and correlation see also Figs. 10
and 12 in Marzo et
al. (1988).
In the Chiriveta location, stratigraphic constraints are limited to the
identification of European Mammals zone MP10 (Badiola et al., 2009), which
provides an age range of 50.73 to 47.4 Ma (GTS2012). This age span is
refined by bio- and magnetostratigraphic data from the Castissent Fm.
outcrops of the Campo location, about 40 km further west
(Kapellos
and Schaub, 1973; Tosquella, 1995; Bentham and Burbank, 1996; Tosquella et
al., 1998; and Payros et al., 2009) (Fig. 3). Because of its outcropping extent,
the Castissent Fm. has been mapped from west to east across these sections
(Nijman
and Nio, 1975; Nijman, 1998; Poyatos-Moré, 2014; and Chanvry et al., 2018).
The low slope of the Castissent Fm. (ca. 2.3×10-4 m m-1; see
Supplement Table S1) indicate an elevation drop of ca. 1 m between the
Chiriveta section and the Campo section. Given average flow depths of
3.75 m in the Castissent channels based on measurement in the Chiriveta and
La Roca sections, we thus assume no significant time lag of deposition
between both sections. In the Campo section, Kapellos
and Schaub (1973) find the transition between the D. lodoensis and the T. orthostylus nannoplankton
(NP) zones at ca. 200 m below the base of the Castissent Fm. and the transition
between the T. orthostylus and the D. sublodoensis NP zones in the transgression ca. 100 m above the
uppermost member of the Castissent Fm. This indicates that the Castissent
Fm. was deposited during NP13. Magnetostratigraphic data of the same section
by Bentham and Burbank (1996) place the transition between
the C22r and C22n magnetozones closely above the top of the Castissent Fm.
We thus used the recent astrochronological age models of
Westerhold et al. (2017), which obtain numerical ages of 50.777±0.01 and 49.695±0.043 Ma for the base and top of C22r, respectively, and obtain a numerical
age of 50.534±0.025 Ma for the base of NP13 based on the Ocean Drilling Program site 1263.
Considering the data available and their resolution, we suggest a
depositional age span between 50.5 and 49.7 Ma for the Castissent Fm.
(reported in green on Fig. 1). According to global isotopic records (Fig. 1), this period was marked by four hyperthermals labelled S/C22rH3, T/C22rH4,
U/C22rH5, and V/C22nH1
(Cramer
et al., 2003; Lauretano et al., 2016; and Westerhold et al., 2017).
Material and methodsSampling
A total of 74 samples were collected from the lower Eocene Chiriveta section
for geochemical studies. All samples consist of floodplain material and were
taken below the weathering depth (∼50 cm), with an average
resolution of 1 m. Resolution was increased by a factor of 2 in specific
horizons such as red beds. When important sandbodies occurred, lateral
equivalent floodplain material or intercalated palaeosol horizons were
sampled. Each sample was split in two aliquots, one for major and trace
element analysis and the other for carbon and oxygen stable isotope analysis
on pedogenic carbonate nodules. The carbonate nodules were extracted from
the bulk palaeosol material by sieving and then cleaned by repeated washes
with deionized water in an ultrasound bath. From each cleaned nodules set,
subsamples of one to four nodules were taken, leading to a total of 149
subsamples of pedogenic carbonate nodules.
Carbon and oxygen stable isotopes
Pedogenic carbonate nodules were crushed and powdered in an agate mortar and
analysed for stable carbon and oxygen isotope composition at the Institute
of Earth Surface Dynamics of the University of Lausanne (Switzerland) using
a Thermo Fisher Scientific (Bremen, Germany) carbonate-preparation device
and Gas Bench II connected to a Thermo Fisher Delta Plus XL isotope ratio
mass spectrometer. The carbon and oxygen isotope compositions are reported
in the delta (δ) notation as the per mil (‰)
isotope ratio variations relative to the Vienna Pee Dee Belemnite standard
(VPDB). The analytical reproducibility estimated from replicate analyses of
the international calcite standard NBS-19 and the laboratory standard
Carrara marble was better than ±0.05 ‰ (1σ)
for δ13C and ±0.1 ‰ (1σ) for
δ18O.
Major and trace element composition
Fifty-two bulk palaeosol samples were analysed for major and trace elements
using X-ray fluorescence (XRF) spectrometry. Crushed bulk powders (< 80 µm) were mounted in a plastic cup covered by a thin polypropylene
film (4 µm thick) and analysed in the laboratory with a Thermo Fisher Niton
XL3t® portable XRF analyser fixed on a test stand. Analyses
were performed with a beam diameter of 8 mm, to determine the concentrations
of 34 major and trace elements (from Mg to Au). Each measurement took 120 s,
consisting of two 60 s cycles on four different filters (15 s on low,
main, high, and light ranges), operating the X-ray tube at different voltages
to optimize the fluorescence and peak/background ratios of the different
elements. The limits of detection were of tens of parts per million for most elements, except
for Mg, Si, and Al which are at wt % level. Sodium is too light to be
detected. The acquired spectra were transferred to a computer using NDT
software version 8.2.1. (Thermo Fisher Scientific, Waltham, Ma, USA). The
same material has been analysed for 23 major and trace elements on
fused and pressed discs, respectively, using a PANalytical PW2400 XRF
spectrometer with a copper (Cu) tube at the University of Lausanne to
cross-calibrate the compositions measured with the Niton XL3t®
portable XRF analyser.
Mean annual precipitation
The mean annual precipitation estimate (MAP) used in this study was
estimated from the empirical relationship between MAP and
CaO/Al2O3 ratio for Mollisols from a national survey of North
American soils according to the following equation: MAP (mm) =-130.9×ln(CaO/Al2O3)+467
(Sheldon et
al., 2002). CaO and Al2O3 concentrations were measured on bulk
palaeosol material. Climate linked to the MAP estimate was classified based
on the following boundaries: arid to semiarid at 250 mm and semiarid to
subhumid at 500 mm (Bull, 1991).
Grain-size estimation
The relative grain-size variation in the sediment samples was estimated from
their major element compositions. Si, Ti, and Zr are more concentrated in the
coarse fraction of the sediment as they are found in larger mineral grains,
whereas Al is more concentrated in the finer fraction of the sediment
because it is mostly linked to clay minerals
(Lupker et al.,
2011, 2012; Croudace and Rothwell, 2015). Grain size variation throughout
the section was estimated using Si/Al, Ti/Al, and Zr/Al ratios, therefore, an
increase in these ratios suggests a relative increase in the proportion of
coarser material in the sample.
Correlation with target curves
The measured δ13C dataset was compared with a time-equivalent
ODP 1263 global δ13C record reported by
Westerhold et al. (2017) using the AnalySeries software (Paillard et al.,
1996). The δ13C record of site 1263 was favoured over those of
ODP 1209 and 1258 covering the Castissent Fm. time period, because it is
continuous and has a higher resolution. Correlations between the δ13C record of site 1263 and the δ13C record of the
Chiriveta section were performed in order to optimize the Pearson
correlation coefficient (r) and to minimize abrupt variations in
sedimentation rates. Well-defined peaks in both δ13C records
were used as tie points for the correlation and the number of tie points was
kept to a minimum (< 10) so as to not force the correlations.
Field images of the Chiriveta section (42∘7′56.57′′ N,
0∘41′19.45′′ E). (a) Outcrop view of Members A and B of the
Castissent Formation. (b) Close-up view of the upper part of Castissent A
Member. Fluvial channel-fill deposits, intercalated in reddish floodplain
and overbank deposits and regional marine incursions (M1). (c) M0, first
marine incursion at the base of the Castissent Fm. described by Marzo et al. (1988) expressed in the Chiriveta section by a tidal-influenced coarse
sandstone with herringbone cross-stratification. (d) Yellow mottled
palaeosol between CIE C and D. (e) Red floodplain interval equivalent of the
CIE C. (f) 2 m thick grey interval interpreted as poorly drained brackish
water facies and equivalent to the marine incursion M1. (g) An ∼6 m thick laterally extensive Castissent B sandbody incised in the underlying
floodplain deposits. (h) Mottled silt, interpreted as pedogenetic fluvial
channel overbank deposits.
ResultsSedimentology of the Castissent Formation at Chiriveta
We describe here the section logged and sampled in this work (Fig. 4). At
Chiriveta, the Castissent Fm. is a palaeosol-rich succession, which shows
greyish-yellow to red-brown mottled floodplain palaeosols (Fig. 4a–b),
corresponding laterally to thick, medium to coarse-grained quartz-rich
channel-fill deposits (width/depth ratio =20–50; Marzo et al., 1988)
and overbank deposits flowing parallel to the main structures of the
growing Pyrenean orogeny
(Marzo et al., 1988).
At the base of the section, the first marine incursion M0 is situated at the
top of a 20 m thick coarse-grained tidal bar deposit with herringbone
cross-stratifications and oyster shells (Fig. 4c). In the Chiriveta section,
the Castissent Member A is a 48 m thick interval comprising two main
medium-grained sandbodies of light colouration of 5.40 and 1.5 m in thickness
respectively. Bedforms observed in the first sandbody have a mean height of
24 cm (n=9). The second marine incursion M1 is located at 48 m just below
the Castissent B Member and consists of a 2 m thick grey interval
interpreted by Marzo
et al. (1988) as brackish–lagoonal water facies (Fig. 4b–f). The Castissent B
Member (Fig. 4g) is a 12 m thick and laterally extensive (width/depth ratio
≥250; Marzo et
al., 1988) amalgamated sandbody with a micro-conglomeratic erosive base.
Grain size is overall larger than in Member A and ranges from fine sand to
large pebbles. Sandbody tops show a fining-upward trend and are capped by
mottled siltstone packages. Mottled siltstone layers are interpreted as
pedogenized overbank deposits based on roots traces and their capping
relationship with underlying sandbody deposits (observed at 26, 76, 89,
and 96 m in Figs. 5 and 4h). More regular and sheet-like sandbodies
interbedded with mottled siltstone layers are observed upwards. The section
ends with a 23 m thick, medium to very coarse, tidally influenced sandstone
deposit interpreted as the equivalent M3 marine incursion by
Marzo et al. (1988).
Although Castissent Member C was not interpreted by
Marzo et al. (1988)
in this section, a 2 m thick fine-grained sandbody at ca 80 m in our
section could be the condensed lateral equivalent of it (Fig. 5).
Isotopic and geochemical data from the Chiriveta section. For the
isotope dataset, the curves passes through the mean values at each sample
position. Samples with minimum in δ13C values below 1 and 2
standard deviations are labelled A to F. Mean annual precipitation (MAP) was
estimated from the empirical relationship between MAP and CaO to
Al2O3 ratio
(Sheldon et
al., 2002).
Stable isotopic record
Carbon and oxygen isotope ratios from the carbonate nodules are presented in
Fig. 5. The δ13C values vary between -10.9 ‰ and
-1.9 ‰ with a mean value and 1 SD of -7.7±1.6 ‰. Six CIEs (named A to F in Fig. 5 and colour coded in
Fig. 6) are more negative than -9.3 ‰ (i.e. the mean
value – 1 standard deviation) amongst which one (CIE D) is below 2 SDs. The
values are -9.6 ‰, -9.8 ‰, -9.9 ‰, -10.9 ‰, -9.9 ‰, and -9.4 ‰
for CIEs A to F respectively. At the bottom of the section, CIE A is
followed by a relatively constant interval of mean δ13C values.
CIE B, situated in the first red bed, marks the beginning of a stepped
δ13C trend (around ±1 SD) leading to the minimum CIE D.
The second part of the section shows two more CIEs separated by the highest
δ13C value at 74 m. CIE F is the least prominent of all CIEs.
The δ18O values vary between -7.0 ‰ and
-5.0‰ with a mean value of -6.0±0.4 ‰, which makes them less dispersed than the δ13C record. Nine negative oxygen isotope excursions are more negative
than the mean value minus 1 SD, amongst which one is below 2 SD reaching a
minimum value of -6.8 ‰ at 19 m. The oxygen isotope
excursions do not correspond with CIEs described above.
Major and trace elements
Titanium (Ti), aluminium (Al), and zirconium (Zr) concentrations measured on
bulk palaeosols are plotted in Fig. 5. These elements are commonly
considered as immobile and are expected to concentrate in more weathered
soils. Ti values vary between 0.18 % and 0.52 % with a mean value of
0.34 % and a standard deviation of 0.08. Al values vary between 3.03 % and
9.35 % with a mean value of 5.85 % and a standard deviation of 1.53. Zr
values vary between 67 and 204 ppm with a mean value of 128 ppm and a
standard deviation of 35. Mean annual precipitation (MAP) estimates values
vary between 185 and 754 mm yr-1 with a mean value of 376 mm yr-1 and a standard
deviation of 111. Ti, Al, Zr, and MAP show a similar trend starting from the
base of the section with a global increase for all values toward CIE C and a
decrease afterwards. All CIEs show higher values of Ti, Al, Zr, and MAP except
CIE F. Based on Bull (1991), an average value of 387 mm yr-1
for the MAP in the Chiriveta section represents a semi-arid climate (Fig. 5).
All CIEs show an increase in MAP.
DiscussionCarbon and oxygen isotopic recordIdentifying the CIE
In continental successions, the carbon isotope composition of pedogenic
carbonate nodules – which consists of calcareous concretions between 1 mm
and 4 cm diameter formed in situ in the floodplain – have been shown to be
sensitive to environmental conditions during their formation (e.g.
Millière
et al., 2011a, b) and are therefore a promising tool to track how
environments respond to carbon cycle perturbation. The carbon isotope
composition of the soil carbonate nodules depend on the δ13C
value of the atmospheric CO2 and soil CO2, which in turn is a
function of the δ13C of the atmospheric CO2 and the
overlying plants, as well as the soil respiration flux and the partial
pressure of atmospheric CO2
(Cerling,
1984; Bowen et al., 2004; Abels et al., 2012; and Caves et al., 2016).
The δ13C vs. δ18O diagram for the pedogenic
carbonate nodules from the Chiriveta section (r=-0.26, n=149)
suggests a good preservation of the primary isotopic signal (Fig. 6), with
an average value of δ13C=-7.7±1.6 ‰ similar to mid-latitude late Palaeocene to Eocene
continental δ13C values (excluding the PETM samples) observed
elsewhere (e.g. McInerney and Wing, 2011; and
references therein) and a spread comparable with δ13C values
from carbonate nodules analysed for the same period in the Bighorn Basin
(Bowen et al., 2001). Figure 6 emphasizes that lower Eocene carbonate nodules
display overall more negative δ13C values than the Holocene
nodules, which is consistent with a large compilation of data from eastern
Eurasia (Caves Rugenstein and
Chamberlain, 2018). Pre-PETM δ18O values from carbonate nodules
from the same area (-4.5±0.4‰)
(Hunger, 2018) show values similar to measurements from the Chiriveta
section (-6.0±0.4‰). Oxygen
and carbon isotopes are not coupled during hyperthermal events in
continental record as already observed by Schmitz
and Pujalte (2003) Bowen et
al. (2001) for the PETM isotopic excursion. Though the precise mechanisms
that produce stable δ18O during CIE are still debated,
mid-latitude precipitation δ18O appears to be relatively
insensitive to changes in atmospheric pCO2 and warming, particularly in
greenhouse climates (Winnick et al., 2015). In
contrast, the stable δ18O values of soil carbonates from the
Pyrenean foreland basin (-5.5±0.9‰) are likely
additionally stabilized by its position close to the coast
(Cerling, 1984; Kukla et al., 2019)
compared for example to those of the Bighorn Basin (-9.0±0.6‰). This is in line with a more continental
palaeogeographical position of the Bighorn Basin compared to the Tremp-Graus
Basin at the time (Seeland, 1998).
Continental δ13C and δ18O values from
the early Eocene Castissent Fm. in the Chiriveta section (this study)
plotted with pre- and syn-PETM δ13C and δ18O
values from the same area (Khozyem Saleh, 2013; Hunger,
2018) and pre-, syn-, and post-PETM values from the Bighorn Basin
(Bowen et al., 2001) as well as
recent pedogenic carbonate isotopic values
(Cerling and Quade, 1993; Gallagher
and Sheldon, 2016).
A hyperthermal event recorded in marine sediments is defined by paired
negative carbon and oxygen stable isotope excursions that are more negative
than the mean value minus 1 SD
(Kirtland-Turner et al., 2014). This
definition may not be applicable to continental deposits, because
continental systems respond differently than marine systems to the carbon
cycle perturbations. Though the marine δ13C record is thought
to record the global CO2δ13C, the δ13C value
of the marine dissolved inorganic carbon is also influenced by dissolution
of carbonates at depth (McInerney and Wing, 2011). In
contrast, δ13C in pedogenic nodules varies with soil
properties, atmospheric and soil pCO2 and δ13C, and the
rate and nature of carbon input and/or output by soil respiration
(Bowen et
al., 2004; Sheldon and Tabor, 2009). These processes create complexities in
estimating CIEs in soil carbonate nodules and in marine carbonates
(McInerney and Wing, 2011). Nevertheless, we used the hyperthermal definition from Kirtland-Turner et al. (2014) as a starting point to filter the high-resolution variations in the Chiriveta section. We identify 16 samples with CIE values more negative
than the mean minus 1 SD. Among these 16 samples, we recognized six discrete CIEs
(named A–F in Figs. 5 and 7). Both marine incursion M1 and M2 show an
abrupt shift from -9 ‰ to -10 ‰ in continental δ13C values towards more (positive) marine values of -4 ‰ to
-2 ‰; this points to a progressive higher contribution
of seawater to the formation of the carbonate nodules.
(a) Scaling of the Chiriveta isotopic section with the time
equivalent interval of site 1263 (Westerhold et al., 2017). The correlation
was calculated using the AnalySeries software (Paillard et
al., 1996) and centred on CIE D and hyperthermal U. Mean, minus 1 and 2 SD
lines on the global record were calculated sets over the selected time
period. The correlation coefficient (r) between the two curves is 0.65. (b) Hyperthermal U amplitude in palaeosol carbonate and benthic foraminifera
(inset B after Abels et
al., 2016).
Six correlation options with the global record were explored in the
time window of the Castissent Fm. (Figs. S2 and S3). The
correlation presented in Fig. 7a was favoured as it shows (i) reasonable
sedimentation rates variations, (ii) a similar amplitude to the CIE in the
global record, and (iii) the highest correlation coefficient (r=0.65, n=71). Moreover, it plots along the same trend regarding hyperthermal CIE
amplitudes in marine and continental environments, suggesting a common
mechanism of global climatic change with events I1, I2, H2, and ETM2 (Fig. 7b). Based on these observations and obtained correlation, we suggest that
only hyperthermal U is preserved in the Chiriveta section and that it is
correlated with CIE D. Sedimentation rate obtained with the favoured
correlation (Fig. 7) varies between 0.1–0.29 mm yr-1, consistent with
sedimentation rates reported for other Eocene floodplain successions
(Kraus and Aslan, 1993). The correlation
coefficient of r=0.65 suggests an overall good signal preservation in the
studied continental section for a 40 kyr climatic event.
Mechanisms causing the CIE
An increase in temperature could potentially release a significant amount of
CO2 into the atmosphere
(Trumbore et al.,
1996; Melillo et al., 2014). The amplitude and duration of Eocene CIEs are
approximately 30 % of the one recorded for the PETM; we hypothesize that
the climatic effects of smaller-scale hyperthermals can be linearly scaled
to the PETM. Based on this assumption and in order to get a rough
approximation without considering a nonlinear sensitivity response, a
smaller-scale hyperthermal would imply a release of approximately 500 to
1500 Gt of carbon to the ocean and atmosphere reservoir and a global
temperature rise of about 1.5–2.5 ∘C. This estimation corresponds
to the 1500–4500 Gt of carbon released during the PETM, causing a rise of
5–8 ∘C (Bowen et al., 2006),
and is in line with previous estimations of ∼3 and
∼2∘C warming for ETM2/H1 and H2 events
respectively (Stap et al., 2010).
A release of 500 to 1500 Gt of carbon in the form of methane would imply a
marine CIE of 0.8 ‰ to 2.3 ‰ or 0.3 ‰ to
0.9 ‰ if the carbon origin is dissolved organic carbon
(DOC) (Sexton et al., 2011).
The latter seems more plausible regarding the observed amplitude of
∼1 ‰ measured in the marine record for
hyperthermal U
(Westerhold et al.,
2017) and the supposed origin linked to the oxygenation of deep-marine DOC
of post-PETM hyperthermals
(Sexton et al., 2011). A
global shift of -1 ‰ in δ13C can however not
fully explain the 3 ‰ shift in δ13C observed
in this study.
Components influencing the δ13C values of pedogenic
carbonate nodules. Mean early Eocene bulk marine carbonate and small-scale
hyperthermal (all except PETM) are from
Westerhold et al. (2018).
Fractionation value between organic matter and carbonate nodules are based
on Sheldon and Tabor (2009). All other
fractionation values are based on
Koch et al. (1995). Mean carbonate
nodule values come from this study.
The δ13C mean value in the Chiriveta section is -7.7±1.6 ‰. This value reflects an overall equilibrium with
a mean atmospheric CO2 of -7 ‰
(Koch et al., 1995) and is coherent with
pre-PETM δ13C values of -7.1±0.9‰
found in the same area (Hunger, 2018; Fig. 6). It is
possible to calculate from the (small-scale) hyperthermal δ13C
excursions in the marine environment the shift to be expected in soil
carbonate nodules by using known fractionation coefficients
(Koch
et al., 1995, 2003); the expected δ13C value in carbonate
nodules, only considering the respiration of organic matter,
is -11 ‰ (Fig. 8). This value is within the range of
those measured from the Chiriveta section, where some nodules reach values
as low as -10.9 ‰. We suggest that the bacterial
respiration of organic matter, enhanced by warmer temperatures
(e.g. Davidson and
Janssens, 2006; Trumbore et al., 1996), may also have contributed to the
lower δ13C values of nodules during the CIEs (Fig. 8). On
geological timescales, soil organic carbon can be considered at steady state
with equal organic carbon inputs and outputs from the soil
(Koven et al., 2017). Respiration (carbon output after
mineralization as CO2) is thought to be more sensitive to global
warming than gross primary productivity (organic carbon input as organic
matter), leading to a depletion of the total soil carbon pool with time
during transient global warming events; although the precise sensitivity of
gross primary productivity remains poorly constrained
(Davidson and Janssens, 2006). Large uncertainties
remain about carbon dynamics and their timescale in the soils during climate
changes. Parameters such as the vegetation type
(Klemmedson, 1989), temperatures (Koven
et al., 2017), soil geochemistry
(Torn et al., 1997; Doetterl et
al., 2015), and soil water content (Davidson
et al., 2000) have been shown to be important controlling factors within
historical timescales.
Considering these caveats, we estimate the maximum possible contribution of
enhanced soil carbon respiration to negative δ13C excursions
during the CIEs. Using typical values for the organic carbon reservoir
comprising fast and slow cycling carbon in soils in arid to semi-arid
ecosystems of 5.6–19.2 kg C m-2
(Klemmedson, 1989; Raich and
Schlesinger, 1992), respiration fluxes starting at a steady-state value of
0.5 kg C yr-1 and a respiration rate sensitivity ca. 5 % per degree
(Raich and Schlesinger, 1992) (Q10=1.5), we estimate that all of the organic carbon in soils would be consumed
within 250 to 850 years, given an increase of 1 ∘C and without
changing the carbon input rate. Though there are a number of assumptions in
this first-order estimate, the timescale of soil carbon depletion is
substantially shorter than our estimate of the timescale of the CIE
(∼36 kyr) (Fig. 7). As evidenced by this calculation, an
increase in soil respiration triggered by warmer temperatures cannot be the
sole mechanism driving the CIE shift over multi-millennial timescales.
Instead, we suggest that during these transient warmings, this mechanism is
associated with a high primary productivity – resulting in a greater input
of carbon to the soil – leading to an overall higher soil respiration of
organic matter. Coupled with lower atmospheric δ13C during
hyperthermals, this mechanism caused a pronounced CIE in soil carbonate
nodules.
Geochemical signature of hyperthermal events
Major and trace elements compositions of floodplain sediments is a function
of river dynamics, climate, and sediment grain-size
(Lupker et al., 2012;
Turner et al., 2015). Based on the CIEs, we defined six intervals showing a
relative enrichment (up 10 % to 30 % compared to the average value) in
immobile elements such as Ti, Al, and Zr (Fig. 5). To ensure that major and
trace concentrations are not grain-size biased, we plotted grain-size
proxies Si/Al, Ti/Al, and Zr/Al
(Lupker et al., 2012;
Turner et al., 2015), which all exhibit a relatively stable trend, not
connected with the immobile element concentrations (Fig. S4). The enrichments in Ti, Al, and Zr suggest mature palaeosols with
potential intense weathering due to enhanced humid climatic conditions; but this
may also correspond to a longer exposure time on a stable floodplain,
allowing leaching of mobile elements and relative enrichment of immobile
elements (Sheldon and Tabor, 2009).
Pedogenic nodules are frequent in well-drained soil profiles associated with
a climate regime where the potential evapotranspiration is greater than the
mean annual precipitation rate
(Slessarev et al., 2016) and with a mean annual precipitation < 800 mm yr-1
(Cerling, 1984; Retallack,
1994; and Sheldon and Tabor, 2009). These conditions correspond to climate
ranging from arid to subhumid conditions
(Hasiotis,
2004; Prochnow et al., 2006; Hyland and Sheldon, 2013). This agrees with MAP
values obtained for the palaeo-precipitation estimate (Fig. 5) and with a
smectite/kaolinite > 1 assemblage dominating some of the studied
soils (Nicolaides, 2017, Table S2); all
suggestive of a semi-arid to subhumid climate with seasonal humidity
(Arostegi et al., 2011). Associated with CIEs C and D in red bed deposits, sub-millimetric iron-oxide
and iron-hydroxide nodules made of concentric hematite and goethite were found
together with carbonate nodules (Fig. S1). This suggest a seasonal climate
as hematite forms under more arid soil condition than goethite
(Kraus and Riggins, 2007). Together,
these observations are in line with an acceleration of the hydrological
cycle and a higher seasonality, as has been observed during the PETM, H1,
H2; I1 and I2 hyperthermals
(Bowen
et al., 2004; Nicolo et al., 2007; Slotnick et al., 2012; and Dunkley Jones et
al., 2018). Therefore, combined with CIEs, we suggest that small-scale
hyperthermals in continental records can be recognized by an increase in the
weathering index (Hessler et al., 2017) and by an
increase in the immobile element concentrations, both related to an increase
in precipitation intensity.
High-resolution hyperthermal signal
The high-resolution isotopic and elemental record of the Chiriveta profile
allow us to highlight the dynamics and variability in a hyperthermal event.
We do not observe a unique peak in δ13C but rather a stepped
isotopic signal suggesting, together with above-discussed geochemical data,
a climatic oscillation alternating with variably intense precipitations and
leaching conditions during a climax spanning ca. 150 kyr (interval CIE B to
D). Such a climatic behaviour, was already described for the PETM, during
the pre-onset excursion (Bowen et al.,
2015) and in the core CIE of the PETM
(Giusberti et al., 2016).
Moreover, the δ13C climax (CIE D) does not correspond to the
highest concentrations of immobile elements or maximum MAP estimates, which
we estimate occur during CIE C, which predates the CIE D by ca. 50 kyr (Fig. 7) . The minimum δ13C value therefore does not seem to be coeval
with the most extreme climatic response, suggesting a complex environmental
response. However, because sedimentation in floodplain depositional settings
is a function of the channel position and flood frequency, the relative
concentration of elements may reflect the changes in river dynamics instead
of climatic variability, which could explain the mismatch between minimum
values in CIE and the climatic response. More high-resolution hyperthermal
studies in coeval continental sections are needed to better understand the
relationships between proxies.
Possible implication for the preservation potential of hyperthermals in
continental sections
Major events such as the PETM have proven to be detectable in both marine
and continental environments
(e.g. Abels et al., 2016; Koch et al., 1992), but the signal and preservation
potential of smaller-scale climatic events (e.g. hyperthermal events L to W
in Lauretano et al., 2016) may be
more difficult to detect (Foreman and Straub, 2017) because of the inherent highly dynamic nature of sedimentation in fluvial deposits. To address this issue in the present case study, we
calculated the compensation timescale (Tc) of the Castissent Fm. Tc is a
timescale characteristic of an alluvial basin below which stratigraphic
signals with shorter durations may be of autogenic origin, thereby giving a
scale below which allogenic forcing should be interpreted carefully
(Wang
et al., 2011; Foreman and Straub, 2017; and Trampush et al., 2017). In other
words, an external forcing signal with a duration smaller than Tc will be
challenging to identify from background variability; the external forcing
must be therefore of a duration longer than Tc and optimally twice Tc
(Foreman and Straub, 2017). The Tc max can be
calculated by dividing the topographic roughness or maximum channel depth by
the average subsidence or deposition rate (Wang et
al., 2011). Using an average sedimentation rate of 0.17 mm yr-1 and an average
channel depth of 3.75 m, we obtained a mean Tc of 22 000 years, which means
that hyperthermal events of 40 kyr duration (timescale of hyperthermal U
and preceding CIE) have the potential to be recorded despite fluvial system
dynamics. Our estimate of preservation potential assumes steady
sedimentation rates throughout the section. But, sedimentation in
terrestrial records is not uniform (steady) but rather highly variable,
resulting in spatial and temporal changes in facies and deposition rates
ranging from < 0.1 to 1–2 mm yr-1
(Marriott
and Wright, 1993; Bowen et al., 2015; Kraus et al., 2015). However, mean
accumulation rates give a reasonable estimate approximating more realistic
(i.e. variable) sedimentation rates as observed in the Bighorn Basin
(Bowen et al., 2015). Additionally, we
analyse the vertical movement of the nearby structures to evaluate their
potential influence on disrupting deposition at Chiriveta during Castissent
times. The Chiriveta section was deposited near or at the axis of the
Tremp-Graus basin (Nijman, 1998),
which is bounded by the Bóixols thrust in the north and the in the south
(Marzo et al., 1988).
The Tremp-Graus basin is transported as a piggyback basin on the Montsec
thrust emerging at the time approximatively 4 km south of the studied
section (Nijman, 1998). In the basin
axis, subsidence is the highest with rates of 0.1 to 0.29 mm yr-1 (this study
and Marzo et al.,
1988). Taking into account a vertical movement rate of the Montsec thrust
of 0.03 to 0.1 mm yr-1 during the Castissent time-interval (based on a
horizontal displacement of 7 km, a period of activity lasting 26 Ma, and a
thrust dip between 6 and 20∘;
Farrell
et al., 1987; Nijman, 1998; Clevis et al., 2004; and Whitchurch et al., 2011),
we estimate that the vertical displacement is no more than equal to
sedimentation rates in the basin axis. This is consistent with the general
absence of growth strata in the basin axis, although growth strata can
indeed be observed closer to the Montsec
(Nijman, 1998).
The rates of accumulation, distance to the main structures, and
characteristic compensation timescale together suggest that hyperthermal
events of ca. 40 kyr duration can be recorded in the Castissent Fm. These
results confirm that, despite its highly dynamic nature, fluvial
sedimentation may contain valuable record of high-frequency events, even in
active tectonic contexts.
Conclusions
A new high-resolution isotopic record from the palaeosol-rich deposits at the
Chiriveta section identified a prominent negative carbon isotope excursion
(CIE) in continental settings. We suggest that the CIE recorded in fluvial
succession of the early Eocene Castissent Formation is the “U” event,
identified for the first time in continental deposits. This climatic event
reaches δ13C values of 2σ (standard deviation) below the
mean and is heralded and followed by several smaller-scale stepped CIEs,
which are interpreted as moments of enhanced primary productivity, leading
to an overall higher soil respiration. We show that all these CIEs are
relatively enriched in immobile elements (i.e. Ti, Zr, and Al) and display
an increase in MAP estimates. These observations coupled with the presence
of iron oxide nodules on an overall weathered succession, suggest an
increase in precipitation rates during these events. The data presented in
this study suggest a period of ca. 150 kyr of contrasted climate alternating
average and above background weathering conditions. Finally, the results of
this demonstrate the importance of hyperthermal events in continental
successions as well as in the preservation potential of such deposits.
Data availability
Isotopic and major and trace element data and tie points used for correlation can be
found in the Supplement (Tables S3 and S4).
The supplement related to this article is available online at: https://doi.org/10.5194/cp-16-227-2020-supplement.
Author contributions
LH led the field work, sampling, sample preparation,
data interpretation, and writing. TA contributed to field work, sampling,
data interpretation, discussion, and writing. JES performed stable isotope
analysis, data interpretation, and writing. JKCR interpreted the data, and
writing. MPM and EC contributed to fieldwork, sampling, discussion, and
writing. CP, JC and AF supervised the fieldwork, discussions, and writing. EV
led discussions on the palaeosols. KK and MH performed the XRF analysis. SC
supervised the project and writing.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
The authors would like to acknowledge the lifetime work of
Josep Serra Kiel, whose research and scientific contributions in the
Pyrenees have been fundamental to this work and much beyond. This study has
benefited from scientific discussion and field work with Marc Perret, Andres Nowak, Charlotte Läuchli, Teodoro Hunger, Joshua Vernier, Kelly Thomson, Margo Odlum, Aymeric Le Cotonnec, and Thierry Maeder.
Financial support
This research has been supported by Equinor (grant no. 4503110279) and by an Augustin Lombard grant from the Société de Physique et d'Histoire Naturelle de Genève (SPHN).
Review statement
This paper was edited by Alberto Reyes and reviewed by three anonymous referees.
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