Changes in the hydrological cycle in tropical East Africa during the Paleocene–Eocene Thermal Maximum

Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

Contents lists available at SciVerse ScienceDirect

Palaeogeography, Palaeoclimatology, Palaeoecology journal homepage: www.elsevier.com/locate/palaeo

Changes in the hydrological cycle in tropical East Africa during the Paleocene–Eocene Thermal Maximum Luke Handley a, Aoife O'Halloran b, Paul N. Pearson c, Elizabeth Hawkins a, Christopher J. Nicholas b, Stefan Schouten d, Ian K. McMillan c, Richard D. Pancost a,⁎ a

Organic Geochemistry Unit, Bristol Biogeochemistry Research Centre, School of Chemistry, University of Bristol, Cantock's Close, Bristol BS8 1TS, UK Department of Geology, Trinity College, University of Dublin, Dublin 2, Ireland School of Earth, Ocean and Planetary Sciences, Cardiff University, Park Place, Cardiff CF10 3YE, UK d NIOZ Royal Netherlands Institute for Sea Research, Department of Marine Organic Biogeochemistry, PO Box 59, 1790 AB Den Burg, Texel, The Netherlands b c

a r t i c l e

i n f o

Article history: Received 6 May 2011 Received in revised form 3 February 2012 Accepted 5 February 2012 Available online 12 February 2012 Keywords: Biomarkers Aridity Hydrogen isotopes Precipitation Kaolinite Weathering

a b s t r a c t The Paleocene–Eocene Thermal Maximum (PETM), at ca. 55.8 Ma, is one of the most studied instances of past greenhouse gas-induced global warming. As such, it provides a rich opportunity to examine the impact of such global change on local climates. The effects of increased continental and sea surface temperatures on local precipitation and humidity during the PETM remain poorly constrained and studies reveal complex, regional differences; whilst some localities appear to experience a net increase in humidity, others exhibit the opposite. Crucially, there are few records of hydrological change from tropical regions. Recent onshore drilling expeditions in Tanzania have yielded expanded sedimentary sections, deposited in a marine environment, that span much of the Late Cretaceous and Paleogene and show exceptionally good preservation of both calcareous microfossils and organic matter. The PETM interval has previously been constrained by both biostratigraphy and carbon isotopic records and spans ca. 7 m of section. Lipid distributions, including various terrestrial, marine and bacterial biomarkers and their hydrogen isotopic compositions, as well as mineralogy, were used to examine East African vegetation and hydrological responses to the global change occurring at the PETM. Although total organic carbon contents decrease, the concentrations of both higher plant (n-alkanes, n-alkanoic acids) and soil bacterial (glycerol dialkyl glycerol tetraethers) biomarkers increase dramatically at the onset of the PETM negative carbon isotope excursion (CIE), suggesting an increased discharge of fluvial sedimentary organic matter. Similarly, mineralogical indicators of terrestrial input – including Ti/Al and Si/Al ratios, quartz contents and, notably, the proportion of kaolinite – also increase at the onset of the CIE. However, higher plant leaf wax n-alkanes (C27, C29 and C31) become more deuterium-enriched throughout the same interval, suggesting a more arid and/or hotter, rather than a more humid, environment. This evidence collectively suggests an East African early PETM climate characterised by overall hot and arid conditions punctuated by intense, perhaps seasonal, precipitation events. These data match observations from other locations at mid-latitudes, suggesting that the humid climate often suggested for the PETM was not globally widespread. © 2012 Elsevier B.V. All rights reserved.

1. Introduction In the Late Paleocene, temperatures increased abruptly by ca. 5 °C globally (e.g. Kennett and Stott, 1991; Bralower et al., 1995; Zachos et al., 2003), and as much as 8 °C locally at high- and mid-latitudes (e.g. Sluijs et al., 2006; Zachos et al., 2006; Weijers et al., 2007). This warming, the Paleocene–Eocene Thermal Maximum (PETM), is thought to have been caused, at least in part, by the catastrophic release of biogenic methane from methane clathrate reservoirs in shallow marine sediments (Dickens et al., 1995, 1997), although other carbon sources (Kent et al., 2003; Kurtz et al., 2003; Svensen et al.,

⁎ Corresponding author. Tel.: + 44 117 928 9178. E-mail address: [email protected] (R.D. Pancost). 0031-0182/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2012.02.002

2004) could have also contributed to the increase in atmospheric CO2 concentrations. Indeed, it has recently been proposed that an increase in pCO2 alone is not sufficient to account for the full extent of such warming, suggesting that mechanisms other than greenhouse gas forcing may also have contributed to the global temperature increase (Zeebe et al., 2009). The PETM is not the only Early Eocene hyperthermal: the Eocene Thermal Maximum 2 (ETM2) or “Elmo” event (ca. 53 Ma) (Lourens et al., 2005) and the informally named “X” event (ca. 52 Ma) (Röhl et al., 2005) are also characterised by increases in sea surface temperature (SST). However, the PETM has the largest degree of warming sustained over the longest time interval. As such, it could provide a useful past model for understanding current climate change, which is driven by increased greenhouse gas emissions and for assessing the consequences of a rapid increase in atmospheric and seawater CO2 concentrations.

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

The increase in temperature at the PETM could have caused fundamental changes to global climate, including perturbations to the hydrological cycle, and the PETM has been thought to be associated with regional (Pagani et al., 2006), potentially global (Bowen et al., 2004), increases in humidity. To date, examinations of changes in clay mineral composition in nearshore settings have formed the greater part of PETM hydrological studies. Large increases in the relative abundance of kaolinite in the clay mineral suite have been recorded during the PETM from several locations: the South Atlantic off Antarctica (Robert and Chamley, 1991), the northeastern Atlantic (Robert and Chamley, 1991), the North Atlantic (Gibson et al., 1993, 2000; Zachos et al., 2006), New Zealand (Kaiho et al., 1996), and the Tethys in northern Spain (Schmitz and Pujalte, 2003, 2007). Such increases in kaolinite content have been linked to increases in humidity (Robert and Kennett, 1994; Gibson et al., 2000; Bolle and Adatte, 2001) on the adjacent continental landmasses. Alternatively, such changes could be due to increased physical weathering that mobilised deeper sediment originally deposited in a more humid paleogeographic setting (Thiry and Dupuis, 2000; Schmitz et al., 2001; Schmitz and Pujalte, 2003, 2007), and it has been argued that the latter hypothesis is more probable given the rapidity of the kaolinite increase relative to the onset of the PETM (b10 kyr) vs. the long-term changes in soil processes that would be required to account for such an increase in production (Thiry and Dupuis, 2000; Schmitz and Pujalte, 2003). Further evidence for a change in the global hydrological cycle, and specifically increased terrestrial sediment discharge, is provided by an increase in the proportion of terrestrial palynomorphs in coastal marine sediments during the PETM (Crouch et al., 2003), increases in coastal productivity as indicated by changes in phytoplankton (Crouch et al., 2003; Crouch and Brinkhuis, 2005) and nannofossil assemblages (Gibbs et al., 2006) and increases in sedimentation rates (e.g. Hollis et al., 2005; Sluijs et al., 2008b). In contrast, mineralogical studies from northern Spain suggest a transition from a semiarid to arid climate before the PETM to seasonally wetter but still generally dry conditions during the event (Schmitz and Pujalte, 2003). Those authors argued that the observed increase in kaolinite in the clay fraction and the inferred high erosion rates provided evidence for enhanced seasonal precipitation over a dry hinterland associated with an increase in the intra-annual humidity gradient. This hypothesis was supported by further studies that revealed the existence of a megafan at the Paleocene/Eocene boundary in northern Spain, the western Tethys, that would require repeated seasonal flash floods (Schmitz and Pujalte, 2007). This theory of a climate with increased intra-annual humidity and seasonal precipitation variation is consistent with model predictions for a contemporaneous strengthened greenhouse climate (Houghton et al., 2001). Likewise, inertinite (charcoal) distributions in the Cobham laminated lignite in southern England suggest episodic fires and post-fire erosion events during the PETM, indicating a persistent fire regime comprising regular drier episodes followed by periods of increased rainfall and runoff (Collinson et al., 2007). Leaf area analyses also suggest that mean annual precipitation declined in the Bighorn Basin, Wyoming, United States, near the onset of the PETM, before recovering in the latter stages of the warming period (Wing et al., 2005). The δD values of higher plant leaf wax components, in particular nalkanes, also provide a means of evaluating perturbations to the hydrological cycle and changes in humidity (e.g. Sauer et al., 2001; Schefuß et al., 2005; Smith and Freeman, 2006). Hydrogen isotope records have been used to study humidity changes during the PETM, with shifts to more deuterium-enriched values being invoked as evidence for increased water transport to the Arctic (Pagani et al., 2006) and transitional dry and wet intervals in central North America (Smith et al., 2007) and New Zealand (Handley et al., 2011). However, global geographical coverage remains very poor for the PETM with no low latitude data available. A wider geographical coverage of paleo-precipitation and humidity records is essential for understanding perturbations to the hydrological cycle during the PETM. The recently recovered Tanzania Drilling Project (TDP) sections, including sediments spanning the

11

PETM (Handley et al., 2008), offer a unique opportunity to study sub-tropical changes in PETM hydrology. We use organic biomarkers, including higher plant n-alkane concentrations and the BIT (Branched vs. Isoprenoid Tetraether) index, and mineralogical analyses to assess changes in the amount and nature of discharged sediment and use compound-specific δD values to reconstruct humidity changes. Specifically, we explore whether the hydrological cycle responded similarly at this location in the tropics as at mid- and high-latitudes or if a more arid climate state arose due to increased export of water to higher latitudes. 2. Material and methods 2.1. TDP Site 14 TDP Site 14 (9°16′59.89″S, 39°30′45.04″E; Fig. 1A) was first drilled in 2004, and TDP Hole 14A extends to 35.20 m total depth, with a sediment recovery of ca. 75%. A range of micropalaeontological evidence indicates that the sediments, which are mainly claystones and clayey siltstones, were deposited in a bathyal outer shelf or upper slope setting at an estimated depth of 300–500 m (Nicholas et al., 2006), although considerable uncertainty exists regarding the deeper end of this range. Further lithological information can be found in Nicholas et al. (2006). The Paleocene/Eocene transition at TDP Site 14 has previously been identified from biostratigraphy (Nicholas et al., 2006), and the PETM interval defined by the characteristic negative carbon isotope excursion recorded in the δ13C values of both higher plant derived n-alkanes and planktonic foraminifer carbonate (Handley et al., 2008). These results place the recovered PETM interval between ca. 12.6 m and 19.2 m depth, lying within the Kivinje Formation of the Kilwa Group, and the base of the fully expressed carbon isotope excursion (CIE) at 19.2 m is used as a reference for the onset of the PETM in the following discussion (Fig. 1B), although more detailed studies may reveal features associated with the ‘onset’ of the full excursion below this level. There is an inferred hiatus near the top of the section, above 12.6 m, which truncates the CIE (Handley et al., 2008). Preceding this hiatus, the carbon isotopic compositions of individual leaf wax n-alkanes remain relatively constant, with δ13C values being 5.5 (C31) to 7 (C29) ‰ more negative relative to preexcursion values (Handley et al., 2008). The hiatus corresponds to an abrupt return to pre-excursion values (Fig. 1B). The majority of PETM sections are characterised by a gradual recovery of the carbon isotope excursion (e.g. Kennett and Stott, 1991; Koch et al., 1995; Zachos et al., 2005, 2006; Pagani et al., 2006), and it is probable that TDP Site 14 contains only the early stages of the CIE and, therefore, the PETM. Unfortunately, a direct result of this depositional hiatus is that it is not yet possible to calculate accurate sedimentation rates, although the recovered portion is expanded (ca. 7 m) relative to deep sea records, and thus still constitutes an exceptional tropical record. Assuming a duration of 170 kyr for the PETM (Röhl et al., 2007), sedimentation rates must have been at least ca. 4 cm/ kyr. If our record includes only the core part of the PETM as represented by the clay layer in marine settings (100 kyr, Röhl et al., 2007), then sedimentation rates were at least 7 cm/kyr. As there is no evidence for a return of δ13C values toward pre-excursion values, the latter estimate seems more valid and sedimentation rates were probably even higher than this. Care must be taken when contrasting and comparing this record with other PETM sections, bearing in mind that these records probably represent only a part of the CIE. 2.2. Bulk rock analyses 2.2.1. Major elements Samples were fused with lithium metaborate and lithium tetraborate at 1000 °C in graphite crucibles. The melt was dissolved in dilute HNO3 and analysed by inductively coupled plasma-optical emission spectrometry (ICP-OES) for major elements, using an internal standard for quantification. Samples were analysed in batches of fifty,

12

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

Fig. 1. Geographical location of Paleocene/Eocene boundary TDP Site 14 in southern coastal Tanzania with biostratigraphy and higher plant derived n-alkane δ13C records. A: (a) current documented extent of the Kilwa Group outcrop, the formation to which the site belongs; (b) location of TDP Site 14 and local towns and villages. B: Compound-specific δ13C values of terrestrial higher plant derived n-alkanes. Stratigraphy of the Paleocene/Eocene boundary is based on δ13C values and planktonic foraminifera and nannofossil zonations are indicated. The shaded area highlights the PETM CIE and the zigzag line represents the possible hiatus truncating the PETM interval. Data are from Handley et al. (2008)

consisting of 44 samples, plus 1 blank and 4 duplicates, at OMAC Laboratories Ltd., Co. Galway, Ireland. 2.2.2. Mineralogy For determination of bulk sediment mineralogy, samples were reduced to a fine powder and 5 g was loaded into a cavity mount slide for X-ray diffraction (XRD) analysis. The X-ray generator used was a Phillips PW1720, equipped with a Phillips PW1050/25 diffractometer and a Phillips PW3313/20 Cu k-alpha anode tube that was operated with standard conditions of 40 kV and 20 mA. A soller slit and a 1° divergence slit were used on the incident X-ray beam, and an anti-scatter slit followed by a 0.25° receiving slit were used on the diffracted beam, in front of the AMR (anisotropic magnetoresistance) detector. The detector controller was a Hilton Brooks. All measurements were taken from 2° to 40° (2θ) at a step size of 0.02°/s. Minerals were identified from the XRD results using the Traces software, as part of the Hilton Brooks package. We note that rigorous clay mineralogical identification entails three analyses under different conditions (e.g. Bout-Roumazeilles et al., 1999), such that identifications of e.g. kaolinite reported here are somewhat tentative.

by elution with DCM/iso-propanol (3:1 v/v; neutral fraction) and 2% (by volume) acetic acid in diethyl ether (acid fraction). The neutral fraction was further split using a column packed with (activated) alumina by elution with hexane (saturated hydrocarbon fraction), hexane/DCM (9:1 v/ v; aromatic fraction) and DCM/MeOH (1:2 v/v; polar fraction).

2.3.1. Bulk organic analyses Sediment samples were gently washed with methanol (MeOH) and powdered with a Retsch PM100 Ball Mill. Total organic carbon (TOC) and inorganic carbon were measured as a mass % of the total sediment using a Carlo Erba EA 1108 and a Coulomat 702 (Strohlein).

2.3.3. Gas chromatography and gas chromatography–mass spectrometry Individual compounds were identified and quantified relative to internal standards (5α-androstane, apolar fraction; hexadecan-2-ol, polar fraction; n-C19 alkane, acid fraction) using gas chromatography (GC) and gas chromatography–mass spectrometry (GC–MS). Prior to analysis, acid fractions were methylated using BF3/MeOH and polar and acid fractions were silylated with BSTFA (N,O-bis(trimethylsilyl)trifluoroacetamide). Typical error in absolute quantification was ±20%. GC analysis was performed on a CarloErba Gas Chromatograph equipped with a flame ionisation detector (FID) and fitted with a Chrompack fused silica capillary column (50 m×0.32 mm i.d.) coated with a CP Sil-5CB stationary phase (dimethylpolysiloxane equivalent, 0.12 μm film thickness). GC–MS analysis was performed on a ThermoQuest Finnigan Trace GC interfaced with a ThermoQuest Finnigan Trace MS operating with an electron ionisation source at 70 eV and scanning over m/z ranges of 50 to 850 Da. The GC was fitted with a fused silica capillary column (50 m×0.32 mm i.d.) coated with a ZB1 stationary phase (dimethylpolysiloxane equivalent, 0.12 μm film thickness). For both GC and GC–MS, 1 μl of sample was injected at 50 °C using an on-column injector. The temperature was increased to 130 °C with an initial ramp of 20 °C/min, then to 300 °C at 4 °C/min, followed by an isothermal for 20 min.

2.3.2. Lipid biomarker extraction and fractionation For biomarker analyses, powdered freeze-dried samples were extracted via Soxhlet apparatus for 24 h using dichloromethane (DCM)/ MeOH (2:1 v/v) as the organic solvent. The total lipid extracts were separated into two fractions on an aminopropyl solid phase extraction column

2.3.4. Gas chromatography–thermal conversion–isotope ratio mass spectrometry Compound-specific stable hydrogen isotope ratios were determined for urea-adducted n-alkanes by gas chromatography–thermal conversion–isotope ratio mass spectrometry (GC–TC–IRMS) using

2.3. Geochemical analyses

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

an Agilent 6890 gas chromatograph equipped with an Agilent split/ splitless injector (splitless mode; 300 °C; purge time = 2 min) coupled to a ThermoQuest Finnigan Delta Plus XL isotope ratio mass spectrometer via a thermal conversion reactor (300 × 0.5 mm i.d.; Al2O3; 1450 °C) and ThermoQuest Finnigan GC Combustion III interface. The GC column used was a fused silica capillary column (30 m × 0.25 mm i.d.) with a ZB1 stationary phase (dimethylpolysiloxane equivalent, 0.25 μm film thickness). An H3+ factor, to correct for H3 ions, was calculated daily and was consistently lower than 5 ppm.V − 1. The temperature programme used consisted of an initial isothermal of 1 min at 40 °C followed by a ramp of 10 °C/min to 300 °C and a final isothermal for 13 min. To condition the reactor, two early eluting compounds (n-pentadecane and ethyl capricate) were coinjected with every run. The analytical accuracy of the system was monitored using a standard suite of 15 n-alkanes (C16–C30; Mixture B, Arndt Schimmelmann, Biogeochemical Laboratories, Indiana University), which was analysed before and after each sample. The root-mean-square error (RMS) for hydrogen isotopic measurements of the standard mixture was calculated for each sample analysed to monitor analytical accuracy, which was typically b5‰. Analytical precision, as represented by 1 standard deviation, was generally b2‰. δD values, based on duplicate analyses and reported in standard per mil notation relative to Vienna Standard Mean Ocean Water (VSMOW), were calculated against a calibrated H2 gas. 2.3.5. Liquid chromatography–mass spectrometry The polar fractions, containing glycerol dialkyl glycerol tetraethers (GDGTs), were dissolved in hexane/iso-propanol (99:1, v/v) and passed through 0.45 μm PTFE filters. Fractions were analysed by high performance liquid chromatography/atmospheric pressure chemical ionisation–mass spectrometry (HPLC/APCI–MS) using an Agilent 1100 series/Hewlett Packard 1100 MSD SL. Normal phase separation was achieved on an Alltech Prevail Cyano column (150 mm × 2.1 mm; 3 μm i.d.) with a flow rate of 0.2 ml.min − 1. The initial solvent was hexane/iso-propanol 99:1 (v/v), eluted isocratically for 5 min, followed by a linear gradient to 1.8% iso-propanol over 45 min. Ion scanning was performed in Single Ion Monitoring (SIM) mode to increase sensitivity and reproducibility and [M + H] + (protonated molecular ion) GDGT peaks were integrated. The ratio of branched GDGTs to crenarchaeol in marine and lacustrine sediments is a function of terrestrial input, expressed as the Branched vs. Isoprenoid Tetraether (BIT) index (Hopmans et al., 2004). 3. Results 3.1. Bulk rock analyses TDP Site 14A, like all TDP cores recovered to date, is dominated by lithogenic sediments. Carbonate contents are low and variable, ranging from 2 to 20% with no systematic trend, and the remainder comprises lithogenic components. Only a partial characterisation of the lithogenic fraction was attempted in this study, and we note that detailed mineralogical analyses are essential to probe these signatures; here they are presented to complement the biomarker interpretation. Quartz abundances vary between ca. 5 and 10% in pre-CIE sediments and increase during the CIE, reaching peak abundance at 16 m depth (Fig. 2). Despite large variations throughout the interval, kaolinite abundance generally increases upcore into the inferred PETM interval. Illite and smectite, however, have variable abundances throughout the cores, with values becoming more variable but also slightly higher during the CIE interval (Fig. 2). Titanium (Ti) and silica (Si) concentrations are reported normalised to aluminium (Al). Pre-PETM values are low and invariant for both (Fig. 2). For Ti/Al ratios, deviation from background begins at the CIE onset; an initial decrease (one datum) is followed by an overall increase throughout the rest of the CIE. A decrease occurs again at

13

the top of the interval (Fig. 2). Si/Al ratios exhibit a gradual increase, peaking between ca. 16 and 14.5 m, followed by a gradual decrease to background values towards the top of the section, similar to changes recorded by quartz abundances (Fig. 2). 3.2. Elemental analysis Total organic carbon (TOC) content is less than 1% (w/w) throughout the entire section, consistent with TOC values in other Tanzanian Cenozoic sediments (van Dongen et al., 2006). However, there is variation within the section; from the bottom at 34.4 m up to 19.2 m, the onset of the CIE, TOC content increases from 0.33 to ca. 0.70% before decreasing and remaining low (b0.50%) up to 12.8 m, above which it once again increases (Fig. 3); hence, the PETM appears to be associated with a decrease in the TOC content of the sediments. 3.3. Biomarkers 3.3.1. Higher plant and algal biomarkers The general composition of the saturated hydrocarbon fraction was reported by Handley et al. (2008); here, we discuss data and trends in greater detail. The saturated hydrocarbon fraction is dominated by a homologous series of n-alkanes, with the long-chain (>C25), predominantly odd-carbon-number homologues being most abundant (Handley et al., 2008), a typical signature for terrestrial higher plant derived n-alkanes (Eglinton et al., 1962; Eglinton and Hamilton, 1967; Kolattukudy, 1976). Individual n-alkane concentrations vary significantly through the portion of the PETM recovered at TDP Site 14. Concentrations of short-chain n-alkanes (C16–C19), typically derived from bacteria and marine algae (e.g. Gelpi et al., 1970; Naraoka and Ishiwatari, 1999; Duan and Wang, 2002), are lower during the CIE, with the exception of two horizons (Fig. 4), than in underlying sediments. In contrast, odd-carbon-number, long-chain n-alkane (C29, C31, C33) concentrations increase through the recovered portion of the CIE, although there is an initial decrease at 18.65 m (Fig. 5). Following the onset of the CIE there is also a change in the distribution of the higher plant derived n-alkanes. Prior to the CIE, C27, C29 and C31 n-alkanes are all present in similar concentrations, but C27 concentrations decrease at the onset of the CIE relative to the other two homologues (Fig. 5). Consequently, the average chain length (ACL) of the C25–C33 n-alkanes increases and remains elevated with respect to background pre-event values for the greater part of the CIE (Fig. 5; Handley et al., 2008). The n-alkane carbon preference indices (CPI) (Bray and Evans, 1961) and odd-to-even predominance (OEP) (Scalan and Smith, 1970) also increase slightly in the early stages of the PETM (Fig. 5).The ratio of long- to shortchain n-alkanes ranges from 1.9 to 6.5 in the deeper part of the core, with an outlier at 31.16 m, but increases to 8.7 and then 13.5 at 17.25 m and 16.15 m, respectively (Fig. 6). Values generally remain elevated but highly variable throughout the recovered CIE. The polar fractions contain a homologous series of C12–C34 nalkanols with a bimodal distribution dominated by the C16 and C28 components and a predominance of the even-carbon numbered homologues. Most acid fractions are dominated by a C16 to C34 homologous series of n-alkanoic acids, exhibiting an even-over-odd carbon number preference and a dominance of the C28 or C30 components. Short-chain n-alkanoic acid concentrations (C16, C18) are relatively constant throughout the section. However, long-chain even-carbonnumber n-alkanoic acid concentrations increase dramatically at 17.25 m and remain high throughout the recovered portion of the CIE, although local minima occur therein (Fig. 5). Driven by this shift in concentrations, there is an increase in both the average chain length and the ratio of high to low molecular weight fatty acid homologues during the recovered PETM interval (Fig. 6). Pristane (Pr) and phytane (Ph) are typical products of the diagenetic and catagenic alteration of the phytyl side chain of chlorophyll

14

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

Fig. 2. Mineralogical composition and major element abundances through the onset and during the early stages of the PETM at TDP Site 14 in Tanzania shown alongside the lithological record for the section. IIlite, smectite and quartz abundances are expressed as a % of the total sediment composition (note that for clay components, this is distinct from previous PETM studies where abundances are reported relative to total clay). Ti and Si are normalised to Al content. The shaded area highlights the CIE, characterised by increased sedimentation rates, and the zigzag line represents the possible hiatus truncating the PETM interval.

a (e.g. Dean and Whitehead, 1961; Rowland, 1990; Rontani and Volkman, 2003), although other sources have been recognised, such as archaea (Chappe et al., 1982). Both biomarkers are present in all samples. Total (Pr + Ph) concentration follows a similar trend to that of short-chain n-alkanes (C16–C19); a marked decrease at the onset of the CIE is followed by generally low values (Fig. 4). A suite of C27 to C29 steranes of various isomeric configurations and trace to

minor amounts of steroids were detected in the saturated fraction and polar fraction, respectively, but only in low concentrations. 3.3.2. Bacterial biomarkers In the lower part of the section, prior to the CIE, a wide variety of C29 to C35 hopanoids are present, including the 17β(H),21β(H) isomers (the configuration in which precursor bacteriohopanepolyols

Fig. 3. n-alkane δ13C values defining the CIE at TDP Site 14, compared to total organic carbon (TOC) content. The shaded area highlights the CIE, characterised by increased sedimentation rates, and the zigzag line represents the possible hiatus truncating the PETM interval.

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

15

Fig. 4. Depth profiles of marine organic matter biomarkers, including (Pr + Ph), low molecular weight (LMW) n-alkane and hopanoid concentrations and hopanoid isomer ratios. The hopanoid isomer ratio is the ratio of all hopanoids with the biological 17β(H),21β(H) configuration to those with the 17β(H),21α(H) and 17α(H),21β(H) configurations. For hopanoid concentrations and ratios, the scale at the top of the plot is for hopanes and the one at the bottom for hopanoic acids. Concentrations are expressed as weight per gram of freeze-dried sediment. The shaded area highlights the CIE and the zigzag line represents the possible hiatus truncating the PETM interval.

are produced by bacteria) as well as the more thermally stable αβ and βα isomers of the C29 to C35 hopanes (including both the 22S and 22R stereoisomers). Also present in pre-CIE sediments are a suite of C29 to C31 hopenes, mostly hop-17(21)-enes but also C29 and C30 hop-13(18)-enes. Concentrations are relatively high; C31ββ hopanoids are as abundant as the n-alkanes in pre-PETM horizons. Assemblages and concentrations remain similar up to 18.65 m, the

onset of the CIE, above which hopenes are absent and concentrations of ββ and βα hopanes, and hopanoids in general, decrease dramatically (Fig. 4). Consequently, the thermally mature 17α(H),21β(H) isomers dominate (although their concentrations are not higher), and C31αβ hopane is the most abundant hopanoid in the majority of CIE sediments. Coinciding with the change in distributions, the C31αβ homohopane isomerisation ratio, the ratio of 22S to 22R

Fig. 5. Depth profiles of higher plant biomarker (high molecular weight (HMW) n-alkane and n-alkanoic acid) concentrations and distribution ratios. Concentrations are expressed as weight per gram of freeze-dried sediment. n-alkane OEP = (C25 + 6 × C27 + C29) / (4 × C26 + 4 × C28) and CPI = 2 × (C25 + C27 + C29 + C31) / (C24 + 2 × (C26 + C28 + C30) + C32). ACL is the average chain length for C25 to C33 n-alkanes. The shaded area highlights the CIE and the zigzag line represents the possible hiatus truncating the PETM interval.

16

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

Fig. 6. Ratios indicative of terrigenous organic matter input, including ratios of high to low molecular weight n-alkanes and n-alkanoic acids and the BIT index. The lack of BIT index data from most pre-PETM horizons is due to the very low concentrations of branched GDGTs. The grey shaded area highlights the CIE and the zigzag line represents the possible hiatus truncating the PETM interval. BIT indices from TDP Sites spanning the Late Paleocene to the Early Eocene are also shown. Data are from van Dongen et al. (2006) and Pearson et al. (2007). The brown shaded area highlights sediments with BIT indices ≥ 0.6.

isomer, expressed as 22S/(22R + 22S), increases from values b0.30 for the pre-CIE samples to ca. 0.50 at 18.65 m. The ratio of ββ to (αβ + βα) homohopanes also decreases at 18.65 m (Fig. 4), such that all data indicate a transition from abundant thermally immature hopanes prior to the PETM to very low concentrations of mature hopanes during the CIE. This is consistent with the depth profiles of C30 to C34 hopanoic acids; these hopanoids also occur in all samples, with the 17β(H),21β(H) isomer being predominant in most samples and the αβ and βα configurations being present in subordinate concentrations. Critically, total hopanoic acid concentrations decrease in the early stages of the CIE followed by a slight recovery further up the section (Fig. 4). The ratio of ββ to (αβ+ βα) hopanoic acids also decreases at the onset of the PETM and remains low (Fig. 4). It is unlikely that these different horizons, separated by only a few metres of sediment, would have experienced different thermal histories and instead the changes in hopane distributions likely reflect changes in organic matter inputs. Hopanoids derive from the bacteriohopanepolyols that constitute bacterial cell walls (e.g. Rohmer and Ourisson, 1976a, 1976b; Ourisson et al., 1979) and, thus, can derive from terrestrial and marine organisms. The similarities between hopanoid and short-chain nalkane concentration depth profiles suggest that, at TDP Site 14, hopanes and especially hopanoic acids and hopenes have a predominantly marine source. Glycerol dialkyl glycerol tetraethers (GDGTs) also occur at TDP Site 14. Although concentrations are low, non-isoprenoidal branched GDGTs, membrane lipids produced by terrestrial bacteria (Hopmans et al., 2004; Weijers et al., 2006), and crenarchaeol, synthesised primarily by pelagic crenarchaeota (Sinninghe Damsté et al., 2002), are present above the limit of quantification in some samples (Hopmans et al., 2004; Schouten et al., 2007). The BIT index measures the amount of terrestrially-derived branched tetraether lipids relative to marine-derived crenarchaeol. Prior to the CIE, at 23.62 m, BIT values are relatively low at 0.16; in other pre-CIE horizons, concentrations of branched GDGTs were too low for BIT indices to be determined. BIT indices increase dramatically to 0.56 at 19.21 m, the onset of the CIE, and values remain elevated throughout the rest of the section, with maximum values of 0.85 to 0.87 between 14.92 and 13.43 m (Fig. 6).

3.3.3. Higher plant δD record δD values were measured for the C27, C29 and C31 n-alkanes, all of inferred terrestrial leaf wax origin (Handley et al., 2008). All records show similar features: in the lower part of the section, from 34.4 to 19.2 m, δD values range from -150‰ to -160‰, with slightly higher values at the very bottom of the section and for the sediment at 20.8 m (Fig. 7). At 18.65 m, there is a shift to higher values by 20 to 40‰: ca. 20‰ for C27 and C29 and 40‰ for C31. Values generally remain elevated in PETM sediments relative to those directly preceding the event. However, there is a large degree of variability within the former, with shifts of up to 18‰ between adjacent sediment horizons (Fig. 7). These shifts occur simultaneously in all n-alkane records and are of similar magnitudes. 4. Discussion 4.1. Interpretation of the δD record 4.1.1. Factors controlling terrestrial n-alkane δD values Hydrogen isotope (δD) values of terrestrial higher plant leaf waxes record: 1) the δD values of meteoric water (Sauer et al., 2001; Chikaraishi and Naraoka, 2003; Sachse et al., 2004; Rao et al., 2009); 2) the extent of D-enrichment during soil evaporation and leaf evapotranspiration, which is dependent on humidity (e.g. Smith and Freeman, 2006); and 3) the biological fractionation that occurs during leaf wax synthesis. Leaf wax n-alkane δD values range from -160 to -115‰ at TDP Site 14. These values are less negative than those reported for mid-latitudes in the Bighorn Basin (- 195 to - 185‰; Smith et al., 2007) and the Arctic (- 230 to - 260‰; Pagani et al., 2006) across the Paleocene/Eocene boundary. This is expected, as TDP Site 14 represents a coastal region, closer to the precipitation source waters, whereas the Bighorn Basin is a continental interior site and the Arctic is a high latitude site: water vapour, formed in the tropics becomes more D-depleted over continental interiors and at high latitude as the heavier isotope is rained out. Thus, the relationships between the δD values recorded at these three sites, before and during the PETM, are consistent with those occurring between such regions in the modern day (Table 1); however, to interpret these differences fully requires consideration of those variables that affect the

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

17

Fig. 7. Terrestrial higher plant n-alkane δD values through the onset and during the recovered PETM at TDP Site 14 in Tanzania, with the record for C29 n-alkane magnified across the CIE interval and compared to total (C27 + C29 + C31) HMW n-alkane concentrations. The n-alkane δD values are from Handley et al. (2008). The error bars in the isotope records represent one standard deviation based on duplicate analyses. Concentrations are expressed as weight per gram of freeze-dried sediment. The blue shaded intervals highlight horizons in which an increase in δD values coincides with a decrease in HMW n-alkane concentration. The grey shaded area highlights the CIE and the zigzag line represents the possible hiatus truncating the PETM interval.

apparent hydrogen isotopic enrichment between n-alkanes and source water (εCn-GW). The value for εC29-GW in C3 plants has been estimated to be between ca. -130‰ and -120‰ using surface sediment n-alkanes across a transect of European lakes (Sachse et al., 2004), and to be -140‰ to -130‰ using surface soil n-alkanes across Eastern China (Rao et al., 2009). These values are consistent with other reported hydrogen isotopic relationships between surface soil n-alkanes and meteoric water (e.g. Chikaraishi and Naraoka, 2003; Bi et al., 2005; Sachse et al., 2006; Smith and Freeman, 2006). Recent studies have highlighted the effect that changing plant communities can have on the isotopic composition of leaf wax lipids (e.g. Chikaraishi and Naraoka, 2003; Hou et al., 2007, 2008; Liu and Huang, 2008). Most of these studies focus on differences between C3 and C4 plant εC29-GW values, but recent work has also revealed differences in εCn-GW values between C3 plants in a semidesert environment (Pedentchouk et al., 2008). The authors reported differences of ca. 50‰ between plants of different classes (angiosperm vs. gymnosperm) and of up to 20‰ in plants of the same class, suggesting that εC29-GW may change significantly between higher plant species. Although these biotic effects cannot be ignored, if εC29-GW is assumed to be similar in C3 plants during the Early Paleogene, ca. -130‰, precipitation would have had a mean δD value of ca. -30 to -10‰ in Tanzania (paleolatitude 18°S, paleolongitude 22°E). This is slightly more negative relative to the modern day annual average estimate of ca. - 11‰ (95% CI: ± 6‰; (Bowen and Revenaugh, 2003; Bowen, 2009), which is similar to VSMOW and consistent with a coastal setting. Of further consideration is the fact that the Early Paleogene was probably an ice-free world: extensive ice sheets at high latitudes are

composed of isotopically light water and therefore δDsw would have been more negative during the Paleogene. Specifically, the decrease in global δDsw is thought to be ca. 8× larger than that observed in δ 18Osw, and this has been observed in records spanning Pleistocene glacial cycles (Schrag et al., 2002). Thus, assuming global δ 18Osw for an ice-free world to be 1.5‰ lower than today (e.g. Shackleton and Kennett, 1975; Shackleton, 1977; Lear et al., 2000; Zachos et al., 2006), seawater δD values during the ice-free Paleogene are expected to be ca. 12‰ lower. This depletion would be reflected in precipitation and meteoric water and could therefore account for the lower δD values recorded at TDP Site 14. Unlike in Tanzania, n-alkane hydrogen isotope compositions from the Bighorn Basin and the Arctic are D-enriched compared to modern values (Table 1). However, in those settings, D-enrichment could reflect less water vapour rainout prior to reaching the site. In fact, this is how Pagani et al. (2006) interpreted their Arctic record, arguing that less rainout during transport was consistent with increased water vapour transport to the Arctic and the observed decrease in water salinity (Sluijs et al., 2006; Sluijs et al., 2008b). They further hypothesised that this must reflect less net precipitation in parts of the low to mid-latitudes. 4.1.2. The increase in δD values during the CIE During the onset of the CIE, there is a positive hydrogen isotope excursion expressed in the δD records of all three higher plant derived n-alkanes. This positive shift could, partly, be due to an increase in the temperature of the source waters (e.g. Gat, 1996), with less fractionation during water vapour formation accounting for some,

Table 1 Average δD values for C29 n-alkane, prior to, during, and after the PETM in Tanzania compared to values from the Bighorn Basin (from Smith et al., 2007) and the Arctic (from Pagani et al., 2006) and inferred δD of precipitation at all three localities for the Paleogene, assuming a εC29-GW value of - 130‰, compared to present day values (Bowen and Revenaugh, 2003; Bowen, 2009). Average C29 n-alkane δD values

δD value of precipitation

Location

Paleolongitude

Paleolatitude

Pre-CIE

Early CIE

Post-CIE

Inferred for the Paleogene

Present day

Pande Peninsula Tanzania Bighorn Basin N. America Lomonosov Ridge Arctic

22E 107W 136E

18S 44N 87N

− − 196 − 210

− 140 − 183 − 165

− 145 − 191 − 210

− 25 to − 10 − 66 to − 53 − 80 to − 35

− 11 ± 6 − 83 ± 9 − 180 ± 34

18

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

but likely not all, of the observed shift (depending on the change in tropical Indian Ocean SSTs, ca. 5‰ of the 15 to 30‰ shift). It could also be due to a change in precipitation source regions, but our estimates for pre-CIE precipitation δD values are already close to VSMOW, so such changes should not have such a pronounced effect. Likewise, it seems unlikely that the increase in δD values at 18.65 m reflects a decrease in the already minimal rainout effect. Instead, it could reflect either a change in vegetation or the extent of evapotranspiration. As discussed above, εC29-GW values can vary amongst plant species, such that a change in vegetation could bring about changes in n-alkane δD values. Pollen records for TDP Site 14 do not reveal any major changes in higher plant assemblages, with no evidence for the presence of gymnosperms (Brinkhuis, personal communication). Although δD values of all n-alkanes exhibit a strong positive shift at the PETM, it is twice as large for the C31 than for the C29 and C27 n-alkanes (Section 3.3.3). This could indicate that there are different sources for the different high molecular weight (HMW) n-alkanes and that they do have different δD values. Consistent with this, n-alkane δD values are rather variable during the PETM as are ACLs, the latter suggesting a possible shift in the higher plant community structure as leaf wax distributions can vary between taxa (Collister et al., 1994; Chikaraishi and Naraoka, 2003; Bi et al., 2005; Sachse et al., 2006). However, shifts in δD values are not always associated with significant changes in HMW n-alkane ACL, and the correlation between the two is poor (R2 = 0.04 for C29 n-alkane). Furthermore, n-alkane δ 13C values remain relatively unchanged and do not fluctuate with δD values across the CIE interval, providing further circumstantial evidence against significant shifts in higher plant community (Handley et al., 2008). Thus, changes in higher plant community structure cannot explain all of the hydrogen isotopic variability. Alternatively, increased δD values could reflect an increase in soil evaporation or evapotranspiration during the early PETM relative to pre-CIE times. This could reflect elevated temperatures or a more arid climate in this East African setting; in any case, the record is clearly inconsistent with a more humid climate. The variability in PETM δD values could, therefore, suggest a highly variable climate state with alternating wet and dry intervals. Further evidence comes from the highly variable leaf wax concentrations and distributions, as discussed below. Thus, horizons characterised by both an increase in δD values and low concentrations of HMW n-alkanes (Fig. 7) could record a combination of arid conditions, and sparse vegetation. Although situated at a different latitude and setting (subtropical and marine vs. mid and continental), alternating wet/dry cycles within the PETM have also been reported by Kraus and Riggins (2007) in the Bighorn Basin. Nonetheless, the overall higher δD values suggest an overall drier climate during the early phase of the PETM. This has also been inferred for the Tethys in Northern Spain (Schmitz and Pujalte, 2003, 2007). It is also consistent with PETM n-alkane δD records from the Bighorn Basin (Smith et al., 2007) and New Zealand (Handley et al., 2011), which show a positive excursion in the initial stages of the PETM, followed by a return to more negative values in the latter stages of the CIE. Those authors interpreted these records as a shift to first drier then wetter conditions, an interpretation consistent, in the case of the Bighorn Basin, with records of plant physiognomy and floral composition (Wing et al., 2005) and the chemical composition and morphology of paleosols (Kraus and Riggins, 2007) from the same location. Because the latter stages of the PETM are probably missing from the TDP Site 14 record, it is not possible to determine whether a similar subsequent increase in humidity occurred in eastern Africa. Nonetheless, these data suggest that tropical Africa became more arid, at least episodically, during the early stages of the PETM. 4.2. Terrigenous inputs 4.2.1. Sedimentological records The base of the Tanzanian CIE is associated with an increase in kaolinite and quartz contents, as well as changes in various elemental

ratios. The increase in specific detrital components could be related to carbonate dilution during the PETM, however, Tanzanian bulk carbonate contents are rather low and variations are not apparently related to the trends shown in Fig. 2. Starting near the base of the Tanzanian CIE, kaolinite contents increase as both a percentage of sediment and as a percentage of total clay. Increased kaolinite contents can be an indicator of enhanced terrestrial input and/or increased chemical weathering, and therefore humidity level (necessary for kaolinite production). Alternatively, increased abundances coinciding with the CIE could reflect increased erosive energy in the regional hydrological cycle, stripping away more kaolinite beds (Thiry and Dupuis, 2000; Schmitz et al., 2001). The coeval increase in other detrital products, including quartz content and detrital element ratios may support increased erosional power, manifested as a quantitative increase in detrital flux. Alternatively, it could reflect a switch to a less mature detrital flux (e.g. coarser grains and more quartz), but the dilution of TOC contents (and marine biomarkers) suggests the former explanation. 4.2.2. Terrestrial organic matter records High and low molecular weight n-alkyl compounds are likely derived from higher plants and marine algae/bacteria, respectively (e.g. Eglinton et al., 1962; Eglinton and Hamilton, 1967; Gelpi et al., 1970; Naraoka and Ishiwatari, 1999; Duan and Wang, 2002). Therefore, the relative concentration and distribution of these compounds in sediments can be used to estimate changes in overall terrestrial contributions to the buried organic matter. Throughout the section, the distribution of n-alkyl compounds reflects a predominantly terrestrial source; however, this signal is more pronounced in sediments within the CIE. The first two horizons that record a significant shift to lower δ 13C values (18.65 and 17.90 m), i.e. the first two from within the PETM CIE, do not show any marked increase in terrestrial higher plant biomarkers. However, at 17.25 m, there is a significant change: the longto short-chain ratios for both n-alkanes and n-alkanoic acids increase as does the average chain length of both compound classes. These changes are driven by increases in the concentration of the high molecular weight homologues, in particular C29 and C31 for the n-alkanes and C28 and C30 for the acids. This increase in terrestrial higher plant biomarkers supports an increase in terrestrial organic matter input to the sediment. However, the sedimentary input of higher plant leaf wax derived n-alkanes is highly variable throughout the recovered CIE: n-alkane ratios shift significantly throughout, mostly driven by increases and decreases in HMW n-alkane concentrations, suggesting an episodic nature for these increases in terrigenous input. The increase in the BIT index during the PETM provides evidence for increased soil organic matter input. Interestingly, the increase in BIT indices occurs at 19.21 m, before the changes seen in the higher plant biomarkers but coinciding with the first lower δ 13C values. An increase from 0.16 to 0.56 is a substantial shift for BIT indices, reflecting a shift from predominantly marine organic matter to a significant soil organic matter contribution (Hopmans et al., 2004; Kim et al., 2006; Weijers et al., 2006). For example, an increase in BIT indices from 0.3 to 0.5 in the Congo fan following the termination of the Younger Dryas was interpreted by Weijers et al. (2009) as a substantial increase in terrestrial organic matter input and correlates with evidence for increased humidity and river discharge (Schefuß et al., 2005). Unfortunately, data from only one sample prior to the PETM is reported here from TDP Site 14. This is due to the fact that, in the bottom part of the section, branched GDGT and crenarchaeol concentrations are so low that accurate and reproducible integration and quantification was not possible using LC–MS, even in SIM mode. The virtual absence of soil-derived GDGTs prior to the PETM further supports the hypothesis that the higher BIT indices during the CIE are primarily due to a dramatic increase in soil organic matter input rather than a decrease in marine organic matter.

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

BIT indices have previously been reported for other TDP Sites (van Dongen et al., 2006; Pearson et al., 2007): for the greater part of the Eocene, indices range from ca. 0.2 to 0.5 (Fig. 6) and are much lower than those recorded at TDP Site 14 during the PETM (ca. 0.6 to 0.8; Fig. 6). Indeed, outside of the PETM interval, BIT indices >0.6 are only recorded in TDP sediments of Early Oligocene (32.3– 31.7 Ma) age (Pearson et al., 2004), and such high values likely record the dramatic sea level decrease associated with the inception or expansion of the East Antarctic ice sheet (e.g. Kennett and Shackleton, 1976; Lear et al., 2000; Zachos et al., 2008). Thus, the magnitude of the increase at the PETM is comparable to that which occurs as a result of a ca. 100 m drop in sea level, further highlighting the substantial increase in terrestrial discharge that occurs at the onset of the CIE. 4.2.3. Causes of increased terrigenous inputs Although an increase in terrestrial organic matter input could be interpreted as a local drop in sea level, we believe this not to be the case at TDP Site 14. Ostracod records from the Tethys, the Gebel Duwi section in Egypt, suggest a sea-level rise during the PETM (Speijer and Morsi, 2002). There is also evidence for a marine transgression caused by eustatic rise driven, at least in part, by the thermal expansion of water from a variety of locations (Sluijs et al., 2008a; Handley et al., 2011). Interestingly, Sluijs et al. (2008a) interpreted decreases in the BIT index to indicate a rise in sea level in sections from New Jersey, the North Sea and New Zealand, with BIT indices decreasing due to the transition to a deeper marine environment. At TDP Site 14, the opposite is observed. Given the wide geographical spread of these other records, it seems likely that the opposite trend observed in Tanzania is not evidence for a sea level fall, but instead reflects an increased transport of terrestrial organic matter from the East African continent into the ocean despite a probable sea level rise. 4.3. Changes in sediment accumulation rates As discussed in Section 3.3.1, the lack of an age model for TDP Site 14 precludes the calculation of accurate mass accumulation rates (MARs). However, changes in marine biomarker concentrations could partially reflect dilution processes and provide some insights into changes in sedimentation rates. Throughout the PETM interval, the concentrations of a range of inferred algal biomarkers (LMW nalkanes, pristane and phytane) decrease markedly compared to preCIE horizons. These decreases in marine organic matter biomarker concentrations could reflect a decrease in primary productivity in the water column. However, this seems unlikely as greater terrestrial input could be associated with greater nutrient delivery to the site and enhanced coastal productivity as suggested, for example, for New Zealand PETM sediments by Crouch et al. (2003, 2005). A more likely explanation is that the decrease in marine biomarker concentrations reflects dilution of organic matter deposited at the sediment surface due to increased terrigenous clay inputs. There are, however, discrete horizons within the PETM interval that are characterised by high LMW n-alkane, pristane and phytane concentrations (Fig. 4); these could reflect periods of enhanced water column productivity driven by increases in fluvial nutrient delivery from the continent. The marked decrease in the hopane 22S/(22S + 22R) and ββ/ (αβ + βα) ratios at 18.65 m probably reflects a dramatic decrease in the concentration of hopanoids derived from sources contemporaneous with the sediments, perhaps caused by dilution due to increased sedimentation rates. The residual, thermally mature hopanes are likely derived from an allochthonous source, such as the erosion of outcropping kerogen on the continent, further supporting an increase in continental weathering and erosion. An increase in sedimentation rates and terrigenous clay inputs could also be the cause of the observed decrease in TOC content during the CIE. Increased sedimentation rates during the PETM have also been observed in the Arctic (Sluijs et al., 2008b) and inferred from continental slope sections

19

from disparate global locations, such as New Zealand (Crouch et al., 2003), the northwestern Atlantic (Lippert and Zachos, 2007) and the Tethys (Schmitz et al., 2001).

5. Implications At TDP Site 14, terrigenous inputs, including both vegetation (e.g. n-alkanes) and soils (e.g. branched GDGTs and kaolinite), are overall higher and more variable in recovered PETM sediments. Such an increase could be caused by the enhanced erosion of kaolinite beds in the terrestrial realm formed prior to this period. The increases in kaolinite input reported here are similar to those observed at other locations during the PETM (Robert and Kennett, 1994; Gibson et al., 2000; Bolle and Adatte, 2001). However, the positive excursion in the nalkane δD records is not consistent with an increase in humidity and even suggests a more arid climate, or at least an episodically arid climate, during the early stages of the PETM, although changes in higher plant community might also contribute. Therefore, although there is evidence for enhanced terrestrial organic matter delivery that likely reflects changes in the hydrological cycle, there is no evidence for an overall increase in humidity in Tanzania during the PETM. This is consistent with evidence for greater water transport to the Arctic at the expense of precipitation at lower latitudes (Pagani et al., 2006; Sluijs et al., 2008b). It is also consistent with an intensification of the hydrological cycle, resulting in enhanced tropical evaporation and water vapour transport to higher latitudes, as suggested by some global precipitation models (Manabe, 1997; Bice and Marotzke, 2002; Speelman et al., 2010). Despite the prevalence of apparently more arid conditions in tropical East Africa during the PETM, there is still an increase in terrestrial inputs, including soil biomarkers, land plant biomarkers, quartz and kaolinite, to the bathyal outer shelf environment. We interpret this as evidence for more intense, perhaps seasonal, episodic precipitation events. Given the timescales involved, these events could also have been pluriannual or millennial and orbitally driven. Such episodic delivery of terrigenous sediments is consistent with the highly variable concentrations of kaolinite, quartz and higher plant biomarkers. Thus, we suggest that East African climate during the PETM was generally semi-arid, with episodic extreme precipitation associated with increased weathering. This could lead to the destabilisation of soils, greater soil erosion and kaolinite deposits in marginal marine settings. Indeed, not only does higher plant input increase but also soil organic matter input, as recorded by the increase in BIT indices. Similar conditions have been inferred from the Tethys in Northern Spain (Schmitz and Pujalte, 2003, 2007) and the New Jersey Shelf (Zachos et al., 2006; Lippert and Zachos, 2007). Interestingly, Allan and Soden (2008) used present day satellite observations and model simulations to evaluate the impact of current increases in SST on tropical precipitation and demonstrated a direct link between warmer climates and extreme precipitation events, a situation analogous to that inferred here for the PETM. Following the global warming at its onset, the PETM lasts for ca. 170 kyr (Röhl et al., 2007), a period over which the released CO2 was gradually sequestered; permanent carbon sequestration possibly occurred through an increase in organic matter burial due to increased sedimentation rates (Sluijs et al., 2008b) and/or the enhanced weathering of silicate rocks (Zachos et al., 2005). Bowen and Zachos (2010) argued that the rate of recovery of the CIE is an order of magnitude greater than would be expected for silicate weathering alone and that the sequestration of organic carbon must also have played an important part. More intense physical weathering and erosion, such as observed here at TDP Site 14, could be linked to an increase in chemical weathering processes, thereby providing an important feedback mechanism for the removal of transient excess greenhouse gas concentrations. Furthermore, sediment flux appears to increase

20

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21

markedly during the CIE but the decrease in TOC content is moderate, which may also reflect an overall increase in organic matter burial. 6. Conclusions TDP Site 14, from eastern Tanzania, is the first recovered PETM subtropical marine margin section and provides insight into the response of the local hydrological cycle during the early stages of the CIE. The portion of the PETM recovered at TDP Site 14 is marked by a large but highly variable increase in higher plant and soil organic matter and inorganic (kaolinite and quartz) matter input to the sediment, which is likely due to more intense precipitation. However, the higher plant leaf wax δD values increase, suggesting that there was a decrease rather than increase in humidity, consistent with reported changes in latitudinal precipitation patterns with greater water transport to high latitudes. The combination of the above observations suggests that increased sediment discharge and soil erosion were likely associated with episodic and perhaps seasonal precipitation events. Together with other records, these results clearly document a dramatic change in the hydrological cycle during the PETM. However, these changes are manifested differently across the globe, highlighting the need for more records from geographically disparate locations. Acknowledgements We are grateful to all those who have been involved in the Tanzania Drilling Project past and present, and for invaluable logistical support from the Tanzania Petroleum Development Corporation (Dar-esSalaam). The Tanzania Commission for Science and Technology and District Commissioner for the Kilwa District provided permission for the research. We would like to thank the Natural Environment Research Council (NERC) for financially supporting the project (grant NE/B503225/1 to P.N. Pearson) and funding Luke Handley's PhD and Dr. Robert Berstan and Dr. Ian Bull of the Bristol Node of the NERC Life Sciences Mass Spectrometry Facility for technical assistance. RDP acknowledges the Royal Society Wolfson Research Merit Award. References Allan, R.P., Soden, B.J., 2008. Atmospheric warming and the amplification of precipitation extremes. Science 321, 1481–1484. Bi, X.H., Sheng, G.Y., Liu, X.H., Li, C., Fu, J.M., 2005. Molecular and carbon and hydrogen isotopic composition of n-alkanes in plant leaf waxes. Organic Geochemistry 36, 1405–1417. Bice, K.L., Marotzke, J., 2002. Could changing ocean circulation have destabilized methane hydrate at the Paleocene/Eocene boundary? Paleoceanography 17. doi:10.1029/ 2001PA000678. Bolle, M.P., Adatte, T., 2001. Palaeocene–early Eocene climatic evolution in the Tethyan realm: clay mineral evidence. Clay Minerals 36, 249–261. Bout-Roumazeilles, V., Cortijo, E., Labeyrie, L., Debrabant, P., 1999. Clay mineral evidence of nepheloid layer contributions to the Heinrich layers in the northwest Atlantic. Palaeogeography, Palaeoclimatology, Palaeoecology 146, 211–228. Bowen, G.J., 2009. The online isotopes in precipitation calculator, version 2.2. Bowen, G.J., Revenaugh, J., 2003. Interpolating the isotopic composition of modern meteoric precipitation. Water Resources Research 39. doi:10.1029/2003WR002086. Bowen, G.J., Zachos, J.C., 2010. Rapid carbon sequestration at the termination of the Palaeocene–Eocene Thermal Maximum. Nature Geoscience 3, 866–869. Bowen, G.J., Beerling, D.J., Koch, P.L., Zachos, J.C., Quattlebaum, T., 2004. A humid climate state during the Palaeocene/Eocene Thermal Maximum. Nature 432, 495–498. Bralower, T.J., Zachos, J.C., Thomas, E., Parrow, M., 1995. Late Paleocene to Eocene paleoceanography of the equatorial Pacific Ocean: stable isotopes recorded at Ocean Drilling Program Site 865, Allison Guyot. Paleoceanography 10, 841–865. Bray, E.E., Evans, E.D., 1961. Distributions of n-paraffins as a clue to recognition of source beds. Geochimica et Cosmochimica Acta 22, 2–15. Chappe, B., Albrecht, P., Michaelis, W., 1982. Polar lipids of archaebacteria in sediments and petroleums. Science 217, 65–66. Chikaraishi, Y., Naraoka, H., 2003. Compound-specific δD-δ13C analyses of n-alkanes extracted from terrestrial and aquatic plants. Phytochemistry 63, 361–371. Collinson, M.E., Steart, D.C., Scott, A.C., Glasspool, I.J., Hooker, J.J., 2007. Episodic fire, runoff and deposition at the Palaeocene–Eocene boundary. Journal of the Geological Society 164, 87–97. Collister, J.W., Rieley, G., Stern, B., Eglinton, G., Fry, B., 1994. Compound-specific δ13C analyses of leaf lipids from plants with differing carbon dioxide metabolisms. Organic Geochemistry 21, 619–627.

Crouch, E.M., Brinkhuis, H., 2005. Environmental change across the Paleocene–Eocene transition from eastern New Zealand: a marine palynological approach. Marine Micropalaeontology 56, 138–160. Crouch, E.M., Dickens, G.R., Brinkhuis, H., Aubry, M.P., Hollis, C.J., Rogers, K.M., Visscher, H., 2003. The Apectodinium acme and terrestrial discharge during the Paleocene– Eocene Thermal Maximum: new palynological, geochemical and calcareous nannoplankton observations at Tawanui, New Zealand. Palaeogeography, Palaeoclimatology, Palaeoecology 194, 387–403. Dean, R.A., Whitehead, E.V., 1961. The occurrence of phytane in petroleum. Tetrahedron Letters 768–770. Dickens, G.R., O'Neil, J.R., Rea, D.K., Owen, R.M., 1995. Dissociation of oceanic methane hydrate as a cause of the carbon isotope excursion at the end of the Paleocene. Paleoceanography 10, 965–972. Dickens, G.R., Paull, C.K., Wallace, P., 1997. Direct measurement of in situ methane quantities in a large gas-hydrate reservoir. Nature 385, 426–428. Duan, Y., Wang, Z., 2002. Evidence from carbon isotope measurements for biological origins of individual longchain n-alkanes in sediments from the Nansha Sea, China. Chinese Science Bulletin 47, 578–581. Eglinton, G., Hamilton, R.J., 1967. Leaf epicuticular waxes. Science 156, 1322–1335. Eglinton, G., Gonzalez, A.G., Hamilton, R.J., Raphael, R.A., 1962. Hydrocarbon constituents of the wax coatings of plant leaves — a taxanomic survey. Phytochemistry 1, 89–102. Gat, J.R., 1996. Oxygen and hydrogen isotopes in the hydrologic cycle. Annual Review of Earth and Planetary Sciences 24, 225–262. Gelpi, V., Schneider, H., Mann, J., Oro, J., 1970. Hydrocarbons of geochemical significance in microscopic algae. Phytochemistry 9, 603–612. Gibbs, S.J., Bralower, T.J., Bown, P.R., Zachos, J.C., Bybell, L.M., 2006. Shelf and open-ocean calcareous phytoplankton assemblages across the Paleocene–Eocene Thermal Maximum: implications for global productivity gradients. Geology 34, 233–236. Gibson, T.G., Bybell, L.M., Owens, J.P., 1993. Latest Paleocene lithologic and biotic events in neritic deposits of Southwestern New-Jersey. Paleoceanography 8, 495–514. Gibson, T.G., Bybell, L.M., Mason, D.B., 2000. Stratigraphic and climatic implications of clay mineral changes around the Paleocene/Eocene boundary of the northeastern US margin. Sedimentary Geology 134, 65–92. Handley, L., Pearson, P.N., McMillan, I.K., Pancost, R.D., 2008. Large terrestrial and marine carbon and hydrogen isotope excursions in a new Paleocene/Eocene boundary section from Tanzania. Earth and Planetary Science Letters 275, 17–25. Handley, L., Crouch, E.M., Pancost, R.D., 2011. A New Zealand record of sea level rise and environmental change during the Paleocene–Eocene Thermal Maximum. Palaeogeography, Palaeoclimatology, Palaeoecology. doi:10.1016/j.palaeo.2011.03.001. Hollis, C.J., Dickens, G.R., Field, B.D., Jones, C.M., Strong, C.P., 2005. The Paleocene–Eocene transition at Mead Stream, New Zealand: a southern Pacific record of early Cenozoic global change. Palaeogeography, Palaeoclimatology, Palaeoecology 215, 313–343. Hopmans, E.C., Weijers, J.W.H., Schefuß, E., Herfort, L., Sinninghe Damsté, J.S., Schouten, S., 2004. A novel proxy for terrestrial organic matter in sediments based on branched and isoprenoid tetraether lipids. Earth and Planetary Science Letters 224, 107–116. Hou, J., D'Andrea, W.J., MacDonald, D., Huang, Y., 2007. Hydrogen isotopic variability in leaf waxes among terrestrial and aquatic plants around Blood Pond, Massachusetts (USA). Organic Geochemistry 38, 977–984. Hou, J., D'Andrea, W.J., Huang, Y., 2008. Can sedimentary leaf waxes record D/H ratios of continental precipitation? Field, model, and experimental assessments. Geochimica et Cosmochimica Acta 72, 3503–3517. Houghton, J.T., Ding, Y., Griggs, D.J., Noguer, M., van der Linden, P.J., Dai, X., Maskell, K., Johnson, C.A., 2001. Climate change—2001: the scientific basis. Cambridge University Press, Cambridge. Kaiho, K., Arinobu, T., Ishiwatari, R., Morgans, H.E.G., Okada, H., Takeda, N., Tazaki, K., Zhou, G.P., Kajiwara, Y., Matsumoto, R., Hirai, A., Niitsuma, N., Wada, H., 1996. Latest Paleocene benthic foraminiferal extinction and environmental changes at Tawanui, New Zealand. Paleoceanography 11, 447–465. Kennett, J.P., Shackleton, N.J., 1976. Oxygen isotopic evidence for development of psychrosphere 38 Myr ago. Nature 260, 513–515. Kennett, J.P., Stott, L.D., 1991. Abrupt deep-sea warming, palaeoceanographic changes and benthic extinctions at the end of the Palaeocene. Nature 353, 225–229. Kent, D.V., Cramer, B.S., Lanci, L., Wang, D., Wright, J.D., Van der Voo, R., 2003. A case for a comet impact trigger for the Paleocene/Eocene Thermal Maximum and carbon isotope excursion. Earth and Planetary Science Letters 211, 13–26. Kim, J.H., Schouten, S., Buscail, R., Ludwig, W., Bonnin, J., Sinninghe Damsté, J.S., Bourrin, F., 2006. Origin and distribution of terrestrial organic matter in the NW Mediterranean (Gulf of Lions): exploring the newly developed BIT index. Geochemistry, Geophysics, Geosystems 7. doi:10.1029/2006GC001306. Koch, P.L., Zachos, J.C., Dettman, D.L., 1995. Stable isotope stratigraphy and paleoclimatology of the Paleogene Bighorn Basin (Wyoming, USA). Palaeogeography, Palaeoclimatology, Palaeoecology 115, 61–89. Kolattukudy, P.E., 1976. The chemistry and biochemistry of natural waxes. Elsevier, Amsterdam. Kraus, M.J., Riggins, S., 2007. Transient drying during the Paleocene–Eocene Thermal Maximum (PETM): analysis of paleosols in the Bighorn Basin, Wyoming. Palaeogeography, Palaeoclimatology, Palaeoecology 245, 444–461. Kurtz, A.C., Kump, L.R., Arthur, M.A., Zachos, J.C., Paytan, A., 2003. Early Cenozoic decoupling of the global carbon and sulfur cycles. Paleoceanography 18. doi:10.1029/ 2003PA000908. Lear, C.H., Elderfield, H., Wilson, P.A., 2000. Cenozoic deep-sea temperatures and global ice volumes from Mg/Ca in benthic foraminiferal calcite. Science 287, 269–272. Lippert, P.C., Zachos, J.C., 2007. A biogenic origin for anomalous fine-grained magnetic material at the Paleocene–Eocene boundary at Wilson Lake, New Jersey. Paleoceanography 22. doi:10.1029/2007PA001471.

L. Handley et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 329–330 (2012) 10–21 Liu, Z., Huang, Y., 2008. Hydrogen isotopic compositions of plant leaf lipids are unaffected by a twofold pCO2 change in growth chambers. Organic Geochemistry 39, 478–482. Lourens, L.J., Sluijs, A., Kroon, D., Zachos, J.C., Thomas, E., Röhl, U., Bowles, J., Raffi, I., 2005. Astronomical pacing of late Palaeocene to early Eocene global warming events. Nature 435, 1083–1087. Manabe, S., 1997. Early development in the study of greenhouse warming: the emergence of climate models. Ambio 26, 47–51. Naraoka, H., Ishiwatari, R., 1999. Carbon isotopic compositions of individual long-chain n-fatty acids and n-alkanes in sediments from river to open ocean: multiple origins for their occurrence. Geochemical Journal 33, 215–235. Nicholas, C.J., Pearson, P.N., Bown, P.R., Dunkley-Jones, T., Huber, B.T., Karega, A., Lees, J.A., McMillan, I.K., O'Halloran, A., Singano, J.M., Wade, B.S., 2006. Stratigraphy and sedimentology of the Upper Cretaceous to Paleogene Kilwa Group, southern coastal Tanzania. Journal of African Earth Sciences 45, 431–466. Ourisson, G., Albrecht, P., Rohmer, M., 1979. The hopanoids. Palaeochemistry and biochemistry of a group of natural products. Pure and Applied Chemistry 51, 709–729. Pagani, M., Pedentchouk, N., Huber, M., Sluijs, A., Schouten, S., Brinkhuis, H., Sinninghe Damsté, J.S., Dickens, G.R., Scientists, Expedition, 2006. Arctic hydrology during global warming at the Palaeocene/Eocene Thermal Maximum. Nature 442, 671–675. Pearson, P.N., Nicholas, C.J., Singano, J.M., Bown, P.R., Coxall, H.K., van Dongen, B.E., Huber, B.T., Karega, A., Lees, J.A., Msaky, E., Pancost, R.D., Pearson, M., Roberts, A.P., 2004. Paleogene and Cretaceous sediment cores from the Kilwa and Lindi areas of coastal Tanzania: Tanzania Drilling Project Sites 1–5. Journal of African Earth Sciences 39, 25–62. Pearson, P.N., van Dongen, B.E., Nicholas, C.J., Pancost, R.D., Schouten, S., Singano, J.M., Wade, B.S., 2007. Stable warm tropical climate through the Eocene Epoch. Geology 35, 211–214. Pedentchouk, N., Sumner, W., Tipple, B., Pagani, M., 2008. δ13C and δD composition of n-alkanes from modern angiosperms and conifers: an experimental set up in central Washington State, USA. Organic Geochemistry 39, 1066. Rao, Z., Zhu, Z., Jia, G., Henderson, A.C.G., Xue, Q., Wang, S., 2009. Compound specific δD values of long chain n-alkanes derived from terrestrial higher plants are indicative of the δD of meteoric waters: evidence from surface soils in eastern China. Organic Geochemistry 40, 922–930. Robert, C., Chamley, H., 1991. Development of early Eocene warm climates, as inferred from clay mineral variations in oceanic sediments. Global and Planetary Change 89, 315–331. Robert, C., Kennett, J.P., 1994. Antarctic subtropical humid episode at the Paleocene– Eocene boundary: clay–mineral evidence. Geology 22, 211–214. Röhl, U., Westerhold, T., Monechi, S., Thomas, E., Zachos, J.C., Donner, B., 2005. The third and final early Eocene Thermal Maximum: characteristics, timing, and mechanisms of the “X” event. Geological Society of America Annual Meeting, 37, p. 264 (Salt Lake City, USA). Röhl, U., Westerhold, T., Bralower, T.J., Zachos, J.C., 2007. On the duration of the Paleocene–Eocene Thermal Maximum (PETM). Geochemistry, Geophysics, Geosystems 8. doi:10.1029/2007GC001784. Rohmer, M., Ourisson, G., 1976a. Bacteriophan derivatives — structural variations and distribution. Tetrahedron Letters 3637–3640. Rohmer, M., Ourisson, G., 1976b. Structure of bacteriohopanetetrols from acetobacterxylinum. Tetrahedron Letters 3633–3636. Rontani, J.F., Volkman, J.K., 2003. Phytol degradation products as biogeochemical tracers in aquatic environments. Organic Geochemistry 34, 1–35. Rowland, S.J., 1990. Production of acyclic isoprenoid hydrocarbons by laboratory maturation of methanogenic bacteria. Organic Geochemistry 15, 9–16. Sachse, D., Radke, J., Gleixner, G., 2004. Hydrogen isotope ratios of recent lacustrine sedimentary n-alkanes record modern climate variability. Geochimica et Cosmochimica Acta 68, 4877–4889. Sachse, D., Radke, J., Gleixner, G., 2006. δD values of individual n-alkanes from terrestrial plants along a climatic gradient — implications for the sedimentary biomarker record. Organic Geochemistry 37, 469–483. Sauer, P.E., Eglinton, T.I., Hayes, J.M., Schimmelmann, A., Sessions, A.L., 2001. Compoundspecific D/H ratios of lipid biomarkers from sediments as a proxy for environmental and climatic conditions. Geochimica et Cosmochimica Acta 65, 213–222. Scalan, R.S., Smith, J.E., 1970. An improved measure of the odd-to-even predominance in the normal alkanes of sediment extracts and petroleum. Geochimica et Cosmochimica Acta 34, 611–620. Schefuß, E., Schouten, S., Schneider, R.R., 2005. Climatic controls on central African hydrology during the past 20,000 years. Nature 437, 1003–1006. Schmitz, B., Pujalte, V., 2003. Sea-level, humidity, and land-erosion records across the initial Eocene Thermal Maximum from a continental-marine transect in northern Spain. Geology 31, 689–692. Schmitz, B., Pujalte, V., 2007. Abrupt increase in seasonal extreme precipitation at the Paleocene–Eocene boundary. Geology 35, 215–218. Schmitz, B., Pujalte, V., Núñez-Betelu, K., 2001. Climate and sea-level perturbations during the Incipient Eocene Thermal Maximum: evidence from siliciclastic units in the Basque Basin (Ermua, Zumaia and Trabakua Pass), northern Spain. Palaeogeography, Palaeoclimatology, Palaeoecology 165, 299–320. Schouten, S., Huguet, C., Hopmans, E.C., Kienhuis, M.V.M., Sinninghe Damsté, J.S., 2007. Analytical methodology for TEX86 paleothermometry by high-performance liquid

21

chromatography/atmospheric pressure chemical ionization–mass spectrometry. Analytical Chemistry 79, 2940–2944. Schrag, D.P., Adkins, J.F., McIntyre, K., Alexander, J.L., Hodell, D.A., Charles, C.D., McManus, J.F., 2002. The oxygen isotopic composition of seawater during the Last Glacial Maximum. Quaternary Science Reviews 21, 331–342. Shackleton, N.J., 1977. Oxygen isotope stratigraphic record of late Pleistocene. Philosophical Transactions of the Royal Society of London. Series B, Biological Sciences 280, 169–182. Shackleton, N.J., Kennett, J.P., 1975. Paleotemperature history of the Cenozoic and the initiation of Antarctic glaciation: oxygen and carbon isotope analyses in DSDP sites 277, 279, and 281. Initial reports of the Deep Sea Drilling Project, 29, pp. 743–755. Sinninghe Damsté, J.S., Schouten, S., Hopmans, E.C., van Duin, A.C.T., Geenevasen, J.A.J., 2002. Crenarchaeol: the characteristic core glycerol dibiphytanyl glycerol tetraether membrane lipid of cosmopolitan pelagic crenarchaeota. Journal of Lipid Research 43, 1641–1651. Sluijs, A., Schouten, S., Pagani, M., Woltering, M., Brinkhuis, H., Sinninghe Damsté, J.S., Dickens, G.R., Huber, M., Reichart, G.J., Stein, R., 2006. Subtropical Arctic Ocean temperatures during the Palaeocene/Eocene Thermal Maximum. Nature 441, 610–613. Sluijs, A., Brinkhuis, H., Crouch, E.M., John, C.M., Handley, L., Munsterman, D., Bohaty, S.M., Zachos, J.C., Reichart, G.J., Schouten, S., Pancost, R.D., Sinninghe Damsté, J.S., Welters, N.L.D., Lotter, A.F., Dickens, G.R., 2008a. Eustatic variations during the Paleocene–Eocene greenhouse world. Paleoceanography 23. doi:10.1029/2008PA001615. Sluijs, A., Röhl, U., Schouten, S., Brumsack, H.J., Sangiorgi, F., Sinninghe Damsté, J.S., Brinkhuis, H., 2008b. Arctic late Paleocene–early Eocene paleoenvironments with special emphasis on the Paleocene–Eocene Thermal Maximum (Lomonosov Ridge, Integrated Ocean Drilling Program Expedition 302). Paleoceanography 23. doi:10.1029/2007PA001495. Smith, F.A., Freeman, K.H., 2006. Influence of physiology and climate on δD of leaf wax n-alkanes from C3 and C4 grasses. Geochimica et Cosmochimica Acta 70, 1172–1187. Smith, F.A., Wing, S.L., Freeman, K.H., 2007. Magnitude of the carbon isotope excursion at the Paleocene–Eocene Thermal Maximum: the role of plant community change. Earth and Planetary Science Letters 262, 50–65. Speelman, E.N., Sewall, J.O., Noone, D., Huber, M., von der Heydt, A., Sinninghe Damsté, J.S., Reichart, G.J., 2010. Modeling the influence of a reduced equator-to-pole sea surface temperature gradient on the distribution of water isotopes in the Early/ Middle Eocene. Earth and Planetary Science Letters 298, 57–65. Speijer, R.P., Morsi, A.M.M., 2002. Ostracode turnover and sea-level changes associated with the Paleocene–Eocene Thermal Maximum. Geology 30, 23–26. Svensen, H., Planke, S., Malthe-Sorenssen, A., Jamtveit, B., Myklebust, R., Eidem, T.R., Rey, S.S., 2004. Release of methane from a volcanic basin as a mechanism for initial Eocene global warming. Nature 429, 542–545. Thiry, H., Dupuis, C., 2000. Use of clay minerals for paleoclimatic reconstructions: limits of the method with special reference to the Paleocene–lower Eocene interval. GFF 122, 166–167. van Dongen, B.E., Talbot, H.M., Schouten, S., Pearson, P.N., Pancost, R.D., 2006. Well preserved Palaeogene and Cretaceous biomarkers from the Kilwa area, Tanzania. Organic Geochemistry 37, 539–557. Weijers, J.W.H., Schouten, S., Spaargaren, O.C., Sinninghe Damsté, J.S., 2006. Occurrence and distribution of tetraether membrane lipids in soils: implications for the use of the TEX86 proxy and the BIT index. Organic Geochemistry 37, 1680–1693. Weijers, J.W.H., Schouten, S., Sluijs, A., Brinkhuis, H., Sinninghe Damsté, J.S., 2007. Warm arctic continents during the Palaeocene–Eocene Thermal Maximum. Earth and Planetary Science Letters 261, 230–238. Weijers, J.W.H., Schouten, S., Schefuß, E., Schneider, R.R., Sinninghe Damsté, J.S., 2009. Disentangling marine, soil and plant organic carbon contributions to continental margin sediments: a multi-proxy approach in a 20,000 year sediment record from the Congo deep-sea fan. Geochimica et Cosmochimica Acta 73, 119–132. Wing, S.L., Harrington, G.J., Smith, F.A., Bloch, J.I., Boyer, D.M., Freeman, K.H., 2005. Transient floral change and rapid global warming at the Paleocene–Eocene boundary. Science 310, 993–996. Zachos, J.C., Wara, M.W., Bohaty, S., Delaney, M.L., Petrizzo, M.R., Brill, A., Bralower, T.J., Premoli-Silva, I., 2003. Paleoclimate: a transient rise in tropical sea surface temperature during the Paleocene–Eocene Thermal Maximum. Science 302, 1551–1553. Zachos, J.C., Röhl, U., Schellenberg, S.A., Sluijs, A., Hodell, D.A., Kelly, D.C., Thomas, E., Nicolo, M., Raffi, I., Lourens, L.J., McCarren, H., Kroon, D., 2005. Rapid acidification of the ocean during the Paleocene–Eocene Thermal Maximum. Science 308, 1611–1614. Zachos, J.C., Schouten, S., Bohaty, S., Quattlebaum, T., Sluijs, A., Brinkhuis, H., Gibbs, S.J., Bralower, T.J., 2006. Extreme warming of mid-latitude coastal ocean during the Paleocene–Eocene Thermal Maximum: inferences from TEX86 and isotope data. Geology 34, 737–740. Zachos, J.C., Dickens, G.R., Zeebe, R.E., 2008. An early Cenozoic perspective on greenhouse warming and carbon-cycle dynamics. Nature 451, 279–283. Zeebe, R.E., Zachos, J.C., Dickens, G.R., 2009. Carbon dioxide forcing alone insufficient to explain Palaeocene–Eocene Thermal Maximum warming. Nature Geoscience 2, 576–580.