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Large Igneous Province Control on Ocean Anoxia and
Eutrophication in the North Sea at the Paleocene–Eocene
Thermal Maximum
Erica Mariani
1
, Sev Kender
1,2
, Stephen P. Hesselbo
1,3
, Kara Bogus
1
, Kate Littler
1,3
,
James B. Riding
2
, Melanie J. Leng
2
, Simon J. Kemp
2
, Karen Dybkjær
4
, Gunver K. Pedersen
4
,
Thomas Wagner
5
, and Alexander J. Dickson
6
1
Camborne School of Mines, Department of Earth and Environmental Sciences, University of Exeter, Cornwall, UK,
2
British Geological Survey, Nottingham, UK,
3
Environment and Sustainability Institute, University of Exeter, Cornwall,
UK,
4
Geological Survey of Denmark and Greenland (GEUS), Copenhagen, Denmark,
5
Lyell Centre, Heriot‐Watt
University, Edinburgh, UK,
6
Department of Earth Sciences, Royal Holloway, University of London, Surrey, UK
Abstract The Paleocene–Eocene Thermal Maximum (PETM) was a global hyperthermal event ∼56 Ma
characterized by massive input of carbon into the ocean–atmosphere system and global warming. A leading
hypothesis for its trigger is the emplacement of the North Atlantic Igneous Province (NAIP), with extensive
extrusion/intrusion of igneous material into nearby sedimentary basins, forcing local uplift and warming‐
inducing carbon emissions. It remains unclear if oceanographic changes in the North Sea–Norwegian Sea–
Arctic basins, such as anoxia and productivity, were causally linked to local NAIP uplift/activity, and at what
time scales these perturbations occurred. To test mechanisms and time scales, we present geochemical proxies
(XRF analysis, clay mineralogy, molybdenum isotopes, and pyrite framboid size distribution) in undisrupted
marine sediment core E8X located in the central North Sea. We find evidence for a rapid onset of anoxia/
euxinia at the negative carbon isotope excursion from redox proxies, followed by a gradual drawdown of
molybdenum/total organic carbon (Mo/TOC) during the PETM main phase indicative of tectonically‐restricted
basin likely from NAIP uplift. A short‐lived increase in Mo, pyrite and TOC occurred during a precursor event
associated with a sedimentary mercury pulse indicative of volcanic activity. We suggest thermal uplift and flood
basalt volcanism tectonically restricted the North Sea and tipped it into an euxinic state via volcanic emission–
oceanographic feedbacks inducing eutrophication. This fine temporal separation of tectonic versus climatic
geochemical proxies, combined with pulsed NAIP volcanism, demonstrates that Large Igneous Province
emplacements can, at least locally, result in ocean biogeochemical feedbacks operating on relatively short
timescales.
1. Introduction
The Paleocene‐Eocene Thermal Maximum (PETM) was a global warming event ∼56 Ma that led to anomalous
sea surface temperature increases of ∼5–9°C (Dunkley Jones et al., 2013). The PETM is associated with in-
jections of
13
C‐depleted carbon into the exogenic carbon pool, reflected in a 2–7‰ negative carbon isotope
excursion (CIE) in marine and terrestrial sediment successions. The relative age model of the event encompassing
the CIE onset and recovery phases is based on two proposed cyclostratigraphic solutions from Svalbard core BH9/
05, whose durations are estimated to be ∼140 ka (option A) and ∼200 ka (option B) (Charles et al., 2011). The
overall event and its rapid onset of ∼3–4 kyr (Kirtland Turner et al., 2017; Kender et al., 2021) can be used to test
climate models and to predict, within limitations, how climate might respond to anthropogenic activity. However,
the causal mechanisms triggering the PETM are still unclear, hampering our ability to model the consequences of
the event at the local scale.
The PETM affected marine environments by reducing water column oxygenation in many locations from high to
low northern and southern latitudes, the paleo‐Tethys, Europe, Asia, and America (Carmichael et al., 2017;
Papadomanolaki et al., 2022). The PETM onset is associated with the largest Mesozoic–Cenozoic benthic
foraminiferal extinction (Thomas, 1998). Multiple causal factors have been proposed, such as temperature in-
crease leading to high metabolic rates in starved waters from increased organic matter remineralization (Ma
et al., 2014), ocean acidification (Thomas, 2007), and oxygen depletion (Thomas, 1998). In NW Europe,
particularly around the North Sea, Denmark, Norway, and Arctic Ocean, laminated sediments and organic
RESEARCH ARTICLE
10.1029/2023PA004756
Key Points:
•Paleoceanographic reconstructions of
the North Sea during the Paleocene‐
Eocene Thermal Maximum via high
resolution geochemical data
•We find that redox, productivity and
basin restriction changes occurred
before and after the carbon isotope
excursion
•We argue that precursor and initial
paleoceanographic changes were
driven by Large Igneous Province
emplacement and regional uplift
Supporting Information:
Supporting Information may be found in
the online version of this article.
Correspondence to:
E. Mariani and S. Kender,
E.Mariani2@exeter.ac.uk;
s.kender@exeter.ac.uk
Citation:
Mariani, E., Kender, S., Hesselbo, S. P.,
Bogus, K., Littler, K., Riding, J. B., et al.
(2024). Large Igneous Province control on
ocean anoxia and eutrophication in the
North Sea at the Paleocene–Eocene
Thermal Maximum. Paleoceanography
and Paleoclimatology,39,
e2023PA004756. https://doi.org/10.1029/
2023PA004756
Received 30 AUG 2023
Accepted 3 APR 2024
Author Contributions:
Conceptualization: Erica Mariani,
Sev Kender
Investigation: Erica Mariani, Sev Kender,
Melanie J. Leng, Simon J. Kemp,
Thomas Wagner, Alexander J. Dickson
Resources: Karen Dybkjær, Gunver
K. Pedersen
Supervision: Sev Kender, James
B. Riding
Visualization: Erica Mariani, Sev Kender
Writing – original draft: Erica Mariani,
Sev Kender, Stephen P. Hesselbo,
Kara Bogus, Kate Littler, James B. Riding,
Melanie J. Leng, Simon J. Kemp,
© 2024. The Authors.
This is an open access article under the
terms of the Creative Commons
Attribution License, which permits use,
distribution and reproduction in any
medium, provided the original work is
properly cited.
MARIANI ET AL. 1 of 18
geochemical proxies have been recorded suggesting widespread anoxia during the PETM (Berndt et al., 2023;
Harding et al., 2011; Kender et al., 2012; Schoon et al., 2015; Sluijs et al., 2006). Although increased hydrological
cycling and primary production has been suggested as the main cause of oxygen depletion (Dickson, Rees‐Owen,
et al., 2014; Ma et al., 2014), areas proximal to the North Atlantic Igneous Province (NAIP), such as
the North Sea, have been subjected to strong local environmental changes related to the emplacement of the NAIP
(Kender et al., 2012). Among these changes, basin restriction of the North Sea via tectonic uplift is a
favored mechanism that may have contributed to oxygen depletion in the North Sea and surrounding areas
(Hartley et al., 2011).
Regional tectonic uplift was primarily controlled by the upward movement of the Icelandic mantle plume between
Greenland and Scotland during the opening of the Northeast Atlantic Ocean (Hartley et al., 2011; Jones, Hoggett,
et al., 2019). The correlation between tectonic uplift and NAIP activity is well documented, with thousands of
intruded sills recognized in 2D and 3D seismic stratigraphy in the Norwegian Sea and northwestern North Sea
(Berndt et al., 2023; Conway‐Jones & White, 2022; Hartley et al., 2011; Jones, Percival, et al., 2019; Svensen
et al., 2004). Increased volcanic activity around the PETM onset is also documented by sedimentary mercury (Hg)
and Hg isotopes in the North Sea, Spain, Svalbard, Denmark, and the Arctic Ocean (Jin et al., 2023; Jones,
Percival, et al., 2019; Kender et al., 2021; Tremblin et al., 2022). The North Sea and the northern Atlantic Ocean
were connected by a shallow seaway, however this connection was possibly cut off by the NAIP uplift (Hartley
et al., 2011; Zacke et al., 2009). NAIP emplacement not only modified the North Sea–North Atlantic connection,
but also reorganized and narrowed the Norwegian–Greenland seaway, which became a shallow water strait
around the PETM (Berndt et al., 2023; Hovikoski et al., 2021; Planke et al., 2023). Since the Norwegian–
Greenland seaway was further connected to the Arctic Ocean (Prøis, 2015), the shallowing of the seaway
would result in the partial isolation of the Arctic Ocean (Jones et al., 2023).
Previous records of bottom water oxygen depletion in the North Sea region are not sufficiently resolved to identify
whether basin restriction facilitated, or drove, the observed anoxia/euxinia around the PETM onset (e.g., Schoon
et al., 2015; Stokke et al., 2021). The exact phasing between the emplacement of the NAIP, North Sea basin
restriction, primary productivity and runoff changes and bottom water anoxia/euxinia has not yet been robustly
determined, mainly due to the lack of high‐resolution records from high quality core material. Previous records
are inconclusive due to incomplete core recovery and drilling disturbance at the CIE onset and the main PETM
phase (Arctic Ocean core 302‐M0004 A, Sluijs et al., 2006), or have been affected by transported organic carbon
from mixed sources and/or reworked terrestrial material around the CIE onset, which may alter the shape of the
CIE (North Sea core 22/10a‐4, Kender et al., 2012; Longyearbyen, Svalbard, Harding et al., 2011; Svalbard core
BH9/05, Cui et al., 2011). The timing of volcanic activity, basin restriction and oceanographic change is not only
critical for understanding the causal mechanisms of North Sea anoxia/euxinia, but also has large implications for
the tipping point of the PETM. Several lines of evidence argue that NAIP activity contributed to the trigger of the
PETM carbon release, including direct proxy‐based evidence for volcanism immediately preceding and during
the CIE onset from high resolution sedimentary Hg (Jones, Percival, et al., 2019,2023; Kender et al., 2021;
Tremblin et al., 2022) and Hg isotope records (Jin et al., 2023). Although the correlation between tectonic uplift
and volcanic activity is well documented (Champion et al., 2008; Conway‐Jones & White, 2022; Hartley
et al., 2011; Storey et al., 2007; Wilkinson et al., 2017), possible bottom water changes resulting from tectonically
driven basin restriction need further investigation.
In this study, we analyze the evolution of hydrography and bottom water oxygen conditions with a particular
focus to the PETM CIE onset, from an exceptionally well‐preserved marine sediment core previously described
from the Danish sector of the North Sea (wellsite E8X; Figure 1; Kender et al., 2021). E8X consists of un-
disturbed marine sediments encompassing the PETM onset, main phase and recovery intervals, and has a CIE
onset that is well defined and thought to represent approximately continuous sedimentation (Kender et al., 2021).
We generated geochemical data, including elemental ratios, clay mineralogy, Mo isotopes, and pyrite framboid
size distribution, to determine the succession of oxygen conditions, primary productivity and terrigenous input at
the transition to the PETM CIE. We use these records to investigate the lead/lag relationships between volcanic
activity, carbon release, euxinia, productivity and inferred basin restriction. We find that precursor PETM basin
restriction likely created the conditions for pulsed volcanism to drive initial reductions in bottom water
oxygenation, followed by climate‐driven productivity increases shortly after the PETM onset.
Karen Dybkjær, Gunver K. Pedersen,
Thomas Wagner, Alexander J. Dickson
Writing – review & editing:
Erica Mariani, Sev Kender, Stephen
P. Hesselbo, Kara Bogus, Kate Littler,
James B. Riding, Melanie J. Leng,
Karen Dybkjær, Gunver K. Pedersen,
Thomas Wagner, Alexander J. Dickson
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MARIANI ET AL. 2 of 18
2. Material and Methods
2.1. Depositional Setting and Stratigraphy of E−8X
Wellsite E8X (55°38’13.42’’N, 04°59’11.96’’E) is located in the central portion of the Danish North Sea
(Figure 1) and the core is described in Kender et al. (2021). The paleobathymetry of the deep central parts of the
northern North Sea (58–62°N), was reconstructed to have been around 600–800 m deep (Kjennerud & Gill-
more, 2003). Since Wellsite E8X is located only 3°S of the northern North Sea in the deep central portion of the
basin, we assume that paleodepths might have not been too dissimilar, with the central part of the North Sea being
>500m. The Paleocene‐Eocene boundary in core E8X was previously identified by a positive excursion on the
gamma‐ray log at the base of the Sele Formation, along with the acme of Apectodinium augustum, a dinoflagellate
cyst biomarker for the PETM, located in the lowermost laminated part of the Sele Formation (Schiøler et al., 2007;
Supplementary Material). The studied core interval consists of 7 m of marine sediments (2021.06–2027.36 m
composite core depth) spanning the uppermost Paleocene and lowermost Eocene. The CIE (2,025.41–
2,025.32 m) was identified by a 5‰ negative stable organic carbon isotope (δ
13
C
org
) excursion in the bulk
material (Kender et al., 2021), whose amplitude ‰ is consistent with the CIE expression at nearby wellsite 22/
10a‐4 (Kender et al., 2012).
Figure 1. Late Paleocene paleogeographic reconstruction (adapted from Jones et al., 2023). the light red, violet and gray areas
represent subaerial, submarine and subsurface volcanism and sill intrusions, respectively. Dark red points represent volcanic
centers. Shallow and deep marine areas are shown as light and dark blue, respectively. FSB: Faroe‐Shetland Basin; NGS:
Norwegian‐Greenland Seaway; NZ: Novaya Zemlya.
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The CIE displays a two‐step trend: the first step is characterized by a persistent 1.6‰ decline in δ
13
C
org
values
between 2,025.4 and 2,025.37 m; the second step is more rapid with an amplitude of around 3.4‰ between
2025.36 and 2025.32 m and reaches minimum δ
13
C
org
values of 30.7‰ (Kender et al., 2021). The “main phase”
of the PETM (2,025.32–2,023.15 m) is an interval assigned to the Sele Formation (Kender et al., 2021; Fig. S1 in
Supporting Information S1) with sustained negative δ
13
C
org
values of 31‰ to 29‰ at E8X. The main phase
is dominated by finely laminated dark mudstone, although a silty interval is observed from ∼2,024.5 to 2,023.2 m.
The recovery Phase 1 (2,023.15–2,022.10 m), and Phase 2 (2,022.10–2,021.06 m) are recognized by inflection
points in the δ
13
C
org
signal with recovery Phase 2 exhibiting higher δ
13
C
org
values than recovery Phase 1, albeit
values across both recovery phases do not go back to pre PETM levels.
2.2. Carbon Isotopes and Total Organic Carbon
To further constrain the two‐step CIE onset recognised in Kender et al. (2021), 39 new sediment samples were
micro‐drilled at a sample spacing of ∼5 mm (Figs. S2, S3, S4 in Supporting Information S1), covering the interval
between 2,025.52 and 2,025.32 m. The δ
13
C
org
and total organic carbon (TOC) analyses were carried out at the
NERC Isotope Geosciences Laboratory at the British Geological Survey. Inorganic carbon was removed by
treating the powdered sediments with 5% HCl overnight. TOC was then determined on bulk samples by com-
bustion in a Costech ECS4010 Elemental Analyzer calibrated against an acetanilide standard.
13
C/
12
C analyses
were performed on bulk rock samples by combustion in a Costech Elemental Analyzer on‐line to a VG TripleTrap
and Optima dual‐inlet mass spectrometer, with δ
13
C
org
values relative to VPDB using three within‐run laboratory
standards calibrated against NBS‐18, NBS‐19, and NBS‐22. The internal standards yield standard deviations of
0.07, 0.05, and 0.1. The additional values were incorporated into the isotope stratigraphy and TOC records of
Kender et al. (2021) that were measured in the same laboratory.
2.3. X‐Ray Fluorescence Spectrometry
Portable X‐ray fluorescence (pXRF) measurements were performed using an Olympus DELTA Professional X‐
ray portable spectrometer at the Camborne School of Mines, University of Exeter, on 621 bulk rock samples to
investigate major and trace elements. The pXRF analytical accuracy was checked by running Standard Reference
Material 2710a every 10 samples, with values of expected and measured major element within 1% (Fig. S5 in
Supporting Information S1). Samples were taken at ∼1 cm resolution and powdered with an agate pestle and
mortar. Homogenised samples with thickness >5 mm were measured in glass vials with non‐PVC cling film and
placed in a shielded chamber that faced the analyzer measuring window. The analysis was carried out in “geo-
chem mode” (abundance%) with a run time of 60 s to allow for reliable trace metal measurements (Dahl
et al., 2013).
To constrain the pXRF data quantitatively, analyses of titanium (Ti), aluminum (Al), iron (Fe) and silicon (Si)
were performed on a subset of the same samples (n=53) with a conventional based XRF (bXRF; Figs. S6, S7 in
Supporting Information S1). The instrument was a PANalytical Axios mAX sequential, fully automatic
wavelength‐dispersive X‐ray fluorescence spectrometer fitted with a 60 kV generator and 4 kW rhodium (Super
Sharp) end–window X‐ray tube. The software used to control the instrument and report data was PANalytical
SuperQ 5.2 A. The samples were prepared by taking 0.9 g of sample and mixing it with 9 g of 66:34 lithium
tetraborate/lithium metaborate flux and fusing for 12 min at 1,200°C. The fused glass bead was analyzed using the
Malvern Panalytical WROXI application. Background factors were calculated by either angular difference (2Ɵ)
or from regression values from “high purity” single oxide standards. Line‐overlap factors were calculated from
“high purity” oxide standards of the interfering analyte on regression. The PANalytical calibration algorithm was
used to fit calibration curves, deriving calibration constants for slope and concentration intercept. Inter‐element
effects were corrected by fundamental parameter (FP) coefficients, calculated by the PANalytical fundamental
parameter method. All backgrounds and peaks were corrected for instrument drift using two external ratio
monitors. Drift correction intensities from the monitors were used to monitor instrumental stability and detect
machine faults. The calibrations were validated by analysis of a number of Reference Materials (Table S1 in
Supporting Information S1).
Total sulfur (S) concentration analysis was performed on 25 samples using a LECO CS230 instrument at
Newcastle University in duplicate after 10 samples with a RSD of 1.5%. Certified calibration carbon/S standards
(0.815/0.0032%, respectively) supplied by Leco UK Ltd, along with analytical blanks were measured every 20
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MARIANI ET AL. 4 of 18
and 25 analytical samples, respectively. Conventional XRF data, and dilution from the
100
Mo/
95
Mo, were pri-
marily used to calibrate the pXRF data, and results were compared through regression analysis (Fig. S6 in
Supporting Information S1). As pXRF data consistently underestimated Al, Si and Mo contents, we corrected
these using linear regression equations (Fig. S7 in Supporting Information S1). We emphasize that our paleo-
environmental reconstructions are based on relative trends rather than absolute values.
2.4. Molybdenum Isotopes and Concentrations
Molybdenum isotope analysis was performed on 18 samples, four of which were taken at 2,026.85, 2,026.5,
2,025.9, and 2,025.47 m corresponding to the pre PETM interval; the sampling resolution was increased across
the CIE by analyzing five samples between 2,025.41 and 2,025.33 m, to capture any change around the CO
2
injection; five samples were selected across the PETM main phase between 2,025.29 and 2,023.5 m; two samples
were taken at 2,023 and 2,022.5 m corresponding to the recovery Phase 1 interval, and two samples were chosen
in the recovery Phase 2 at 2,022 and 2,021.5 m (Figs. S6, 7 in Supporting Information S1). Homogenised sample
powders were mixed with a precise mass of a
100
Mo‐
97
Mo double spike solution and digested in a mixture of 3:1
HNO
3
/HCl. Molybdenum was separated from matrix elements using the anion column method of Pearce
et al. (2010) as modified by Dickson et al. (2016). Purified sample solutions were measured on a Neptune Plus
MC‐ICP‐MS at Royal Holloway, University of London, using 10 ppb solutions introduced to the mass spec-
trometer via an Aridus III desolvator. Measurements consisted of 80 ×8.4s integrations. Voltages were back-
ground corrected with a short measurement of 3% HNO
3
prior to each measurement. Sample‐spike mixtures were
solved offline using an Excel‐based routine. Isotope compositions are expressed relative to NIST 3134 as:
δ98/95Mo =(( 98 Mo/95 Mosample
98Mo/95MoNIST3134 )1∗1000)+0.25‰
Uncertainties are propagated from the 2 S.E. counting statistics and the bracketing NIST 3134 solutions and are
generally similar to the long‐term external reproducibility of ±0.06‰ determined from multiple digestions and
measurements of the SDO‐1 USGS shale. A sample of SDO‐1 prepared in the same batch as the North Sea
samples had a composition of 1.09 ±0.03‰, which is identical to values from previous studies (Goldberg
et al., 2013, p. 1.05 ±0.2‰; Dickson et al., 2019, p. 1.04 ±0.08‰). A procedural blank measured by isotope
dilution was 120 pg, which is insignificant compared to the mass of Mo processed for each sample (∼400–
500 ng). Molybdenum concentrations were calculated by isotope dilution from the
100
Mo/
95
Mo ratio.
2.5. Scanning Electron Microscopy
Thirteen samples were chosen to study the size distribution of pyrite framboids as a proxy for bottom water redox
conditions, using a TESCAN VEGA3 scanning electron microscope (SEM) at Camborne School of Mines,
University of Exeter. The samples were chosen so that the sample spacing throughout the main phase, recovery
Phase 1 and 2 was constantly set to ∼30 cm to avoid sampling biases. However, this wasn't always achievable due
to scarceness of rock samples available, and in some instances sample spacing spanned from ∼10 to ∼50 cm. With
regards to the pre PETM interval, only two samples were chosen due to the generally low TOC and non‐laminated
sediments, suggesting that the pre PETM interval sedimented in mostly oxic conditions and possibly be devoid of
pyrite. One pre PETM sample was chosen at ∼10 cm before the CIE to investigate if the onset of anoxia might
have occurred before the CO
2
injection. Sediment samples were cut into sections and mounted in resin blocks,
polished with diamond paste, and carbon‐coated. The diameter of 100 pyrite framboid specimens per sample, a
representative number selected in accordance with previous global‐scale pyrite framboid studies across Permian‐
Triassic boundary sections (Bond & Wignall, 2010), were measured in backscattered electron mode (BSE).
2.6. Clay Mineralogy
The clay mineralogy of a total of 77 samples were investigated using X‐ray diffraction (XRD) analysis at the
British Geological Survey laboratories, Keyworth. About 4 g of sample was crushed and passed through a 2 mm
sieve. The <2 mm fraction was then dispersed in deionized water using an ultrasonic bath and reciprocal shaker;
the resulting suspension was sieved on a 63 μm sieve and placed in a 250 ml cylinder with few drops of 0.1 M Na₆
[(PO₃)₆] to prevent flocculation and to disperse the clay particles. The suspension was left to settle for periods
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determined by Stokes' Law, after which the “clay” (<4 μm) and “silt” (4–63 μm) fractions were dried at 55°C and
stored in glass vials. Although clays are classified as <2 μm particles, a larger particle diameter (i.e., <4 μm) is
preferred in smectite‐rich sediments (Kemp et al., 2016). The production of oriented mounts, as well as XRD
analysis and data modeling followed the methods outlined by Kemp et al. (2016).
3. Results and Discussion
3.1. Development of Euxinia During the PETM
Terrigenous elements Al and Ti display their highest relative abundance values during the pre‐PETM interval,
followed by a decrease coinciding with the CIE onset (∼2,025.4 m; Figure 2). The relative abundances slightly
increase during the recovery phases (∼2,023–2,021 m), but do not reach pre‐PETM values. Si, like Al and Ti, is
present in the aluminosilicate phase although it can also originate from quartz silts and sands, and biogenic silica.
During the pre‐PETM, Si is constantly around 30% followed by a decrease of ∼5% at the CIE onset (∼2,025.4 m;
Figure 2). The highest relative abundance values coincide with the silty interval observed between 2,024.5 and
2,023.1 m (Figure 2).
Iron relative abundance remains overall constant throughout the interval with the exception of an increase at the
CIE onset (∼2,025.4 m) and the main phase between 2,025.4 and 2,024.5 m (Figure 2). Sulfur exhibits its lowest
values of the entire record during the pre‐PETM, followed by a major increase at the CIE onset (∼2,025.4 m), and
by a decrease in the lower part of the main phase (∼2,025.2 m–∼2,024.3 m). Sulfur gradually decreases
throughout the recovery phases, reaching lowest levels during recovery Phase 2 when compared to recovery
Phase 1 and main phase, albeit not declining to pre PETM levels (2,022–2,021 m; Figure 2). The sharp increase of
S at the CIE onset may suggest incorporation under reducing conditions (Croudace & Rothwell, 2015; Sluijs
et al., 2008; Figure 2).
Major and trace elements are normalized against Al and Ti to account for the dilution effects of terrestrial
sediment input rate (Tribovillard et al., 2006). The most significant geochemical changes in core E8X are
observed at the CIE onset (∼2,025.4 m, Figures 2and 3) where all redox proxies point to a shift from oxic /suboxic
Figure 2. E8X lithological log, δ
13
C
org
and elemental analysis. Black curves represent data acquired with pXRF (adjusted
using bXRF for Al, Si, and Mo, see Section 2), red dots correspond to datapoints collected with bXRF and
100
Mo/
95
Mo.
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to anoxic/euxinic bottom water settings. For instance, the Fe/Al increase at the CIE onset (Figure 3) suggests that
a portion of Fe is present in pyrite rather than in detrital material (Lyons & Severmann, 2006).
Aqueous Mo, another diagnostic redox‐sensitive proxy, is abundant in modern oxygenated marine basins where it
is unreactive (Algeo & Lyons, 2006). However, under anoxic conditions Mo is transferred into the sediment (Helz
et al., 1996,2011; Supplementary Material). The rapid increase in Mo/Ti at the early onset of the PETM CIE is
characteristic/diagnostic of a fast transition to euxinic settings (Figure 3). Molybdenum isotope ratios (δ
98/95
Mo)
have been applied as a seawater redox proxy during the PETM in the Arctic and Tethys Oceans (Dickson
et al., 2012; Dickson, Cohen, & Coe, 2014; in Supporting Information S1), as
95
Mo is preferentially removed
during Mo absorption onto Mn oxides in oxic settings, which are fractionated from seawater by 3‰ (Dick-
son, 2017). Anoxic environments of intermediate reducing conditions, with H
2
S<11 μmol l
1
, display a near
constant seawater isotopic fractionation of ∼0.7‰ due to non‐quantitative Mo drawdown (Dickson, 2017). When
H
2
S>11 μmol l
1
, a geochemical switch point is activated under which Mo exists as tetrathiomolybdate and can
be removed from the water column (quantitative Mo drawdown), yielding low seawater isotopic fractionations of
0.5‰ (Dickson, 2017). The shift from light (0.11‰ to 0.84‰) to heavy (1.27–1.99‰) sediment δ
98/95
Mo
isotopic composition in E8X indicates broadly oxic conditions before the PETM (note a temporary shift to
heavy values at 2025.93 m where the sediment becomes briefly laminated and enriched in TOC; Figure 3) that
rapidly changed to sulfidic conditions at the CIE. The first half of the main phase, between ∼2025.2 m and
∼2024.5 m, records a slight return to lighter δ
98/95
Mo values. The second half of the main phase between
∼2024.5 m and ∼2023.2 m is characterized by heavy values around 2‰. Although a slight decrease is observed
throughout the recovery phases, δ
98/95
Mo values remain above 1‰.
Pyrite framboid diameters have also been used as a proxy for redox conditions (Bond & Wignall, 2010; Wilkin
et al., 1997). Pyrite framboids form via two pathways based on bottom water oxygen levels, with mean diameters
between 3 and 5 μm occurring in euxinic settings, and larger diameters up to 50 μm typical of suboxic conditions
(Wilkin et al., 1997; Bond & Wignall, 2010; in Supporting Information S1). The pyrite framboid size distribution
Figure 3. Data from pXRF, Mo isotopes and pyrite framboids from core E8X. Core photos and simplified lithological log
are from Kender et al. (2021). The δ
13
C
org
and TOC (wt%) values presented herein are a compilation of data from this study
and previous measurements from Kender et al. (2021). The box and whisker plots of framboid size consist of first quartile
(25th percentile) and third quartile (75th percentile). The vertical line within each box is the median value, and the X
represents the average value. Whiskers symbolize minimum and maximum values, excluding outliers (individual dots).
Outliers are identified through the 1.5 interquartile range rule. The dashed blue line indicates the maximum average diameter
for syngenetically formed framboids. An additional dashed blue line along the Mo isotope curve indicates a δ
98/95
Mo value
of ∼1.94‰, corresponding to the peak values observed in Site M0004 A (Arctic Ocean) and interpreted as global paleo‐sea
water isotopic composition (Dickson et al., 2012). The age model shows ‘option A and B' from core BH9/05 (Charles
et al., 2011) (see Kender et al., 2021).
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at the CIE onset supports syngenetic growth under euxinic conditions (Figures 3and 4a). No pyrite framboids are
observed under the scanning electron microscope in the sample at ∼2026 m and the pre PETM interval in general
contains much lower pyrite abundance as inferred from low S% and Fe/Al (Figures 2and 3). Framboids are
observed just prior to the CIE onset (2025.5 m, Figure 3), yielding a wide range of sizes and an average diameter
of ∼6 μm indicating suboxic conditions. Larger pyrite framboids are present throughout the recovery phase in-
tervals, where diameters reach up to 20 μm, typical of anoxic/suboxic environments (Figures 3and 4b, c). The
framboids become more sporadic at the recovery Phase 2, suggesting an attenuation of anoxia (Figures 4b and 4c).
The presence of pyrite is further corroborated by the positive correlation between Fe% and S%, whereas the strong
positive correlation between Fe/Al and S/Al suggests that some of Fe and S is bound to the clay source (Figures 4f
and 4g).
To summarize, the pre‐PETM is characterized by predominantly weakly laminated sediments, low pyrite (S and
Fe) and Mo abundances generally below the limit of detection (Figure 2). We interpret these proxy patterns (lack
Figure 4. Pyrite framboid images, spectral analysis, elemental maps and cross plots. (a) Main phase sample (MPA 63928, 2025.05 m) displaying numerous packed
framboids. (b) Recovery Phase 1 sample (MPA 66450, 2022.51 m) showing isolated and, occasionally, clustered framboids. (c) Recovery Phase 2 sample (MPA 66351,
2021.5 m) with rare framboids and pyritized silicofossil fragments. (d) Spectral analysis of framboids in the main phase sample (spectrum d in a). (e) Elemental mapping
of framboid clusters in the recovery Phase 1 sample (cluster ‘e’ in b), clearly showing pyrite framboids in a clay‐dominated matrix. (f) Handheld XRF Fe% versus S%.
(g) Handheld XRF Fe/Al versus S/Al.
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of clearly laminated sediments, low pyrite content and no appreciable Mo concentrations) typical of oxygenated
bottom water conditions. Short suboxic episodes appear to have occurred, consistent with the observation of
occasional laminations, high Mo abundance horizons (up to 20 ppm, concentrations significantly above crustal
values), and framboid sizes typical of suboxic conditions (Figures 2and 3). Unlike the pre‐PETM, the CIE onset
and PETM main phase record continuous sediment lamination, high pyrite (Fe and S), small diameter pyrite
framboids (<5 μm), rapid increase in Mo abundance and heavier Mo isotopes, typical of euxinic conditions and
significant trace metal depletion (Figures 2and 3). This profound redox transition is consistent with previous
observations from Fur Island (Figure 1), and Arctic Ocean Site M0004 A, where laminated sediment and S‐bound
isorenieratane presence is indicative of photic zone euxinia (Schoon et al., 2015; Sluijs et al., 2006; Stokke
et al., 2021).
Comparison of PETM sites across the North Sea‐Norwegian Sea‐Arctic Ocean (Figure 5) shows that laminated
sedimentation appeared approximately simultaneously at the onset of the PETM across the region, which may be
indicative of anoxia/euxinia as we have found for E8X. Laminated sediments at the PETM onset are reported
from Fur Island, Denmark (Jones, Percival, et al., 2019), which is closest to our site. In addition, recent sites
U1567 and U1568, a collection of five holes from International Ocean Discovery Program (IODP) Expedition
396 (Planke et al., 2023) forming a transect across the Modgunn hydrothermal vent complex (HTVC) in the Møre
Basin, all show a change to laminated mudstone coeval with the onset of the PETM (Berndt et al., 2023).
However, sites reverted to unlaminated sedimentation at varying times within the main and recovery phases. This
may have been driven predominantly by a shrinking oxygen minimum zone (OMZ) and expanding mixed depth
that caused shallower sites to become oxic before deeper sites. Although we cannot accurately reconstruct paleo
water depth, the sites with lower sedimentation rates appear to retain laminations for a longer period of time (e.g.,
E8X) which may indicate that deeper sites were affected by anoxia/euxinia for longer as sedimentation rates
tend to be highest in shallow and nearshore settings (e.g., Svalbard site BH9/05, where palynology interpretations
indicate marginal marine conditions, Harding et al., 2011; Fur, which has been interpreted as outer neritic, Schoon
et al., 2015). Although we are unable to correlate Sites U1567 and U1568 with the other records in Figure 5as the
CIEs are of slightly lower resolution in places (Berndt et al., 2023), we do note that these sites revert to unla-
minated sedimentation during the PETM main phase which is consistent with their interpreted shallow marine
water depth (Berndt et al., 2023). By contrast, the rapid infill of the Modgunn hydrothermal vent complex on the
Figure 5. Comparison of PETM sites across the North Sea, Norwegian Sea and Arctic. (a) Map showing the location of North Sea sites (22/10a‐4, E8X and Fur), IODP
396, Spitsbergen section BH9/05 and Lomonosov Ridge site M0004 A. Figure adapted from Jones et al. (2023), with Arctic Ocean reconstruction based on
Golonka (2009). (b) δ
13
C curves from Fur (black line), E8X (blue line), 22/10a‐4 (red line), BH9/05 (orange line) and M0004 A (gray line). (c) Simplified lithological
logs of PETM sites showing onset of laminations around the CIE. (d) Schematic North Sea and Arctic N‐S section showing changes in redox conditions as a function of
water depth across the North Sea, Norwegian Sea and the Arctic.
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Norwegian continental margin means that the water depth of the lowermost‐PETM strata may be considerably
greater than at the top of the vent infill, which is still within the PETM, because Hole U1568A contains ∼80 m of
syn‐PETM infill (Berndt et al., 2023). Therefore, the potential effect of water depth on recording anoxia may be
affected by the rapid infill of the vent crater that was initiated as a significant bathymetric depression. Our
interpretation of a rapid onset of regional euxinia at the PETM onset, followed by gradually contracting OMZ and
expansion of the mixed layer (perhaps via reduced salinity stratification: Zacke et al., 2009; Kender et al., 2012),
suggests the North Sea–Norwegian Sea–Arctic basins were highly tectonically restricted and sensitive to further
tectonic uplift and hydroclimate‐driven productivity changes.
3.2. Basin Restriction and NAIP Emplacement
Our Mo/TOC data provides evidence for basin restriction of the North Sea over the PETM (Figure 6). It has been
shown that in silled basins the limited aqueous Mo availability results in reduced sedimentary burial of Mo per
unit of organic matter (the ‘basin reservoir effect’; Algeo & Lyons, 2006; in Supporting Information S1), and
consequently the Mo/TOC ratio can be used as a basin restriction proxy. Mo/TOC values in E8X are highest
(∼0.4 ×10
2
) during the CIE onset (∼2025.4 m) and gradually decrease throughout the main phase of the PETM to
values between 0.05 and 0.15 (Figure 6). We suggest that the high Mo/TOC values at the CIE onset resulted from
the initial drawdown of aqueous Mo from the previously less restricted North Sea basin. Similar behavior has
been observed at the start of Ocean Anoxic Event (OAE) two in the hydrographically restricted Proto‐North
Atlantic Ocean (Owens et al., 2016). Mo/TOC values in core E8X remain low for the main phase of the
PETM consistent with a strongly restricted basin with deep water renewal times >500 years (Algeo &
Lyons, 2006). This is further supported by Osmium (Os) isotope records from nearby Fur Island pointing to strong
basin restriction during the PETM (Jones et al., 2023).
Additional evidence of basin restriction during the PETM is found in the E8X δ
98/95
Mo isotopic signature:
assuming global seawater composition of ∼1.95‰ for the PETM (Dickson et al., 2012), the E8X δ
98/95
Mo
values of ∼2‰ during the main phase (Figure 3) are similar to Arctic Ocean marine sediments (Dickson
et al., 2012). Such heavy isotopic values are achieved under quantitative Mo drawdown in euxinic conditions,
which result in direct capture of seawater isotopic composition. However, quantitative Mo removal in euxinic
Figure 6. δ
13
C
org
, TOC (wt%), Mo/TOC, Si/Al ratio and clay minerals of core E8X. The TOC (wt%) from Kender
et al. (2021) indicates relative changes in productivity and Mo/TOC as a basin restriction proxy. Si/Al and clay mineralogy
are used to show the influence of coarse and fine‐grained sediments, respectively. Mercury data from Kender et al. (2021).
The age model option A and B are from Charles et al. (2011) (see Kender et al., 2021).
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settings can be hindered if exchange with open ocean water is high. Consequently, sedimentary δ
98/95
Mo
approaching the global ocean value is usually associated with euxinic and restricted basins (Dickson, 2017). As
E8X Mo isotope values are similar to those observed in the Arctic Ocean (Dickson et al., 2012), we use these to
interpret comparable depositional conditions, characterized by euxinia and trace metal drawdown under strong
hydrographic restriction. Box modeling has been carried out on the fractional removal of Mo in euxinic sediments
as a function of the isotopic composition of modern, PETM, OAE 2, and Toarcian‐OAE seawater and a steady
riverine input (Dickson, 2017). According to the model, lower seawater Mo isotope compositions require higher
removal of Mo into sediments, and therefore a larger extent of seafloor euxinia compared to the present.
We suggest that the North Sea basin restriction during the PETM was caused by NAIP emplacement. Several lines
of evidence support regional uplift and relative sea level fall associated with NAIP emplacement, including
lithostratigraphy and seismic stratigraphy of Paleocene–Eocene successions around the North Sea (Hartley
et al., 2011; Maclennan & Jones, 2006; Svensen et al., 2004). The uplift was primarily caused by pulsating upward
flow of the Icelandic mantle plume between Greenland and Scotland, which spread outwards beneath the NW
continental margin (Hartley et al., 2011; Jones, Hoggett, et al., 2019; Figure 1). It is important to note that
sedimentary Hg and Hg/TOC values in E8X and nearby cores 22/10a‐4 (Kender et al., 2021) and 23/16b‐9 (Jin
et al., 2023), in addition to Hg records from Fur and Svalbard (Jones et al., 2019a,2023), indicate that volcanic
activity was enhanced in the lead up to and during the CIE onset (Figure 6), suggesting a direct and causal as-
sociation of NAIP volcanism and the PETM. Additional sedimentary Hg enrichments coinciding with the PETM
have been documented from other study areas globally and have been related to NAIP volcanic activity (Liu
et al., 2019; Tremblin et al., 2022). Thousands of intruded sills have been recognised in 2D and 3D seismic
reflection profiles between the Faroe‐Shetland basin (offshore NW Scotland) and the Vøring and Møre basins
(Norwegian Sea), with up to 400 m thickness (Berndt et al., 2023; Jones, Percival, et al., 2019; Svensen
et al., 2004). Although the vertical displacement and uplift rates are yet to be constrained, we hypothesize that the
thick NAIP sills, which may have been emplaced rapidly (possibly down to a few years, Jones, Percival,
et al., 2019), must have had a significant topographic impact on the northwestern North Sea basin and on the
northern North Atlantic region. Further evidence of a shallow and restricted northern North Atlantic comes from
IODP Sites U1567 and U1568, where five holes through the Modgunn hydrothermal vent complex were drilled,
and 3D seismic mapping was performed. The results suggest that the Modgunn vent complex formed a few
thousand years before the PETM onset under shallow marine waters, whereas other hydrothermal vent complexes
in the Modgunn Arch might have developed under short‐lived subaerial settings (Berndt et al., 2023), supporting
the assumption that the proto‐Northeast Atlantic was possibly very susceptible to topographic changes. The
interpretation of sedimentological, palynological and 2D seismic stratigraphic data propose that a complex
fragmentation of both the northeastern North Atlantic (Norway‐Greenland seaway) and the western North
Atlantic (Greenland‐Scotland Ridge) occurred at the Paleocene‐Eocene, as a result of regional tectonic uplift
(Hovikoski et al., 2021).
Basin restriction via uplift and relative sea level fall may explain our Si/Al and clay mineralogy data (Figure 6).
High Si/Al values in E8X could reflect biogenic silica from diatoms, and/or quartz content in the silty interval as
observed in the fine fraction XRD analysis. During the main phase the clay assemblage exhibits higher illite
within a smectite‐dominated background, and trace occurrence of kaolinite (Figure 6). Kaolinite is composed of
coarser grains, with illite and then smectite composed of progressively finer grain sizes. In the Paleogene North
Sea basin, high abundances of kaolinite and illite at the basin margins and a high abundance of smectite in the
central area were interpreted as a consequence of current sorting during transport (Nielsen et al., 2015; Sup-
plementary Material). Sediment size coarsening during the main phase is supported by increased illite content
versus smectite and peaks of Si/Al in the silty interval between 2,024.5 and 2,023.2 m (Figure 6). We argue that
local relative sea level fall from tectonic uplift could have shifted the paleo shorelines making E8X closer to
continental sediment sources, where coarser particles reached the site. Alternatively, increased grain size may
reflect an increased flux of terrigenous material, as suggested by lithium isotope excursion records at the PETM
onset interpreted to be a response to enhanced hydrological cycle due to global warming (Pogge von Stradmann
et al., 2021; Jones et al., 2023).
3.3. Precursor Signals and an NAIP Trigger for the PETM
The PETM was a complex event: in addition to global temperature increase (Dunkley Jones et al., 2013), some sea
surface temperature and continental proxies hint to a transient warming prior to the CIE onset (Secord et al., 2010;
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Sluijs et al., 2007), while marine environments display highly variable regional and local water column salinity
and stratification (Sluijs & Brinkhuis, 2009). Areas proximal to the NAIP emplacement area, such as the North
Sea, may have faced far more direct and local consequences to the tectonic uplift and sill intrusions, such as
precursor cooling from volcanic SO
2
injections (Stokke et al., 2021), which are not observed in other distal lo-
cations globally. Alternatively, the tectonic uplift resulting from NAIP emplacement might have caused an
episodic cooling event by modifying the positioning of the Northern Hemisphere jet stream (Vickers et al., 2024).
These proximal areas may therefore hold the key to separate trigger and response during the initial onset of the
CIE, specifically the role of volcanic uplift and emissions as causal triggers for the development of environmental
change during the PETM main phase. The North Sea–Norwegian Sea–Arctic Ocean were restricted with a strong
low salinity halocline in place (Harding et al., 2011; Kender et al., 2012; Zacke et al., 2009), and the OMZ would
have been sensitive to changes in tectonic restriction, and riverine influxes of nutrients causing plankton blooms
perhaps in some ways similar to the modern Black Sea (Bologa et al., 1999). Plankton blooms during the PETM
have been documented in the northwestern North Sea (Eldrett et al., 2014; Kender et al., 2012) and offshore the
Møre Basin, Norway (Berndt et al., 2023).
Before the PETM, in E8X a stratigraphically short occurrence of laminated sedimentation (∼2,025.94 m;
marked as “E1” in Figure 7) occurred with increased Mo, pyrite (S and Fe), δ
98/95
Mo, TOC and negative δ
13
C.
This indicates a brief period of anoxia/euxinia driven by elevated productivity (TOC), which may be linked to
volcanism as there is a coeval increase in Hg/TOC. Event E1 is unlikely to have been caused directly by basin
restriction from NAIP uplift, as its short‐lived nature would imply rapid subsequent tectonic relaxation and
subsidence in contrast to evidence for gradually increasing sill emplacement and volcanic intensity (Berndt
et al., 2023; Jones, Hoggett, et al., 2019; Kender et al., 2021). We speculate that volcanic activity releasing Hg and
carbon (negative δ
13
C spike) caused oceanographic change which translated into eutrophication of the North Sea.
This event may also be expressed in 22/10a‐4 (Kender et al., 2012) where there are short‐lived precursor δ
13
C
spikes associated with eutrophication, but it has not been detected outside the North Sea region although we note it
is short‐lived and could have been missed at sites with lower sampling resolution.
We suggest that this apparently local (to the North Sea) precursor eutrophication E1 signal was unlikely the result
of massive greenhouse gas release causing global warming and an increased hydrological cycle (as suggested for
the PETM). One line of evidence comes from terrestrial palynological data around the North Sea, which are
inconsistent with pre‐PETM warming and rather depict a vegetation dominated by temperate conifers up to the
Figure 7. Carbon release, productivity, redox and volcanic proxies in core E8X. Precursor events E1 and E2, and the two‐
step CIE are shown.
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CIE onset (Collinson et al., 2009; Kender et al., 2012; Lenz & Riegel, 2001). Conversely, the CIE onset is
associated with data pointing to climatic perturbations possibly due to the carbon input: higher plant biomarkers
and their hydrogen isotopic composition at Vasterival, Southern North Sea, suggest a shift from dryer to moister
conditions due to increased meteoric water input at the CIE onset (Garel et al., 2013). A change from dry to
warmer and wetter conditions at the CIE onset is also reconstructed from carbon isotopes of the biomarker
hopanoids at Cobham, Southern North Sea (Pancost et al., 2007). Likewise, Cobham palynological data records a
shift from dry fire‐prone pre‐PETM vegetation to wetter conditions and decreased fire regime during the PETM
(Collinson et al., 2009), suggesting a profound climatic change consistent with the CIE onset. As an increase of
Hg/TOC is consistent with event E1 (Figure 7), and given the paucity of evidence suggesting major climate
changes during the pre‐PETM interval, we argue that E1 was rather controlled by volcanic NAIP‐related phe-
nomena perhaps from atmospheric S and ash affecting regional rainfall and riverine discharge, and/or nutrient
delivery from massive hydrothermal venting directly into ocean water (Svensen et al., 2004). A subsequent
possible PETM precursor event occurs in E8X (∼2,025.65 m; “E2” in Figure 7), which is similar to E1 (negative
δ
13
C spike, TOC and pyrite increase) but smaller and with no increase in Mo, is associated with increased Hg/
TOC and may have had similar causal mechanisms. Alternatively, increased productivity during E1 and E2 may
have been the result of Milankovitch‐forced local insolation and rainfall‐related changes interacting with a basin
close to a tipping point, with the Hg increases being coincidental. Either way, the presence of E1 and E2 indicate
that the North Sea basin had already become restricted by uplift from previous activity of the NAIP (evidenced by
precursor Hg spikes, Kender et al., 2021; geochronological constrains of ash layers and volcanic sequences across
Greenland, Faeroe Islands, UK, and Norway predating the PETM onset; Storey et al., 2007; Wilkinson
et al., 2017), such that an ephemeral increase in productivity was able to cause OMZ expansion and euxinia.
Further evidence for basin restriction prior to the PETM comes from the Northern North Sea, where at site 22/10a‐
4 low salinity/eutrophic dinoflagellate cysts gradually increase in proportion suggesting early oceanic eutro-
phication from enhanced productivity (Kender et al., 2012). As we find no early long‐term increase in produc-
tivity (TOC) in the Southern North Sea (E8X), we proposed the 22/10a‐4 signal was related to continued uplift
of the NAIP causing local increases in riverine runoff perhaps from an uplifted Faeroe–Shetland platform (Kender
et al., 2012). There is no evidence for increased warming prior to the PETM in the North Sea basin as a
mechanism to force such a change; instead, sea surface temperature proxy data indicates a possible cooling
episode, although this is complicated by high BIT index values and associated uncertainties of interpreting the
data (Schoon et al., 2015; Stokke et al., 2020).
At the onset of the PETM (CIE step 1; Figure 7) we observe elevated δ
98/95
Mo, pyrite and Mo/TOC, whereas the
increase in TOC lags δ
98/95
Mo, pyrite and Mo/TOC by one sample, corresponding to ∼300 years time step ac-
cording to our sample resolution and simple age model (Figure 3; Kender et al., 2021), again occurring with
elevated Hg/TOC. Although this apparent time lag could be explained by further tectonic uplift and restriction of
the North Sea causing euxinia before any major climate‐driven eutrophication, it is possible that the onset of
bottom water euxinia changed pore water redox chemistry several cm below the sea floor. Therefore, it is
consistent that enhanced productivity forced the early OMZ expansion (as in precursor events E1 and E2) either
by global climate change at the PETM onset, or more likely local eutrophication from volcanic processes as in the
precursor events given this change was so early at the CIE onset. CIE step 2 corresponds to a further release of
CO
2
that may have been triggered from initial volcanism, and the pulsed increases in TOC suggest global
warming and increased global hydrological cycling was continuing to more heavily impact the productivity
regime in the North Sea (Figure 7).
3.4. Recovery From the PETM
A partial attenuation of both basin restriction (Mo/TOC) and euxinia (S, Fe/Al, and pyrite framboid size) is
observed from the base of recovery Phase 1, ∼70 kyr after the main carbon isotope excursion (Figures 3and 6). A
decrease of S and Fe/Al compared to main phase values suggests anoxia‐suboxia, in contrast to the high Mo and
Mo/Ti values typical of persistent euxinia. We interpret this discrepancy during recovery Phase 1 as Mo/TOC
values indicating a shift from strong to moderate basin restriction, which would have allowed at least partial
renewal of aqueous Mo. Thus, we suggest that high Mo and Mo/Ti values in recovery Phase 1 are, for the most
part, a result of Mo renewal rather than a response to changing redox conditions. Mo renewal could have been
caused by tectonic relaxation and minor subsidence, and/or sea level rise from climatic changes and ocean
warming. Sulfidic conditions likely still existed at this time, as supported by high δ
98/95
Mo, but were located
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closer to the seafloor given the increase in pyrite framboid size distributions that indicate dominantly authigenic
rather than syngenetic processes.
Both Mo and Mo/Ti indicate anoxic‐suboxic conditions during the recovery Phase 2, a paleoenvironmental
setting which is supported by the pyrite framboid size distribution. We suggest that at this time, the Mo/TOC
signal loses its reliability because the proxy is applicable only to systems where at least weakly sulfidic conditions
existed (Algeo & Lyons, 2006). The gradual decrease in δ
98/95
Mo values during both recovery phases may
represent a gradual deepened chemocline either following partial opening of the North Sea basin, and/or reduced
primary productivity due to partial climate recovery, combined with overall lower H
2
S availability (Figure 3).
Relatively high Mo concentrations throughout the recovery phases, however, suggest that H
2
S was still present at
the local sediment‐water interface.
Weakened basin restriction (rising Mo/TOC values) and a decrease in coarser sediment (low Si /Al values, a return
to pre PETM proportions of illite and smectite) is consistent with a relative sea level rise (tectonic and/or climate
driven), defining the local response of the North Sea region to the recovery Phase 1 of the PETM. In the North
Sea, anoxic‐suboxic bottom water conditions persisted throughout the recovery phases at E8X despite possible
sea level rise. We speculate this could be related to the slow (Myr) NAIP subsidence from waning effects of
thermal uplift (Hartley et al., 2011) which kept the North Sea moderately restricted for large parts of the recovery
phases with intermittent anoxic‐suboxic periods, and the gradual reduction in the extent of the OMZ affecting
deeper sites for longer (Figure 5). It is important to note that our suggested hypothesis of partial attenuation of
basin restriction during the recovery phases is somewhat in contrast with other sites in the North Sea. Osmium
isotope records (
187
Os/
188
Os) from Fur Island indicate a strong deviation from the global ocean signal coinciding
with the start of the recovery phase, suggesting that the North Sea became restricted after the hyperthermal event
(Jones et al., 2023). However, the Os isotopic shift at the end of the PETM correlates with a variation in lithium
isotopes (δ
7
Li), a proxy for siliciclastic input and erosion of ash deposits. This suggests an increase in basalt‐
derived fluxes during the recovery phase, and if these fluxes have different Os isotopic signatures, then the
reliability of
187
Os/
188
Os as a proxy for water connectivity might be hindered. Although quantifying the extent of
basin restriction relaxation during the recovery phases is beyond the scope of this study, it is important to note that
restriction changes might have been more subtle than our proposed shift from ‘strongly restricted’ to ‘moderately
restricted’ North Sea after the PETM main phase. This would be more aligned with salinity records (Zacke
et al., 2009) and uplift reconstructions (Conway‐Jones & White, 2022) that suggest the North Sea remained
somewhat restricted during the early Ypresian.
In core E8X, the main phase and recovery phases may cover a time span of ∼140 kyr. According to our data, the
bottom water environment responded close to instantaneously (∼300 years) to the hydrographic restriction
threshold, and never fully recovered to pre PETM oxic conditions for at least ∼140 kyr. Sustained anoxic con-
ditions are prone to carbon burial and when coupled with increased silicate weathering, required to resolve the
short duration of the PETM recovery (Penman & Zachos, 2018), our data support that North Sea and nearby
basins could have been important carbon sinks throughout the observed interval (Papadomanolaki et al., 2022;
Stokke et al., 2021). Future modelling studies will allow assessment of the role this local and relatively small
carbon sink played as part of the global PETM perturbance.
4. Conclusions
New high‐resolution geochemical data from exceptionally well‐preserved marine sediment core E8X are used
to reconstruct hydrography and bottom water conditions during the PETM in the North Sea basin. Bottom water
oxygen depletion is commonly associated with climatic responses to atmospheric warming, such as increased
hydrological cycling, water column stratification, and eutrophication due to increased nutrient input. However,
we find that North Sea redox, productivity and basin restriction changes occurred before the major carbon in-
jection and global strong climatic changes associate with the PETM. Short‐lived precursor PETM events (E1 and
E2) co‐occur with low δ
13
C, high productivity, low oxygen and volcanic pulses (Hg/TOC), evidencing a
tectonically restricted basin where Large Igneous Province (LIP)‐related activity caused eutrophication, forcing
the North Sea past a tipping point toward oceanic anoxia. This could be related to atmospheric ash and S releases
causing local changes in hydrology, river runoff and nutrient discharge, and/or hydrothermal venting injecting
nutrients along with Hg directly into the North Sea. A similar scenario is interpreted for the PETM onset where
rapid early eutrophication and euxinia demonstrate LIP activity control over North Sea oceanography before the
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largest global warming and climatic changes occurred. Rapidly declining Mo/TOC with continued high δ
98/95
Mo
in the main phase of the PETM further evidences a restricted basin associated with activity of the NAIP. Early
PETM recovery is associated with a slightly less restricted North Sea Basin, which could have been due to minor
tectonic subsidence after the NAIP uplift associated with the PETM onset, and/or sea level rise driven by climatic
changes and ocean warming. Nevertheless, persistent high productivity and low oxygen throughout our records
show the North Sea Basin never fully recovered to pre PETM oxic conditions for at least ∼140 kyr.
Our findings suggest LIP emplacements should not be disregarded as a driver of significant biogeochemical
changes at time scales similar to modern climatic change, although limited to a regional scale and proximal
localities. The timing and lead‐lags between carbon injection, eutrophication and basin restriction provide further
evidence that NAIP activity was a causal trigger for the carbon cycle disruption. This fine temporal separation of
geochemical proxy evidence for volcanic versus climatic mechanisms, combined with the pulsed nature of NAIP
volcanism, demonstrates that deep mantle dynamics related to LIP emplacements can, at least locally, result in
strong ocean biogeochemical feedbacks, at time scales commonly associated with climatic response. This concept
can be tested for other LIP‐shelf sea regions to understand the cascade of trigger–response mechanisms that may
lead to global hyperthermal events, or failed events.
Data Availability Statement
Data files are publicly available at PANGAEA – Data Publisher for Earth & Environmental Science (Mariani
et al., 2024).
References
Algeo, T. J., & Lyons, T. W. (2006). Mo–total organic carbon covariation in modern anoxic marine environments: Implications for analysis of
paleoredox and paleohydrographic conditions. Paleoceanography,21(1), PA1016. https://doi.org/10.1029/2004PA001112
Berndt, C., Planke, S., Alvarez Zarikian, C. A., Frieling, J., Jones, M. T., Millett, J. M., et al. (2023). Shallow‐water hydrothermal venting linked to
the Palaeocene‐eocene thermal maximum. Nature Geoscience,16(9), 803–809. https://doi.org/10.1038/s41561‐023‐01246‐8
Bologa, A. S., Frangopol, P. T., Vedernikov, V. I., Stelmakh, L. V., Yunev, O. A., Yilmaz, A., & Ogun, T. (1999). Distribution of planktonic
primary production in the Black Sea. In S. T. Beşiktepe, Ü. Ünlüata, & A. S. Bologa (Eds.), Environmental degradation of the Black Sea:
Challenges and remedies,NATO Science Series (Vol. 56, pp. 131–145). Springer. https://doi.org/10.1007/978‐94‐011‐4568‐8_9
Bond, D. P. G., & Wignall, P. B. (2010). Pyrite framboid study of marine Permian‐Triassic boundary sections: A complex anoxic event and its
relationship to contemporaneous mass extinction. Geological Society of America Bulletin,122(7–8), 1265–1279. https://doi.org/10.1130/
B30042.1
Carmichael, M. J., Inglis, G. N., Badger, M. P. S., Naafs, B. D. A., Behrooz, L., Remmelzwaal, S., et al. (2017). Hydrological and associated
biogeochemical consequences of rapid global warming during the Paleocene‐Eocene Thermal Maximum. Global and Planetary Change,157,
114–138. https://doi.org/10.1016/j.gloplacha.2017.07.014
Champion, S. M. E., White, N. J., Jones, S. M., & Lovell, J. P. B. (2008). Quantifying transient mantle convective uplift: An example from the
Faroe‐Shetland basin. Tectonics,27(1), TC1002. https://doi.org/10.1029/2007TC002106
Charles, A. J., Condon, D. J., Harding, I. C., Pälike, H., Marshall, J. E. A., Cui, Y., et al. (2011). Constrains on the numerical age of the Paleocene‐
Eocene boundary. Geochemistry, Geophysics and Geosystem,12(6), Q0AA17. https://doi.org/10.1029/2010GC003426
Collinson, M., Steart, C. C., Harrington, G. J., Hooker, J. J., Scott, A. C., Allen, L. O., et al. (2009). Palynological evidence of vegetation dynamics
in response to palaeoenvironmental change across the onset of the Paleocene–Eocene thermal maximum at Cobham, Southern England. Grana,
48(1), 38–66. https://doi.org/10.1080/00173130802707980
Conway‐Jones, B. W., & White, N. (2022). Paleogene buried landscapes and climatic aberrations triggered by mantle plume activity. Earth and
Planetary Science Letters,593, 117644. https://doi.org/10.1016/j.epsl.2022.117644
Croudace, I. W., & Rothwell, R. G. (2015). Micro‐XRF studies of sediment cores: Applications of a non‐destructive tool for the environmental
sciences. Springer Science+Business Media. https://doi.org/10.1007/978‐94‐017‐9849‐5
Cui, Y., Kump, L. R., Ridgwell, A. J., Charles, A. J., Junium, C. K., Diefendorf, A. F., et al. (2011). Slow release of fossil carbon during the
paleocene‐eocene thermal maximum. Nature Geoscience,4(7), 481–485. https://doi.org/10.1038/ngeo1179
Dahl, T. W., Ruhl, M., Hammarlund, E. U., Canfield, D. E., Rosing, M. T., & Bjerrum, C. J. (2013). Tracing euxinia by molybdenum con-
centrations in sediments using handheld X‐ray fluorescence spectroscopy (HHXRF). Chemical Geology,360–361, 241–251. https://doi.org/10.
1016/j.chemgeo.2013.10.022
Dickson, A. J. (2017). A molybdenum‐isotope perspective on Phanerozoic deoxygenation events. Nature Geoscience,10, 721–726. https://doi.
org/10.1038/ngeo3028
Dickson, A. J., Cohen, A. S., & Coe, A. L. (2012). Seawater oxygenation during the paleocene eocene thermal maximum. Geology,40(7),
639–642. https://doi.org/10.1130/G32977.1
Dickson, A. J., Cohen, A. S., & Coe, A. L. (2014). Continental margin molybdenum isotope signatures from the early Eocene. Earth and Planetary
Science Letters,404, 389–395. https://doi.org/10.1016/j.epsl.2014.08.004
Dickson, A. J., Idiz, E., Porcelli, D., & Van den Boorn, S. H. J. M. (2019). The influence of thermal maturity on the stable isotope compositions
and concentrations of molybdenum, zinc and cadmium in organic‐rich marine mudrocks. Geochimica et Cosmochimica Acta,287, 205–220.
https://doi.org/10.1016/j.gca.2019.11.001
Dickson, A. J., Jenkyns, H. C., Porcelli, D., Van den Boorn, S., & Idiz, E. (2016). Basin‐Scale controls on the molybdenum‐isotope composition of
seawater during oceanic anoxic event 2 (late Cretaceous). Geochimica et Cosmochimica Acta,178, 291–306. https://doi.org/10.1016/j.gca.
2015.12.036
Acknowledgments
The authors acknowledge Ian Mounteney
(BGS) for his assistance with the XRD
analyses. SJK, MJL and JBR publish with
the approval of the Executive Director,
British Geological Survey (NERC). The
authors thank Malvern Panalytical for
based XRF analysis, and the Geological
Survey of Denmark and Greenland
(GEUS) for granting access to the E8X
core. This publication contains data
acquired during EM's PhD programme
funded by the College of Engineering,
Mathematics and Physical Sciences,
University of Exeter, and the British
Geological Survey (BGS) under the
programme GW4+, whose support is
gratefully acknowledged. This work was
supported by NERC Isotope Geoscience
Steering Committee (NIGFSC) Grants IP‐
1547‐0515 and IP‐1915‐0619 (to S.K.).
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MARIANI ET AL. 15 of 18
Dickson, A. J., Rees‐Owen, R. L., März, C., Coe, A. L., Cohen, A. S., Pancost, R. D., et al. (2014). The spread of marine anoxia on the northern
Tethys margin during the Paleocene‐Eocene Thermal Maximum. Paleoceanography,29(6), 471–488. https://doi.org/10.1002/2014PA002629
Dunkley Jones, T., Lunt, D. J., Schmidt, D. N., Ridgwell, A., Sluijs, A., Valdes, P. J., & Maslin, M. (2013). Climate model and proxy data
constraints on ocean warming across the Paleocene–Eocene Thermal Maximum. Earth‐Science Review,125, 123–145. https://doi.org/10.1016/
j.earscirev.2013.07.004
Eldrett, J. S., Greenwood, D. R., Polling, M., Brinkhuis, H., & Sluijs, A. (2014). A seasonality trigger for carbon injection at the Paleocene–Eocene
Thermal Maximum. Climate of the Past,10(2), 759–769. https://doi.org/10.5194/cp‐10‐759‐2014
Garel, S., Schnyder, J., Jacob, J., Dupuis, C., Boussafir, M., Le Milbeau, C., et al. (2013). Paleohydrological and paleoenvironmental changes
recorded in terrestrial sediments of the Paleocene‐Eocene boundary (Normandy, France). Palaeogeography, Palaeoclimatology, Palae-
oecology,376, 184–199. https://doi.org/10.1016/j.palaeo.2013.02.035
Goldberg, T., Gordon, G. W., Izon, G., Archer, C., Pearce, C. R., McManus, J., et al. (2013). Resolution of inter‐laboratory discrepancies in Mo
isotope data: An intercalibration. Journal of Analytical Atomic Spectrometry,28(5), 724–735. https://doi.org/10.1039/C3JA30375F
Golonka, J. (2009). Phanerozoic paleoenvironment and paleolithofacies maps: Cenozoic. Geologia /Akademia Górniczo‐Hutnicza im, Stanisława
Staszica w Krakowie,35, 507–587.
Harding, I. C., Charles, A. J., Marshall, J. E. A., Pälike, H., Roberts, A. P., Wilson, P. A., et al. (2011). Sea level and salinity fluctuations during the
Paleocene–Eocene thermal maximum in Arctic Spitsbergen. Earth and Planetary Science Letters,303(1–2), 97–107. https://doi.org/10.1016/j.
epsl.2010.12.043
Hartley, R. A., Roberts, G. G., White, N., & Richardson, C. (2011). Transient convective uplift of an ancient buried landscape. Nature Geoscience,
4(8), 562–565. https://doi.org/10.1038/ngeo1191
Helz, G. R., Bura‐Nakić, E., Mikac, N., & Ciglenečki, I. (2011). New model for molybdenum behavior in euxinic waters. Chemical Geology,
284(3–4), 323–332. https://doi.org/10.1016/j.chemgeo.2011.03.012
Helz, G. R., Miller, C. V., Charnock, J. M., Mosselmans, J. F. W., Pattrick, R. A. D., Garner, C. D., & Vaughan, D. J. (1996). Mechanism of
molybdenum removal from the sea and its concentration in black shales: EXAFS evidence. Geochimica et Cosmochimica Acta,60(19),
3631–3642. https://doi.org/10.1016/0016‐7037(96)00195‐0
Hovikoski, J., Fyhn, M. B. W., Nøhr‐Hansen, H., Hopper, J. R., Andrews, S., Barham, M., et al. (2021). Paleocene‐Eocene volcanic segmentation
of the Norwegian‐Greenland seaway reorganized high‐latitude ocean circulation. Communications Earth & Environment,2(1), 172. https://doi.
org/10.1038/s43247‐021‐00249‐w
Jin, S., Kemp, D. B., Yin, R., Sun, R., Shen, J., Jolley, D. W., et al. (2023). Mercury isotope evidence for protracted North Atlantic magmatism
during the paleocene‐eocene thermal maximum. Earth and Planetary Science Letters V.,602, 117926. https://doi.org/10.1016/j.epsl.2022.
117926
Jones, M. T., Percival, L. M. E., Stokke, E. W., Frieling, J., Mather, T. A., Riber, L., et al. (2019). Mercury anomalies across the Palaeocene–
eocene thermal maximum. Climate of the Past,15(1), 217–236. https://doi.org/10.5194/cp‐15‐217‐2019
Jones, M. T., Stokke, W., Rooney, A. D., Frieling, J., Pogge von Strandmann, P. A. E., Wilson, D. J., et al. (2023). Tracing North Atlantic
volcanism and seaway connectivity across the paleocene–eocene thermal maximum (PETM). Climate of the Past,19(8), 1623–1652. https://
doi.org/10.5194/cp‐19‐1623‐2023
Jones, S. M., Hoggett, M., Greene, S. E., & Dunkley Jones, T. (2019). Large igneous province thermogenic greenhouse gas flux could have
initiated paleocene‐eocene thermal maximum climate change. Nature Communications,10(5547), 1–16. https://doi.org/10.1038/s41467‐019‐
12957‐1
Kemp, S. J., Ellis, M. A., Mounteney, I., & Kender, S. (2016). Palaeoclimatic implications of high‐resolution clay mineral assemblages preceding
and across the onset of the Paleocene‐Eocene thermal maximum, North Sea Basin. Clay Minerals,51(5), 793–813. https://doi.org/10.1180/
claymin.2016.051.5.08
Kender, S., Bogus, K., Pedersen, G. K., Dybkjær, K., Mather, T. A., Mariani, E., et al. (2021). Paleocene/Eocene carbon feedbacks triggered by
volcanic activity. Nature Communications,12(1), 5186. https://doi.org/10.1038/s41467‐021‐25536‐0
Kender, S., Stephenson, M. H., Riding, J. B., Leng, M. J., Knox, R. W. O. ’B., Peck, V. L., et al. (2012). Marine and terrestrial environmental
changes in NW Europe preceding carbon release at the Paleocene‐Eocene transition. Earth and Planetary Science Letters,353–354, 108–120.
https://doi.org/10.1016/j.epsl.2012.08.011
Kirtlan Turner, S., Hull, P. M., Kump, L. R., & Ridgwell, A. (2017). A probabilistic assessment of the rapidity of PETM onset. Nature Com-
munications,8(1), 353. https://doi.org/10.1038/s41467‐017‐00292‐2
Kjennerud, T., & Gillmore, G. K. (2003). Integrated Palaeogene palaeobathymetry of the northern North sea, Pet. Geoscience Series,9(2),
125–132. https://doi.org/10.1144/1354‐079302‐510
Knox, R. W. O. ’B. (1996). Correlation of the early Paleogene in northwest Europe: An overview, Sp. In R. W. O. ’B. Knox, R. M. Corfield, &
R. E. Dunay (Eds.), Correlation of the early Paleogene in northwest Europe (Vol. 101, pp. 1–11). Geological Society Special Publication.
https://doi.org/10.1144/GSL.SP.1996.101.01.01
Lenz, O. K., & Riegel, W. (2001). Isopollen maps as a tool for the reconstruction of a coastal swamp from the Middle Eocene at Helmstedt
(Northern Germany). Facies,45(1), 177–194. https://doi.org/10.1007/BF02668111
Liu, Z., Horton, D. E., Tabor, C., Sageman, B. B., Percival, L. M. E., Gill, B. C., & Selby, D. (2019). Assessing the contributions of Comet impact
and volcanism toward the climate perturbations of the Paleocene‐Eocene thermal maximum. Geophysical Research Letters,46(24),
14798–14806. https://doi.org/10.1029/2019GL084818
Lyons, T. W., & Severmann, S. (2006). A critical look at iron paleoredox proxies: New insights from modern euxinic marine basins. Geochimica
et Cosmochimica Acta,70(23), 5698–5722. https://doi.org/10.1016/j.gca.2006.08.021
Ma, Z., Gray, E., Thomas, E., Murphy, B., Zachos, J., & Paytan, A. (2014). Carbon sequestration during the Palaeocene–Eocene thermal
maximum by an efficient biological pump. Nature Geosciences,7(5), 382–388. https://doi.org/10.1038/ngeo2139
Maclennan, J., & Jones, S. M. (2006). Regional uplift, gas hydrate dissociation and the origins of the Paleocene–Eocene thermal maximum. Earth
and Planetary Science Letters,245(1–2), 65–80. https://doi.org/10.1016/j.epsl.2006.01.069
Mariani, E., Kender, S., Riding, J. B., Leng, M. J., Kemp, S. J., Dybkjær, K., et al. (2024). Geochemical records (d13C, TOC, XRF, SEM, total
sulfur, molybdenum isotopes, clay mineralogy) from marine sediment core E‐8X, central North Sea, across the Paleocene‐Eocene Thermal
maximum. [Dataset]. PANGAEA.https://doi.org/10.1594/PANGAEA.965401
Nielsen, O. B., Rasmussen, E. S., & Thyberg, B. I. (2015). Distribution of clay minerals in the northern North Sea Basin during the Paleogene and
neogene: A result of source‐area geology and sorting processes. Journal of Sedimentary Research,85(6), 562–581. https://doi.org/10.2110/jsr.
2015.40
Paleoceanography and Paleoclimatology
10.1029/2023PA004756
MARIANI ET AL. 16 of 18
Owens, J. D., Reinhard, C. T., Rohrssen, M., Love, G. D., & Lyons, T. W. (2016). Empirical links between trace metal cycling and marine
microbial ecology during a large perturbation to Earth’s carbon cycle. Earth and Planetary Science Letters,449, 407–417. https://doi.org/10.
1016/j.epsl.2016.05.046
Pancost, R., Steart, D., Handley, L., Collinson, M. E., Hooker, J. J., Scott, A., et al. (2007). Increased terrestrial methane cycling at the Palaeocene–
Eocene thermal maximum. Nature,449(7160), 332–335. https://doi.org/10.1038/nature06012
Papadomanolaki, N. M., Sluijs, A., & Slomp, C. P. (2022). Eutrophication and deoxygenation forcing of marginal marine organic carbon burial
during the PETM. Paleoceanography and Paleoclimatology,37(3), e2021PA004232. https://doi.org/10.1029/2021PA004232
Pearce, C. R., Burton, K. W., Pogge van Strandmann, P. A. E., James, R. H., & Gíslason, S. R. (2010). Molybdenum isotope behaviour
accompanying weathering and riverine transport in a basaltic terrain. Earth and Planetary Science Letters,295(1–2), 104–114. https://doi.org/
10.1016/j.epsl.2010.03.032
Penman, D. E., & Zachos, J. C. (2018). New constraints on massive carbon release and recovery processes during the Paleocene‐Eocene Thermal
Maximum. Environmental Research Letters,13(10), 105008. https://doi.org/10.1088/1748‐9326/aae285
Planke, S., Berndt, C., & Alvarez Zarikian, C. A., & the Expedition 396 Scientists. (2023). Mid‐Norwegian margin magmatism and paleoclimate
implications. Proceedings of the International Ocean Discovery Program (Vol. 396). International Ocean Discovery Program. https://doi.org/
10.14379/iodp.proc.396.2023
Pogge von Strandmann, P. A. E., Jones, M. T., West, A. J., Murphy, M. J., Stokke, E. W., Tarbuck, G., et al. (2021). Lithium isotope evidence for
enhanced weathering and erosion during the Paleocene‐Eocene Thermal Maximum. Science Advances,7(42). https://doi.org/10.1126/sciadv.
abh4224
Prøis, B. M. (2015). Late Paleocene – earliest Eocene prograding system in the SW Barents sea, Department of Geosciences (p.76). Universityof
Oslo. Retrieved from https://www.duo.uio.no/handle /10852/45547
Schiøler, P., Andsbjerg, J., Clausen, O. R., Dam, G., Dybkjær, K., Hamberg, L., et al. (2007). Lithostratigraphy of the Palaeogene–lower neogene
succession of the Danish North Sea. Geol. Surv. Den. Green. Bull.,12, 77. https://doi.org/10.34194/geusb.v12.5249
Schoon, P. L., Heilmann‐Clausen, C., Schultz, B. P., Sinninghe Damsté, J. S., & Schouten, S. (2015). Warming and environmental changes in the
eastern North Sea Basin during the Paleocene–Eocene thermal maximum as revealed by biomarker lipids. Organic Geochemistry,78, 79–88.
https://doi.org/10.1016/j.orggeochem.2014.11.003
Secord, R., Gingerish, P. D., Lohmann, K. C., & MacLeod, K. G. (2010). Continental warming preceding the Paleocene‐Eocene thermal
maximum. Nature,467(7318), 955–958. https://doi.org/10.1038/nature09441
Shaw Champion, M. E., White, N. J., Jones, S. M., & Lovell, J. P. B. (2008). Quantifying transient mantle convective uplift: An example from the
Faroe‐Shetland basin. Tectonics,27(1), TC1002. https://doi.org/10.1029/2007TC002106
Sluijs, A., & Brinkhuis, H. (2009). A dynamic climate and ecosystem state during the paleocene–eocene thermal maximum: Inferences from
dinoflagellate cyst assemblages on the New Jersey shelf. Biogeosciences,6(8), 1755–1781. https://doi.org/10.5194/bg‐6‐1755‐2009
Sluijs, A., Brinkhuis, H., Schouten, S., Bohaty, S. M., John, C. M., Zachos, J. C., et al. (2007). Environmental precursors to light carbon input at the
Paleocene/Eocene boundary. Nature,450(7173), 1218–1221. https://doi.org/10.1038/nature06400
Sluijs, A., Röhl, U., Schouten, S., Brumsack, H. J., Sangiorgi, F., Sinninghe Damste, J. S., & Brinkhuis, H. (2008). 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, 1–17. https://doi.org/10.1029/2007PA001495
Sluijs, A., Schouten, S., Pagani, M., Woltering, M., Brinkhuis, H., Sinninghe Damsté, J. S., et al. (2006). Subtropical Arctic Ocean temperatures
during the paleocene/eocene thermal maximum. Nature,441(7093), 610–613. https://doi.org/10.1038/nature04668
Stokke, E. W., Jones, M. T., Riber, L., Haflidason, H., Midtkandal, I., Schultz, B. P., & Svensen, H. H. (2021). Rapid and sustained environmental
responses to global warming: The paleocene–eocene thermal maximum in the eastern North Sea. Climate of the Past,17(5), 1989–2013. https://
doi.org/10.5194/cp‐17‐1989‐2021
Stokke, E. W., Jones, M. T., Tierney, J. E., Svensen, H. H., & Whiteside, J. H. (2020). Temperature changes across the paleocene‐eocene thermal
maximum – A new high‐resolution TEX
86
temperature record from the eastern North Sea basin. Earth and Planetary Science Letters,544,
116388. https://doi.org/10.1016/j.epsl.2020.116388
Storey, M., Duncan, R. A., & SwisherIII, C. C. (2007). Paleocene‐Eocene thermal maximum and the opening of the Northeast Atlantic. Science,
316(5824), 587–589. https://doi.org/10.1126/science.1135274
Svensen, H., Planke, S., Malthe‐Sørenssen, A., Jamtveit, B., Myklebust, R., Rasmussen Eidem, T., & Rey, S. S. (2004). Release of methane from a
volcanic basin as a mechanism for initial Eocene global warming. Nature,429(6991), 542–545. https://doi.org/10.1038/nature02566
Thomas, E. (1998). The biogeography of the late Paleocene benthic foraminiferal extinction. In M. P. Aubry, S. Lucas, & W. A. Berggren (Eds.),
Late Paleocene‐early Eocene biotic and climatic events in the marine and terrestrial records (pp. 214–243). Columbia University Press.
Thomas, E. (2007). Cenozoic mass extinctions in the deep sea: What perturbs the largest habitat on Earth? In S. Monechi, R. Coccioni, & M. R.
Rampino (Eds.), Large ecosystem perturbations: Causes and consequences (Vol. 424, pp. 1–23). Geological Society of America. https://doi.
org/10.1130/2007.2424(01)
Tremblin, M., Khozyem, H., Adatte, T., Spangenberg, J. E., Fillon, C., Grauls, A., et al. (2022). Mercury enrichments of the Pyrenean foreland
basins sediments support enhanced volcanism during the Paleocene‐Eocene thermal maximum (PETM). Global and Planetary Change,212,
103794. https://doi.org/10.1016/j.gloplacha.2022.103794
Tribovillard, N., Algeo, T. J., Lyons, T., & Riboulleau, A. (2006). Trace metals as paleoredox an paleoproductivity proxies: An update. Chemical
Geology,232(1–2), 12–32. https://doi.org/10.1016/j.chemgeo.2006.02.012
Vickers, M. L., Jones, M. T., Longman, J., Evans, D., Ullmann, C. V., Stokke, E. W., & IODP Expedition 396 Scientists. (2024). Paleocene–
Eocene age glendonites from the Mid‐Norwegian Margin – Indicators of cold snaps in the hothouse? Climate of the Past,20(1), 1–23.
https://doi.org/10.5194/cp‐20‐1‐2024
Wilkin, R. T., Arthur, M. A., & Dean, W. E. (1997). History of water‐column anoxia in the Black Sea indicated pyrite framboid size distributions.
Earth and Planetary Science Letters,148, 17–525. https://doi.org/10.1016/S0012‐821X(97)00053‐8
Wilkinson, C. M., Ganerød, M., Hendriks, B. W. H., & Eide, E. A. (2017). Compilation and appraisal of geochronological data from the North
Atlantic igneous Province (NAIP) (Vol. 447, pp. 69–103). Geological Society, London, Special Publications. https://doi.org/10.1144/SP447.10
Zacke, A., Voigt, S., Joachimski, M. M., Gale, A. S., Ward, D. J., & Tütken, T. (2009). Surface‐water freshening and high‐latitude river discharge
in the Eocene North Sea. Journal of the Geological Society,166(5), 969–980. https://doi.org/10.1144/0016‐76492008‐068
Paleoceanography and Paleoclimatology
10.1029/2023PA004756
MARIANI ET AL. 17 of 18
References From the Supporting Information
Deegan, C. E., & Scull, B. J. (1977). A standard lithostratigraphic nomenclature for the central and northern North Sea. Represent Institute of
Geological Science,77/25, 36.
Horni, A. J., Hopper, J. R., Blischke, A., Geisler, W. H., Stewart, M., McDermott, K., et al. (2017). Regional distribution of volcanism within the
North Atlantic Igneous Province. Geological Society, London, Special Publications,447(1), 105–125. https://doi.org/10.1144/sp447.18
Jones, S. M., & White, N. (2003). Shape and size of the starting Iceland plume swell. Earth and Planetary Science Letters,216(3), 271–282.
https://doi.org/10.1016/S0012‐821X(03)00507‐7
Meyers, S. R., Sageman, B. B., & Lyons, T. W. (2005). Organic carbon burial rate and the molybdenum proxy: Theoretical framework and
application to Cenomanian‐Turonian oceanic anoxic event 2. Paleoceanography,20(2), 169–189. https://doi.org/10.1029/2004PA001068
Nadin, P., Kusznir, N., & Cheadle, M. (1997). Early tertiary plume uplift of the North Sea and Faroe–Shetland basins. Earth and Planetary
Science Letters,148(1–2), 109–127. https://doi.org/10.1016/S0012‐821X(97)00035‐6
Nakagawa, Y., Takano, S., Firdaus, M. L., Norisuye, K., Hirata, T., Vance, D., & Sohrin, Y.(2012). The molybdenum isotopic composition of the
modern ocean. Geochemistry Journal,46(2), 131–141. https://doi.org/10.2343/geochemj.1.0158
Smallwood, J. R., & Jill, C. E. (2002). The rise and fall of the Faroe–Shetland basin: Evidence from seismic mapping of the Balder formation.
Journal of the Geological Society,159(6), 627–630. https://doi.org/10.1144/0016‐764902‐064
Thiry, M. (2000). Palaeoclimatic interpretation of clay minerals in marine deposits: An outlook from the continental origin. Earth‐Science Re-
views,49(1–4), 201–221. https://doi.org/10.1016/S0012‐8252(99)00054‐9
Paleoceanography and Paleoclimatology
10.1029/2023PA004756
MARIANI ET AL. 18 of 18
