PUBLICATIONS
Geochemistry, Geophysics, Geosystems
RESEARCH ARTICLE
10.1002/2013GC005074
Key Points:
We measure concentrations of Mn
and U in marine PETM sediments
North Atlantic bottom waters were
suboxic relative to the Pacific during
the PETM
The source of methane release
during the PETM was likely in the
North Atlantic
Deep-sea redox across the Paleocene-Eocene thermal
maximum
€like1,2,3, Margaret L. Delaney4, and James C. Zachos5
Cecily Pa
1
Institute of Geosciences, Goethe-University Frankfurt, Frankfurt, Germany, 2Biodiversity and Climate Research Centre
(BIK-F), Frankfurt, Germany, 3Now at MARUM—Center for Marine Environmental Sciences, University of Bremen, Bremen,
Germany, 4Ocean Sciences Department, Institute of Marine Sciences, University of California, Santa Cruz, California, USA,
5
Earth and Planetary Sciences Department, University of California, Santa Cruz, California, USA
Abstract Large amounts of 13C-depleted carbon were released to the oceans and atmosphere during a
Supporting Information:
Read_me
Figure S1
Table S1–S15
Correspondence to:
C. P€alike,
[email protected]
Citation:
P€alike, C., M. L. Delaney, and J. C.
Zachos (2014), Deep-sea redox across
the Paleocene-Eocene thermal
maximum, Geochem. Geophys. Geosyst.,
15, 1038–1053, doi:10.1002/
2013GC005074.
period of abrupt global warming at the Paleocene-Eocene thermal maximum (PETM) (55 Ma). Investigations of qualitative sedimentologic and paleontologic redox proxies such as bioturbation and benthic
assemblages from pelagic and hemipelagic sections suggest transient reductions in bottom water oxygen
during this interval, possibly on a global scale. Here, we present bulk sediment manganese (Mn) and uranium (U) enrichment factors (EF) in Atlantic and Pacific deep-sea cores to constrain relative paleoredox
changes across the PETM. Mn EF range from 1 to 9 in Atlantic sites, 1 to 35 in Southern Ocean sites, and are
at crustal averages (EF 5 1) in Pacific sites. U EF range from 1 to 5 in Atlantic sites, 1 to 90 in Southern Ocean
sites, and are at crustal averages in Pacific sites. Our results indicate suboxic conditions prior to, during, and
in the recovery from the PETM at intermediate depth sites in the Atlantic and Southern Ocean while the
Pacific sites remained relatively oxygenated. The difference in oxygenation between the Atlantic and Pacific
sites leads us to suggest the source for isotopically light carbon release during the PETM was in the Atlantic.
1. Introduction
Received 2 OCT 2013
Accepted 16 FEB 2014
Accepted article online 24 FEB 2014
Published online 11 APR 2014
The Paleocene-Eocene thermal maximum (PETM) is characterized by extreme warming coupled with a
major perturbation to the global carbon cycle. The ocean’s surface temperature increased by 5–7 C, in less
than 20 kyr [Sluijs et al., 2006; Zachos et al., 2003]. This event was the first and largest of a series of hyperthermal events in the early Eocene, identified by sharp negative d13C excursions [Lourens et al., 2005; Stap
et al., 2009; Zachos et al., 2010]. Several mechanisms have been posited to explain the globally synchronous
3& negative carbon isotope excursion (CIE) and implied carbon release: the rapid dissociation of methane
release from gas hydrates [Dickens, 2000, 2011], volcanic release of CH4 and CO2 in the North Atlantic [Svensen et al., 2004], and various terrestrial sources of carbon [DeConto et al., 2012; Higgins and Schrag, 2006;
Kurtz et al., 2003; Pancost et al., 2007].
The posited release of a large mass of reduced carbon, such as methane, directly into the ocean
coupled with rapid or even slow warming of the upper ocean could alter processes critical to maintaining the dissolved oxygen content of the deep sea. This includes changes in the rate and style of meridional overturning circulation (MOC), which would be highly sensitive to high-latitude warming [Lunt
et al., 2010]. Indeed, multiple indicators or proxies of bottom water redox conditions and model studies
suggest shifts toward lower redox state of bottom waters in multiple locations throughout the oceans
[Dickson et al., 2012; Nicolo et al., 2010; Winguth et al., 2012]. For example, benthic foraminifer assemblages show reduction in the abundance of species that were intolerant of suboxic conditions [Thomas,
2007]. Documenting the temporal and spatial patterns of redox change during the PETM is thus critical
to testing models for the origin of the event, as well as for assessing the sensitivity of ocean overturning
to rapid warming in general.
Previous studies have either used dissolution patterns of seafloor calcium carbonate (CaCO3) records along
with sediment models to evaluate the pattern and location of carbon release through the ocean basins
[Panchuk et al., 2008; Zeebe and Zachos, 2007] or benthic foraminifera carbon isotope records [Nunes and
Norris, 2006]. We previously documented changes in the distribution of redox-sensitive trace metals in a
depth transect of Paleocene-Eocene Boundary (PEB) sections from Walvis Ridge, in the south Atlantic,
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1038
Geochemistry, Geophysics, Geosystems
180˚ 210˚ 240˚ 270˚ 300˚ 330˚
90˚
0˚
30˚
60˚
10.1002/2013GC005074
90˚
120˚ 150˚ 180˚
90˚
60˚
60˚
401
30˚
30˚
1001
999 1258
0˚
1209
0˚
1221
-30˚
1266
1262
1263
-30˚
-60˚
738
690
-90˚
180˚ 210˚ 240˚ 270˚ 300˚ 330˚
0˚
30˚
60˚
90˚
-60˚
-90˚
120˚ 150˚ 180˚
55 Ma Reconstruction
Figure 1. Plate tectonic reconstruction for 55 Ma showing the location of the study sites. This map was generated by the Ocean Drilling
Stratigraphic Network (http://www.odsn.de/).
finding evidence of a transient shift toward suboxic conditions [Chun et al., 2010]. Our current study broadens the spatial patterns of change in the same redox-sensitive trace metals, specifically Mn and U. One
objective is to use these patterns to assess the methane dissociation hypothesis. If CH4 is oxidized in the
oceans as posited, this should produce a pattern of oxygen deficiency in the deep oceans that is partly
determined by the flow of carbon and ocean circulation [Dickens, 2000, 2003]. We use our proxies for bottom water redox as one test of the methane hydrate hypothesis. While alternate sources of carbon release
to the ocean and atmosphere produce similar d13C and ocean temperature records, warming alone is insufficient for major changes in dissolved oxygen [Dickens, 2000].
In the modern ocean, dissolved inorganic carbon (DIC) and bottom water oxygen are influenced in opposite
ways along the path of deepwater circulation. The youngest deep waters are oxygen rich and DIC poor.
Along the pathway of deep-ocean circulation, oxygen concentrations decrease as organic matter oxidation
occurs, increasing DIC concentrations. The increase in DIC results in deep waters that are more corrosive to
CaCO3. Depletion of bottom water oxygen should be most pronounced in the reservoir where methane oxidation occurred [Dickens, 2000]. Methane oxidation in the water column will also cause DIC values to
increase, causing deep waters to appear temporarily older.
In this study, we test the theory that thermal dissociation of marine gas hydrates contributed to rapid climate change at the PETM. We also aim to determine the geographic location of this gas hydrate reservoir.
Our results are from a wide distribution of sites: two in the North Atlantic, two from the Southern Ocean,
and two from the Pacific Ocean (Figure 1 and Table 1). Redox-sensitive trace metal enrichment factors are
measured to infer the paleoredox conditions at time of sediment burial. We compare our results to other
proxies of bottom water carbon chemistry and oxygenation across the PETM.
2. Background
2.1. Proxies
Trace element enrichment factors of manganese (Mn) and uranium (U) are used to characterize the paleoredox
environment of the host sediments at or near the time of burial [Calvert and Pedersen, 1993]. Because of the
complementary redox sensitivity of Mn and U in their solid and dissolved states, we use bulk trace element
ratios to determine the position of the oxic/suboxic boundary [Mangini et al., 2001]. However, early diagenesis of
subsurface sediments has the potential to remobilize primary redox signals [Froelich et al., 1979; Tribovillard et al.,
2006]. We use modern day redox state to assess the likelihood of postdepositional diagenesis.
In general, Mn enrichments relative to crustal abundances represent Mn oxides (MnO2) indicating welloxygenated conditions [Calvert and Pedersen, 1993]. However, Mn enrichments may also represent
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1039
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
Table 1. Site Characteristics
Site (Leg)
Bay of
Biscaya
401 (48)
Demerara
Riseb
1258 (207)
Walvis Ridgec
1262 (208)
1263 (208)
1266 (208)
Maud Rised
690 (113)
Kerguelen
Plateaue
738 (119)
Equatorial
Pacificf
1221 (199)
Shatsky
Riseg
1209 (198)
Holes used in this study
NA
A
B
A/C/D
B/C
B
C
C
B
Water depth
2495
3192
4755
2717
3798
2914
2252
5174
2387
Paleodepth
2000
2500
3600
1500
2600
2100
1350
3300
2400
9 26N
27 11S
28 32S
28 33S
65 09S
62 43S
12 02N
32 39N
Latitude
47 26N
54 44W
1 35E
2 47E
2 21E
1 12E
82 47E
143 42W
158 30E
Longitude
8 49W
Depth range of samples used
198.54–202.98 170.55–174.99 124.40–128.58 282.53–288.73 260.00–272.49 166.90–171.73 283.54–286.30 153.40–154.80 195.10–197.21
in this study (mbsf)
Sulfate minimum (mM)
24.3
0.7
21.19
23.32
19.13
20.1
22.8
29.2
23.5
Depth to sulfate minimum (mbsf)
220
399.25
140.4
9.75
229
238.85
57.45
80.4
247.15
Sulfate at depth of PETM (estimate)
24
15
22
24
22
22
22.8
NA
24
a
Ellis et al. [1979] and Montadert et al. [1979].
Erbacher et al. [2004] and Sexton et al. [2006].
c
Zachos et al. [2004, 2005].
d
Barker and Kennett [1988] and Kennett and Stott [1990].
e
Barron et al. [1989] and Barerra and Huber [1991].
f
Lyle et al. [2002].
g
Bralower et al. [2002].
b
authigenic Mn carbonates (MnCO3) that precipitate within suboxic pore waters [Boyle, 1983; Gingele and
Kasten, 1994; Pedersen and Price, 1982]. We use a reductive cleaning procedure to operationally differentiate
between MnO2 (oxic environment) and MnCO3 (suboxic environment) [Boyle, 1983]. In conjunction with Mn
EF, U enrichments above crustal averages represent uraninite (UO2), reflecting suboxic conditions [Calvert
and Pedersen, 1993; Morford and Emerson, 1999]. Although these proxies have not been directly calibrated
to oxygen concentrations, they can still provide useful information about the qualitative redox conditions
with time and among sites (see Chun et al. [2010] for more details).
2.2. Site Selection
This study expands on our previous work in the South Atlantic to include sites from the North Atlantic,
Southern Ocean, and Pacific Ocean. We selected six Deep Sea Drilling Program (DSDP) and Ocean Drilling
Program (ODP) sites where the PETM has been identified (Figure 1 and Table 1). The sites range in paleowater depth from 1350 m on Kerguelen Plateau to 3300 m in the Equatorial Pacific, thus providing constraints on intermediate to deepwater characteristics. Burial depths of late Paleocene age samples from
these sites range from 150 meters below seafloor (mbsf) to 280 mbsf, sufficiently deep that convective
flow through the sediments with oceanic crust has ceased [McDuff, 1981].
2.3. Age Model
€hl et al., 2000, 2007].
Orbitally tuned age models exist for the PETM sections at ODP Sites 690 and 1263 [Ro
These age models are based on the identification of precession and eccentricity cycles in sediment lithology, Fe, and Ca; and they assign ages relative to the PETM to inflection points in the bulk calcite d13C. We
adopted the published tie points by Kaiho et al. [2006] for Site 1209 and by Paytan et al. [2007] for Site
€hl et al. [2007]. We visually correlated all other sites to the orbitally tuned
1221, reassigning the ages to Ro
age model using the major inflection points in bulk calcite d13C (Figure 2 and Table 2). The convention of
subdividing the PETM into 50 kyr ‘‘preevent,’’ 80 kyr ‘‘core-CIE,’’ and 80–170 kyr ‘‘recovery’’ intervals is based
on the carbon isotope records and adapted from Chun et al. [2010].
To constrain time outside of the PETM interval, we used magnetostratigraphic and biostratigraphic datums
where available with ages calculated by Westerhold et al. [2007] which are consistent with the time frame of
€hl et al. [2007]. For the oldest age pick, we used the boundary between magnetothe PETM as defined in Ro
chron 24 reversal and magnetochron 25 normal (C24r/C25n) representing 1110 kyr prior to the onset of the
PEB [Westerhold et al., 2007]. For an upper boundary past the recovery of the PETM, we used the base of
planktonic foraminifera Zone P6 as defined by the last occurrence of Morozovella velascoensis as 1743.60
kyr relative to the onset of the PETM [Zachos et al., 2004]. In all our sites, we assumed linear sedimentation
rates between tie points for age assignments.
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1040
Geochemistry, Geophysics, Geosystems
1263
δ C (‰, vPDB)
10.1002/2013GC005074
1209
δ C (‰, vPDB)
13
13
-1 0 1 2 3
-1 0 1 2 3
332
194
210.5
164
333
195
210.7
Depth (meters)
166
G
196
334
211.0
168
F
197
335
170
211.2
C
A
198
?
336
A-
?
211.4
172
199
-1 0 1 2 3
13
δ C (‰, vPDB)
1258
-1 0 1 2 3
13
δ C (‰, vPDB)
690
Figure 2. Bulk sediment carbon isotope records for Sites 1258, 1263, 690, and 1209 versus depth. Letters and associated lines indicating
lines of correlation are based on inflections in the carbon isotopes after Zachos et al. [2005]. Bulk isotope data for Site 1258 are out of range
of the plot (this study), 1263 from Zachos et al. [2005], Site 690 from Bains et al. [1999], Site 1209 from Colossimo et al. [2006]. vPDB, Vienna
PeeDee Belemnite; Sites 1258 and 1209 given in meters composite depth (MCD); Site 690 given in meters below seafloor (MBSF); and Site
1263 given in revised meters composite depth (RMCD).
Where bulk d13C values were not available (Site 401) we used benthic foraminifera stable carbon isotopes.
Recent work has shown that the benthic foraminifera stable isotope records display a more abrupt CIE than
bulk carbonate records [McCarren et al., 2008]. Therefore, using d13C values from benthic foraminifera species (such as Nuttallides truempyi) that were not affected by the extinction event is the most suitable
replacement for bulk carbonate d13C records.
The benthic foraminifera carbon isotope record of Nuttallides truempyi from Site 401 [Nunes and Norris,
2006] was correlated to the benthic records from Sites 1262, 1263, and 690 [Kelly et al., 2005; Kennett and
€hl et al. [2007] (Figure 3 and Table 2). To
Stott, 1991; McCarren et al., 2008] using the assigned ages from Ro
assign ages to depths below the onset of the PETM at Site 401, we used the sedimentation rate calculated
by Nunes and Norris [2006] for this portion of their age model, and extrapolated ages from the onset of the
event.
2.4. Samples and Analytical Procedures
Sediment samples of 1 cm3 volume were taken from all sites for trace metal analysis. Sampling resolution
varied among the sites and ranged from 1 sample every 1.5 cm (1 sample/1.6 kyr) at ODP 738 to 1
sample every 13 cm (1 sample/13 kyr) at ODP 1258.
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1041
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
Table 2. Carbon Isotope Tie Points From ODP Site 690 and Assigned Ages Used for Correlation and Dating the Corresponding P-E Boundary Sectionsa
Depth
Tie Points
Site 690
(mbsf)
Time Relative to
PETM CIE (6kyr)
[R€
ohl et al., 2007]
LO Morozovella velascoensis
H
G
F
E
D
C
B
A
PEB
ABC24rC25n boundary
318.08b
166.13
167.12
169.05
169.39
169.56
170.02
170.33
170.63
170.64
171.24
172.81
153.11c
743.60
196.87
153.5
94.23
81.17
71.25
42.38
21.9
0.75
0
230.8
2102.53
21110.00
Site 401 (mbsf)
Site 1258 (mcd)
199.79
192.21
200.72
201.48
Site 738
(mbsf)
Site 1209 (mcd)
[Kaiho et al., 2006]
Site 1221 (mbsf)
[Paytan et al., 2007]
208.80
195.47
196.38
202.58
283.78
284.30
284.62
153.40
211.04
285.16
211.22
285.48
198.15
286.02
211.26
211.31
153.90
154.00
154.10
154.30
154.40
216.28
a
mbsf 5 meters below seafloor and mcd 5 meters composite depth.
Depth for LO Morozovella velascoensis from Site 1263 in revised meters composite depth (rmcd).
c
for C23r 25n boundary from 1262 in mcd.
b
Bulk sediment samples were first freeze dried, crushed, and sieved with a 150 lm polypropylene mesh.
Total sediment digestions and reductive cleaning of sample splits were processed in the same way as
detailed in Chun et al. [2010]. Total concentrations of manganese (Mn), titanium (Ti), and uranium (U) were
measured on a Finnigan Element high-resolution inductively coupled plasma-mass spectrometer (HR-ICPMS). Aluminum (Al) was measured on a Perkin-Elmer Optima 4300 DV inductively coupled plasma-optical
emission spectrometer (ICP-OES). All digestions and analyses took place at the University of California, Santa
Cruz, USA.
250
401
200
Depth (meters)
200
G
150
?
201
100
F
Time relative to PEB (kyr)
199
50
202
0
PEB
203
1262 benthic
1263 benthic
690 benthic
1263 bulk
-50
-3 -2 -1 0 1 2 3 -3 -2 -1 0 1 2 3
13
δ C (‰, vPDB)
13
δ C (‰, vPDB)
Figure 3. Benthic foraminiferal isotope correlation of Site 401 to Sites 1262,
1263, and 690 versus depth (meters below seafloor). Benthic foraminiferal isotope records from Site 401 Nunes and Norris [2006], Sites 1262 and 1263 from
McCarren et al. [2008], Site 690 record from Kennett and Stott [1991] and Kelley
et al. [2005]. Bulk isotope record from Site 1263, shown for comparison, from
Zachos et al. [2005]. Depth of section for Site 401 is plotted on the left axis and
relative time before and after the PEB (Paleocene-Eocene Boundary) on the
right axis. Ages for Sites 1262, 1263, and 690 from R€
ohl et al. [2007]. vPDB,
Vienna PeeDee Belemnite.
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
Two internal standards composed of a
homogenized mixture of ODP Leg 202
and ODP Leg 208 samples were processed and analyzed with each set of
digestions to evaluate reproducibility.
Mean concentrations 6 1 standard deviation of the ODP Leg 202 consistency
standard were 133 6 23 lmol g21 for Al,
2.4 6 0.4 lmol g21 for Ti, 6.1 6 0.7 lmol
g21 for Mn, and 1.1 6 0.3 nmol g21 for U
(Table 3). Mean concentrations 6 1
standard deviation of the ODP Leg 208
consistency standard were 154 6 22
lmol g21 for Al, 4.7 6 0.8 lmol g21 for
Ti, 4.8 6 0.6 lmol g21 for Mn, and
0.30 6 0.2 nmol g21 for U (Table 3).
For sites that did not have published
records of weight percent (wt %) CaCO3
or bulk stable carbon isotopes, we measured them in this study. We measured
wt % CaCO3 on sediments from Site 401
using a UIC, Inc. Coulometrics Model
5012 CO2 coulometer at the University
of California, Santa Cruz, USA. Relative
standard deviations of the mean of multiple measures of a pure CaCO3 standard
were <1%. Weight % CaCO3
1042
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
Table 3. Analytical Figures of Merit
Element Concentration (lmol/g sediment)
b
Detection limits
Reproducibility Leg 202 (consistency standard replicates)c
Reproducibility Leg 208 (consistency standard replicates)c
Al
Ti
Mn
Ua
18.52
132.98 6 23
154.26 6 22
0.32
2.4 6 0.4
4.71 6 0.81
0.78
6.1 6 0.7
4.82 6 0.55
0.44
1.09 6 0.28
0.30 6 0.15
a
nmol/g sediment.
Defined as three times the standard deviation of replicate measures of a blank of the same matrix as each digestion, and expressed
in equivalent concentration for a typical size sediment sample.
c
Reproducibility defined as 61 standard deviation of multiple replicates of a solid sample included in each analytical run (n 5 14).
b
measurements for Site 1258 were made simultaneously during stable isotope measurements on the VG
Micromass Optima ratio mass spectrometer with precision of 66%.
Bulk stable carbon isotopes were measured on 33 homogenized and crushed sediment samples from ODP
Site 1258. Analyses were performed on the VG Micromass Optima ratio mass spectrometer at the University
of California, Santa Cruz, USA. Results are reported in & relative to Vienna PeeDee Belemnite (vPDB) and
calibrated via NBS-19. Precision of d13C measurements is better than 0.1&.
3. Results
The results of bulk sample elemental analyses are presented in supporting information Tables S1–S15 and
Figure S1. Average bulk Mn concentrations range from 3 lmol g21 at Site 738 to 171 lmol g21 at Site
1221. Average U concentrations range from 0.5 nmol g21 at Site 738 to 5 nmol g21 at Site 1258. We normalize our results to Ti to isolate the enrichment over crustal values and to provide downcore comparisons
among the sites. Trace metal enrichment factors (EF) for Mn and U were calculated as EF 5 (metal/Ti)sample/
(metal/Ti)crust using the bulk crustal ratios Mn/Ti (mol/mol) 5 0.156 and U/Ti (mmol/mol) 5 0.061 [Rudnick
and Gao, 2003].
Mn enrichment factors exhibit variation across the PETM. At the Atlantic and Southern Ocean Sites 401,
1258, 690, and 738, Mn EF for bulk sample digestions have average values between 2 and 6 preevent and
generally decrease during the core of the event to average values between 1 and 4 (Figure 4). Then in the
carbon isotope recovery, Mn EF for bulk sample digestions are at average values between 2 and 20. Site
690 displays a gradual peak with increasing values from 95 kyr from onset of the PEB to a peak value of
37 at 106 kyr, then a gradual decline. At the Pacific Sites 1221 and 1209, Mn EF for bulk sample digestions
average between 2 and 13 preevent, 21 and 47 during the core of the event, and 15 and 17 in the recovery.
Site 1221 exhibits a sharp peak in Mn EF at 20 kyr to values of 350, lasting 4 kyr.
We compare Mn EF from bulk sediment digested before and after a reductive cleaning procedure to determine the presence or absence of Mn oxides [Chun et al., 2010]. When Mn EF are >1 in bulk sediment before
the reductive cleaning procedure and are then at crustal averages (Mn 1) after the cleaning and there is
no U enrichment, the difference is interpreted to be Mn oxides representing an oxygenated environment.
When Mn EF are >1 in bulk sediment before the reductive cleaning and remain relatively unchanged after
the cleaning, this represents Mn carbonates (MnCO3), indicating an early diagenetic environment of suboxic
pore waters.
At the Atlantic and Southern Ocean Sites 401, 1258, and 738 Mn EF on the subset of samples that were
reductively cleaned prior to total digestion show little change compared to Mn EF before reduction. Site
738 displays Mn EF values from reductively cleaned samples that are slightly higher than the untreated bulk
samples. We suspect that this was due to sample loss during the reductive cleaning procedure as we did
not weigh samples again after the cleaning and prior to the sediment digestion. Average values at these
sites are between 1 and 7 before and after reduction. At Site 690, the gradual peak in Mn EF remains
unchanged in samples after reductive cleaning. Mn EF values postreductive cleaning at Site 690 are from
10 to 30 between 90 and 150 kyr relative to the PEB. At the two Pacific Sites 1221 and 1209, Mn/Ti on
the reductively cleaned samples are reduced to crustal averages throughout the sample section. At Site
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1043
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
Figure 4. Manganese enrichment factors (Mn EF) calculated from measured Mn concentrations and the crustal molar Mn/Ti of 0.156 from
Rudnick and Gao [2003], plotted versus relative time before and after the PEB (Paleocene-Eocene Boundary) [R€
ohl et al., 2007]. Scales for
Mn EF vary for each site. Site number is located at the top of each plot. Sites 1262, 1263, and 1266 are from Chun et al. [2010]. Dashed vertical line is EF 5 1 (no enrichment or depletion relative to presumed crustal source). Closed symbols indicate EF for bulk digests without
reductive cleaning. Open symbols indicate EF postreductive cleaning. Reductively cleaned samples are at lower resolution than the bulk
digests. Gray shaded region indicates duration of core-carbon isotope excursion (CIE) 80 kyr.
1221, the sharp Mn EF peak at 20 kyr is reduced to near crustal averages, Mn EF 3, postreductive
cleaning.
U enrichment factors remain at or near crustal averages (U EF 2) at all sites during the PETM, with the
exception of Site 690 (Figure 5). U EF are near crustal averages preevent and during the core at Site 690.
Then, starting at 90 kyr from the onset of the PEB, U EF increase to a gradual peak of 90 at 106 kyr, coeval with the Mn EF peak, then return to crustal averages.
As a part of this study, we also measured CaCO3 concentrations for Sites 401 and 1258. CaCO3 concentrations for Site 401 range between 56 and 68 wt % preevent (Figure 6). At the onset of the PETM, values drop
to 45 wt % and remain in the range of 42–48 wt %. Just prior to the recovery interval, CaCO3 concentrations
decrease to 28 wt % and remain between 30 and 40 wt % throughout the recovery. CaCO3 concentrations
for Site 1258 average 33 wt % pre-CIE, then dropped to values of <1 wt % during the core of the event, and
gradually return to values between 30 and 55 wt % in the recovery (Figure 6). Throughout the record,
CaCO3 concentrations remain <55 wt %. We also generated bulk carbon and oxygen isotope data for Site
1258. Bulk d13C values measured at Site 1258 average 1.58& pre-CIE, then decrease to values as low as
26.3& at the base of the clay layer. During the recovery, d13C values average 1.02& (Figure 7).
4. Discussion
4.1. Paleoredox Changes
Comparing Mn EF before and after a reductive cleaning procedure, combined with U EF measurements,
shows paleoredox changes recorded in marine sediments across a global set of sites that record the PETM
(Table 4). In the following, we evaluate how the geographic response of each site contributes to our understanding of how the oceans responded to this abrupt global warming event.
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1044
Geochemistry, Geophysics, Geosystems
Time relative to PEB (kyr)
401
1258
0 25 50 75100
1262
1263
1266
0 2 4 5 7
250
690
738
1221
1209
200
200
150
150
100
100
50
50
0
0
-50
0 2 4 5 7
Time relative to PEB (kyr)
0 2 4 5 7
250
10.1002/2013GC005074
-50
0 2 4 5 7
0 2 4 5 7
0 2 4 5 7
Uranium Enrichment Factor
Figure 5. Uranium enrichment factors (U EF) calculated from measured U concentrations and the crustal molar U/Ti of 0.061 from Rudnick
and Gao [2003], plotted versus relative time before and after the PEB (Paleocene-Eocene Boundary) [R€
ohl et al., 2007]. Site number is
located at the top of each plot. Sites 1262, 1263, and 1266 are from Chun et al. [2010]. Dashed vertical line is EF 5 1 (no enrichment or
depletion relative to presumed crustal source). Gray shaded region indicates duration of core-carbon isotope excursion (CIE) 80 kyr.
Figure 6. Bulk sediment weight % carbonate content plotted versus time (kyr) relative to the Paleocene-Eocene Boundary (PEB) using the
age model of R€
ohl et al. [2007]. Site number is located at the top of each plot. Site 401 (this study), Site 1258 (this study), Sites 1262, 1263,
and 1266 from Zachos et al. [2005], Site 690 from Farley and Eltgroth [2003] (smoothed line); Kelly et al. [2005] (open triangles), and this
study (black X). Site 1221 from Murphy et al. [2006] and Site 1209 from Colossimo et al. [2005]. Note change in axis scale for Sites 690 and
1209. Gray shaded region indicates duration of core-carbon isotope excursion (CIE) 80 kyr.
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1045
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
We have previously demonstrated
the utility of these proxies to provide a paleoredox history on a
depth transect on Walvis Ridge
during the PETM [Chun et al., 2010].
We found that prior to the
Paleocene-Eocene Boundary (PEB),
sites deeper than 2500 m paleowater depth were oxygenated
while the shallow 1500 m site was
suboxic. During the core-CIE, the
oxygen minimum zone expanded
to include our deepest site (3600 m
paleowater depth), then returned
to preevent conditions during the
recovery.
In the modern oceans, methane
hydrates are predominately stored
on continental margins at intermediate water depths (500–1000 m).
Figure 7. Bulk carbon isotope records for Site 1258 plotted versus depth (meters comDSDP Site 401 in the Bay of Biscay
posite depth) and time (kyr) relative to the Paleocene-Eocene Boundary (PEB) using
the age model of R€
ohl et al. [2007]. vPDB, Vienna PeeDee Belemnite.
represents the northernmost Atlantic end-member in our global array
of PETM sites. We selected intermediate water depth sites from the Atlantic basin to test the theory of methane hydrate dissociation in the
North Atlantic [Dickens et al., 1995; Dickens, 2011]. Prior to the CIE, Mn EF at Site 401 are >1 and show no
change with reductive treatment, indicating the presence of Mn-carbonates (Figure 4). During the core of
the CIE and recovery intervals, Mn EF values average <2 while U EF is 2 (Figure 5) throughout the interval
suggesting that Site 401 pore waters became more reducing and Mn-carbonates were no longer produced.
Thus, inferring relatively lower oxygen concentrations in the overlying deep water. Our results are in agreement with planktic foraminiferal and ostracode assemblage studies across the PETM at this site, which
record an increase in low oxygen tolerant taxa in the core of the CIE [Pardo et al., 1997; Yamaguchi and Norris, 2012].
Calcium carbonate at Site 401 displayed a unique pattern not measured at the other PETM sites (Figure 6).
The initial decrease in calcium carbonate at the onset of the CIE was expected as the global calcite compensation depth (CCD) shoaled in response to rapid acidification of the oceans due to the release of carbon to
Table 4. Summary of Proxy Results (Mean Values)
Location
Site (Leg)
Walvis Ridgec
Bay of Demerara
Maud Rised
Riseb
Biscaya
401 (48) 1258 (207) 1262 (208) 1263 (208) 1266 (208) 690 (113)
Pre-CIE (250 to 0 kyr)
Mn EF before reduction
5.81
Mn EF after reduction
6.31
U EF
1.62
Redox state of bottom waters Suboxic
Core-CIE (0–80 kyr)
Mn EF before reduction
1.84
Mn EF after reduction
2.52
U EF
1.53
Redox state of bottom waters Suboxic
Recovery (80–170 kyr)
Mn EF before reduction
1.58
Mn EF after reduction
1.17
U EF
1.22
Redox state of bottom waters Suboxic
€
PALIKE
ET AL.
Kerguelen Equatorial Shatsky
Plateaue
Pacificf
Riseg
738 (119) 1221 (199) 1209 (198)
1.51
1.28
2.96
Suboxic
5.37
1.37
1.22
Oxic
2.49
3.22
2.52
Suboxic
5.33
2.05
1.17
Oxic
3.33
4.33
1.46
Suboxic
5.69
6.15
2.04
Suboxic
2.52
NA
2.25
Oxic
12.71
6.28
4.26
Oxic
0.91
0.79
1.97
Suboxic
2.66
0.88
1.05
Suboxic
1.09
1.20
1.11
Suboxic
3.30
1.24
1.20
Suboxic
2.75
3.35
1.69
Suboxic
4.50
4.38
1.59
Suboxic
46.97
2.08
2.06
Oxic
21.24
6.16
3.06
Oxic
2.65
2.34
2.68
Suboxic
5.57
1.81
1.50
Oxic
7.28
8.41
1.30
Suboxic
5.52
2.22
1.39
Oxic
19.50
15.24
46.91
More suboxic
6.83
7.97
1.76
Suboxic
17.17
1.70
2.11
Oxic
19.06
8.02
3.26
Oxic
C 2014. American Geophysical Union. All Rights Reserved.
V
1046
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
the oceans and atmosphere. However, at most of the other measured PETM sites, calcium carbonate
returned to pre-CIE concentrations (or overshoots them) in the recovery as the source of light carbon ceases
and the CCD gets deeper again [Kelly et al., 2005; Zachos et al., 2005; Zeebe and Zachos, 2007]. Site 401
exhibited the initial decrease in calcium carbonate at the onset of the CIE, but then calcium carbonate
remained low for the remainder of the event and in the recovery (between 30 and 50 wt % calcium carbonate). One possibility could be that this site was situated at a water depth near the lysocline during and after
the PETM, such that dissolution began to occur, but was not sufficient to form a clay layer.
Demerara Rise, ODP Site 1258, exhibits stable carbon isotopes of bulk carbonate that have more negative
values than any of the PETM records in our study (Figures 7 and supporting information Figure S1). We suspect that these negative values in the clay layer are a result of early diagenesis under sulfate reduction and
do not reflect the primary signal. Sulfate reduction occurs at this site today [Erbacher et al., 2004] and given
the location and sedimentation rates, might have altered the primary signal. Nunes and Norris [2006] also
suggested diagenetic overprinting of the d13C values of benthic foraminifera from across the PETM at
Demerara Rise because of the unusually low values and observations of calcite overgrowths on benthic foraminifer tests.
The geographical proximity of Demerara Rise to ODP Sites 999 and 1001 in the Caribbean, which also
exhibit thick clay layers and very low bulk carbonate d13C values at the PEB [Bralower, 1997], indicates that
this diagenetic feature could be related to a regional influence. The paleodepth of the two Caribbean sites,
€hl and Abrams, 2000], are similar to that of Site 1258 during the PETM, which means they
2000–2500 m [Ro
may have been influenced by the same water mass(es). With intensification of the O2 minimum zone at
intermediate depths in the region, all three sites would have experienced enhanced sulfate reduction during the PETM, possibly to the point of driving carbonate authigenesis. Given the lack of biogenic carbonate
preservation, this authigenic carbonate dominates the carbon isotope signature of the bulk carbonate, leading to the anomalously low values. Our EF data suggest that Site 1258 was close to suboxic throughout the
late Paleocene-early Eocene. Mn EF before and after reductive cleaning displays no change in the pre-CIE
and recovery intervals, indicating the presence of authigenic MnCO3 formation (Figure 4). Mn EF values are
at crustal averages during the core-CIE. U EF are slightly enriched (3) during the sampled interval, also suggesting suboxic conditions throughout the PETM (Figure 5). Thus, just a slight drop in dissolved O2 levels
would have been sufficient to result in sulfate reduction.
ODP Site 690 located on Maud Rise in the Southern Ocean represents our southernmost end-member of
the South Atlantic. Both Mn EF and U EF exhibit large broad peaks starting at 90 kyr after the onset of the
PEB and lasting 50 kyr (Figures 4 and 5). The U EF peak in the range of 65–90 is the largest U EF measured
in all our global PETM Sites. Combined with Mn EF values in the range 20–35 that show no change after
reductive cleaning, these indicate highly suboxic conditions during the recovery phase at this site. Benthic
foraminifera assemblage studies at 690 also indicate low oxygen conditions [Thomas, 2007].
Large Mn and U peaks in the recovery are coincident with a large shift in sedimentation rates at this site,
identified through the accumulation of extraterrestrial 3He [Farley and Eltgroth, 2003] (Figure 8e). Sedimentation rates are 2 cm/kyr pre-CIE, within the core of the CIE they are 1.5 cm/kyr, and double in the recov€hl et al., 2007]. The increase in sedimentation rate for the recovery is considered part of
ery to 3 cm/kyr [Ro
a negative feedback mechanism, brought on by enhanced continental weathering and runoff in response
to increased warming during the CIE [Dickens et al., 1997; Kelly et al., 2010], (Figures 8a–8d).
Our paleoredox data for Site 690 have important implications for hypotheses of local marine productivity
changes during the PETM. Specifically, an increase in mass accumulation rates (MAR) of biogenic barite at
this site (Figure 8f) has been proposed as evidence of an increase in local productivity and the rate of the
biological pump, thus acting as a strong negative feedback for the drawdown of atmospheric CO2 [Bains
et al., 2000]. Increases in Sr/Ca in coccoliths also indicate a rise in growth rates and productivity [Stoll and
Bains, 2003; Stoll et al., 2007]. However, it has been suggested that the proposed peak in barite accumulation is not productivity related but instead caused by the thermal dissociation of gas hydrates. Dissociation
of gas hydrates would create sulfate-reducing conditions prompting dissociation of barite crystals and reprecipitation of authigenic barite higher in the sedimentary column [Dickens et al., 2003]. Other data such as
a shift from nannoplankton assemblages indicating productive surface waters to more oligotrophic conditions are also in conflict with the notion of enhanced productivity during the CIE [Bralower, 2002; Kelly,
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1047
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
Figure 8. Geochemical data from ODP Site 690 plotted versus relative time before and after the PEB (Paleocene-Eocene Boundary) from
R€
ohl et al. [2007]. (a) Carbon isotope record from Bains et al. [1999], vPDB, Vienna PeeDee Belemnite. (b) Bulk sediment weight % carbonate
content from Farley and Eltgroth [2003] (smoothed line); Kelly et al. [2005] (open triangles), and this study (black X). (c) Weight percent
coarse fraction showing decrease of foraminiferal shells during recovery interval from Kelly et al. [2005]. (d) Changes in planktic foraminiferal shell fragmentation from Kelly et al. [2005]. (e) 3HeET displays an increase in flux pre and core-CIE and a decrease in the recovery interval, from Farley and Eltgroth [2003]. (f) biogenic-barium mass accumulation rate from Bains et al. [2000]. (g) Uranium enrichment factors (U
EF) calculated from measured U concentrations and the crustal molar U/Ti of 0.061 from Rudnick and Gao [2003]. Dashed vertical line is
EF 5 1 (no enrichment or depletion relative to presumed crustal source). (h) Manganese enrichment factors (Mn EF) calculated from measured Mn concentrations and the crustal molar Mn/Ti of 0.156 from Rudnick and Gao [2003]. Closed symbols indicate samples measured
before reductive cleaning. Open symbols indicate samples measured after reductive cleaning. Dashed vertical line is EF 5 1 (no enrichment or depletion relative to presumed crustal source). Gray shaded region indicates duration of core-carbon isotope excursion (CIE) 80
kyr.
2002]. Finally, the peaks in biogenic barite and Sr/Ca occur during the core of the CIE and not in the recovery (Figure 8f) [Torfstein et al., 2010].
Our trace metal EFs indicate bottom water at Site 690 during the core of the CIE and recovery was suboxic
(Figures 8g and 8h). We interpret the biogenic barite peak, which is coincident with the peak in Mn and U
EF (Figure 8f), as an artifact of the sudden change in accumulation rates and bottom water redox conditions.
This is somewhat analogous to the U EF peak documented at Walvis Ridge Site 1263 (Figure 4) [Chun et al.,
2010], and is in agreement with modeling profiles of interstitial sulfate concentrations and barite fronts,
where peaks in barite occur immediately below the interval of sulfate reduction [Dickens, 2001; Dickens
et al., 2003]. Modern interstitial sulfate concentrations at this site for the depth of the PETM are 22 mM
[Barker and Kennett, 1988] indicating sulfate reduction given modern seawater sulfate averages 28 mM.
However, if average seawater Eocene sulfate concentrations were lower (14–23 mM) [Horita et al., 2002],
then this site would not be sulfate reducing during the Eocene. Even if there is sulfate depletion in the pore
waters presently, this represents sulfate reduction some time after initial deposition of the sediments.
Finally, if paleoredox conditions were sulfate reducing, then barite would have been remobilized and possibly reprecipitated in nonsulfate-reducing depth horizons [Paytan et al., 2002].
The paleoredox features of ODP Site 738 on Kerguelen Plateau (paleodepth 1350 m [Barrera and Huber,
1991]) resemble those of Walvis Ridge, ODP Site 1263; both are at intermediate paleowater depths and suboxic throughout the PETM. However, unlike Site 1263, Site 738 exhibits a gradual increase in Mn EF (average
value 4) during the core-CIE (Figure 4). Calcium carbonate concentrations at Site 738 do not go to zero,
~a et al., 2012], similar to Site 690, indicating this site remained
with minimum values of 70 wt % [Larrasoan
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1048
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
above the CCD throughout the sampled interval. Benthic foraminifera fauna also suggest decreased oxygen
concentrations at this site [Thomas, 1998]
ODP Site 1221 located in the equatorial Pacific is one of two Pacific sites in our global study of PETM locations. Mn EF values of 185–350 during the core of the CIE are the highest measured in any of our PETM sections (Figure 4). They were most likely caused by the presence of small manganese nodules as confirmed by
shipboard measurements and visual inspection of the sediment [Lyle et al., 2002]. Biogenic barite crystals
were extracted from this section and close inspection of their morphology and sulfur isotopes do not suggest they were precipitated within the sediments during early diagenesis [Faul and Paytan, 2005]. These
indicators suggest the sediments were buried in an oxygenated environment with no postdepositional
alteration.
Our geochemical results at ODP Site 1209 on Shatsky Rise indicate oxic bottom waters throughout the
PETM and are contrary to benthic foraminiferal assemblage changes [Takeda and Kaiho, 2007]. The benthic
foraminifera oxygen indicators display a shift in the percentage of ‘‘oxic’’ benthic foraminifera during the
core-CIE interval and suggest decreased oxygen levels caused the benthic foraminifera extinction. The test
size of planktonic foraminifera also increased during the core of the CIE, a phenomenon that is attributed to
surface water warming and enhanced stratification and oligotrophic conditions [Takeda and Kaiho, 2007;
Zachos et al., 2003], a conclusion that is in agreement with independent studies of phytoplankton assemblages [Gibbs et al., 2006] and Sr/Ca measurements in individual coccoliths [Stoll et al., 2007].
While our paleoredox indicators disagree with the benthic foraminifera assemblage study at Site 1209, we
also point out that low sedimentation rates at this site (0.25 cm/kyr) and oligotrophic surface water conditions during the PETM would favor oxygenated bottom waters and surface sediments. Reduced sedimentation and decreased export of organic matter would lead to lower rates of organic matter oxidation and
higher bottom water oxygen concentrations. One alternative mechanism for the benthic extinction event at
Site 1209 is decreased food supply [Thomas, 1998, 2003], with bottom waters remaining oxygenated. The
other is a short-lived episode of suboxic conditions that was not captured by our redox indicators.
4.2. Testing the Theory of Methane Hydrate Release at the PETM
The oxygen concentration of bottom waters is sensitive to changes in methane oxidation, deep-ocean ventilation and circulation, and primary productivity. Separating the contributions of each of these factors to
the oxygen concentration of bottom waters at any given location, however, requires additional constraints
that do not yet exist for this interval. Methane hydrate dissociation and oxidation as a cause for the PETM is
testable if, for example, single specimen carbon isotopes of planktonic and benthic foraminifera from ODP
Site 690 have been used to infer a sequence of carbon release and oxidation that would be consistent with
direct emission to the atmosphere followed by gradual mixing into the deep sea [Thomas et al., 2002]. Our
study is the first to use redox-sensitive trace metals to assess bottom water oxygen concentrations across
the PETM at multiple locations in different ocean basins. With our spatial array of sites it should be possible
to distinguish local from global effects.
In theory, methane oxidation should occur first and most rapidly in the reservoir where methane is released
[Dickens, 2000]. Previous empirical and theoretical observations point toward a possible Atlantic source of
methane release during the PETM [McCarren et al., 2008; Zeebe and Zachos, 2007]. Our redox proxies indicate Atlantic intermediate water depths (Sites 1258 and 1263) experienced suboxic bottom waters during
the PETM (Figures 6 and 7 and Table 4), and were thus most proximal to the methane source. The other
intermediate water depth Atlantic locations (Sites 401 and 690) were also suboxic throughout the PETM,
but calcium carbonate at these sites indicate they remained above the CCD (Figure 4), therefore further
from the source of isotopically light carbon. Below the CCD, the deeper Atlantic Walvis Ridge Sites (Sites
1262 and 1266) remained well oxygenated both pre-CIE and in the recovery, suboxic only during the coreCIE [Chun et al., 2010]. Intermediate water depth sites (401 and 690) remain suboxic throughout the recovery, much longer than the direct influence of methane oxidation. We suggest these sites were situated in
an oxygen minimum zone throughout the duration of the PETM. Ideally, comparing multiple sites on a
depth transect (such as Walvis Ridge) would test this idea.
Moving away from the Atlantic, our intermediate depth (Site 738) in the Indian Ocean sector of the Southern Ocean was also bathed by suboxic bottom waters, which is consistent with findings based on a
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1049
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
bioturbation index of bottom water oxygen in South Pacific intermediate water sites [Nicolo et al., 2010].
Sites further north in the Pacific Ocean (Sites 1221 and 1209) experienced oxygenated bottom waters
throughout the PETM and thus were most distal from an inferred methane source. In short, these patterns
could be attributed partly to methane oxidation, though it is likely that the oxygen minimum zone
expanded as well in response to factors other than methane oxidation. Recent modeling work has shown
that with increasing atmospheric CO2 during the PETM, there was enhanced stratification and a decline in
ventilation of the Atlantic and Pacific, with the Atlantic deep waters more depleted in oxygen than Pacific
deep waters [Winguth et al., 2012].
We have used Mn and U EF paleoredox proxies to directly test the methane hydrate hypothesis during the
PETM. However, these deep-ocean redox proxies are influenced by both methane oxidation as well as
changes in ocean circulation and ventilation. There are conflicting observational evidence as to whether
there was a switch to Northern Hemisphere overturning at the start of the PETM [Nunes and Norris, 2006] or
a predominant Southern Hemisphere overturning throughout the PETM [Thomas et al., 2003]. Recent modeling studies suggest predominantly Southern Hemisphere sources with sluggish overturning associated
with deep-ocean warming [Lunt et al., 2011; Winguth et al., 2010]. We cannot form conclusions about deepocean circulation with our proxies alone.
In testing the size and pattern of carbon release during the PETM, Panchuk et al. [2008] have been able to
reproduce global patterns of calcium carbonate concentrations by reducing bioturbation in the Atlantic
during the CIE. Our interpretation of decreased bottom water oxygen concentrations in the Atlantic relative
to the Pacific would also imply a lower intensity of bioturbation in the Atlantic, in agreement with the
parameters set for their sediment model.
5. Conclusions
Acknowledgments
We thank R. Franks and the IMS for
analytical support. E. Stokes made the
carbonate measurements for Site 401
and B. Murphy assisted with stable
isotope measurements for Site 1258. B.
Young and W. Kordesch provided help
in sample preparation. C.P. thanks R.
Zeebe for pointing out the critical
need for a global redox PETM study. B.
Gill and an anonymous reviewer are
thanked for their careful comments on
the manuscript. This research used
samples and/or data provided by the
Ocean Drilling Program (ODP) and
Integrated Ocean Drilling Program
(IODP). IODP is sponsored by the U.S.
National Science Foundation (NSF) and
participating countries under
management of the Consortium for
Ocean Leadership. Funding for this
research was provided by NSF grant
EAR 0120727 to J.C.Z. and M.L.D. C.P.
was also supported by GAANN and
ARCS fellowships through UCSC Ocean
Sciences and a ‘‘Young scientists
in Focus’’ research grant at the
Goethe-University Frankfurt.
€
PALIKE
ET AL.
Using paleoredox conditions from a range of global sites, we infer the presence or absence of oxygenated
bottom waters and comment upon the possible release of methane from gas hydrates. Evidence from Mn
and U EF in marine sediments indicate that North Atlantic intermediate water depth sites were the most
suboxic during the PETM. Coupled with the pattern of CaCO3 during this interval, we suggest the source of
methane release was located in the North Atlantic. Both of our Pacific Ocean sites are oxygenated before,
during, and after the PETM. Therefore, we suggest the Pacific was not the source of methane release. Our
interpretation of the present data reveals that intermediate water depth sites in the Atlantic are the most
suboxic and likely closest to the site of methane hydrate release, while sites that are well oxygenated are
furthest from the source (Pacific locations). Furthermore, our study shows that Mn and U EF can be applied
as qualitative paleoredox proxies across transient climate states such as the PETM.
References
Bains, S., R. M. Corfield, and R. D. Norris (1999), Mechanisms of climate warming at the end of the Paleocene, Science, 285, 724–727, doi:
10.1126/science.285.5428.724.
Bains, S., R. D. Norris, R. M. Corfield, and K. L. Faul (2000), Termination of global warmth at the Palaeocene/Eocene boundary through productivity feedback, Nature, 407(6801), 171–174, doi:10.1038/35025035.
Barker, P. F., and J. P. Kennett (1988), Proceedings of the Ocean Drilling Program, Initial Rep., vol. 113, Ocean Drill. Program, College Station,
Tex., doi:10.2973/odp.proc.ir.113.1988.
Barrera, E., B. T. Huber (1991), Paleogene and early Neogene oceanography in the southern Indian Ocean: Leg 119, foraminifer stable isotope results, in Proceedings of the Ocean Driling Program, Sci. Results, edited by J. A. Barron et al., Ocean Drill. Program, College Station,
Tex., pp. 693–717.
Barron, J., B. Larsen, et al. (1989), Proc. ODP. Init. Repots, College Station, TX (Ocean Drilling Program), 119.
Boyle, E. A. (1983), Manganese carbonate overgrowths on foraminifera tests, Geochim. Cosmochim. Acta, 47(10), 1815–1819, doi:10.1016/
0016-7037(83)90029-7.
Bralower, T. J. (1997), High-resolution records of the late Paleocene thermal maximum and circum-Caribbean volcanism: Is there a causal
link?, Geology, 25, 963–966, doi:10.1130/0091-7613(1997)025<0963:HRROTL>2.3.CO;2.
Bralower, T. J. (2002), Evidence of surface water oligotrophy during the Paleocene-Eocene thermal maximum: Nannofossil assemblage
data from Ocean Drilling Program Site 690, Maud Rise, Weddell Sea, Paleoceanography, 17(2), doi:10.1029/2001PA000662.
Calvert, S. E., and T. F. Pedersen (1993), Geochemistry of recent oxic and anoxic marine sediments: Implications for the geological record,
Mar. Geol., 113(1–2), 67–88, doi:10.1016/0025-3227(93)90150-T.
Chun, C. O. J., M. L. Delaney, and J. C. Zachos (2010), Paleoredox changes across the Paleocene-Eocene thermal maximum, Walvis Ridge
(ODP Sites 1262, 1263, and 1266): Evidence from Mn and U enrichment factors, Paleoceanography, 25, PA4202, doi:10.1029/
2009PA001861.
C 2014. American Geophysical Union. All Rights Reserved.
V
1050
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
Colosimo, A. B., T. J. Bralower, and J. C. Zachos (2006), Evidence for lysocline shoaling at the Paleocene/Eocene Thermal Maximum on Shatsky Rise, northwest Pacific, in Proc. ODP, Sci. Results, 198 [Online], edited by Bralower, T. J., Premoli Silva, I., and Malone, M. J. Available
from World Wide Web: <http://www-odp.tamu.edu/publications/198_SR/112/112.htm>.
DeConto, R. M., S. Galeotti, M. Pagani, D. Tracy, K. Schaefer, T. Zhang, D. Pollard, and D. J. Beerling (2012), Past extreme warming events
linked to massive carbon release from thawing permafrost, Nature, 484(7392), 87–91, doi:10.1038/nature10929.
Dickens, G. R. (2000), Methane oxidation during the late Palaeocene thermal maximum, Bull. Soc. Geol. Fr., 171(1), 37–49.
Dickens, G. R. (2001), Sulfate profiles and barium fronts in sediment on the Blake Ridge: Present and past methane fluxes through a large
gas hydrate reservoir, Geochim. Cosmochim. Acta, 65(4), 529–543, doi:10.1016/S0016-7037(00)00556-1.
Dickens, G. R. (2003), Rethinking the global carbon cycle with a large, dynamic and microbially mediated gas hydrate capacitor, Earth
Planet. Sci. Lett., 213(3–4), 169–183, doi:10.1016/S0012-821X(03)00325-X.
Dickens, G. R. (2011), Down the Rabbit Hole: Toward appropriate discussion of methane release from gas hydrate systems during the
Paleocene-Eocene thermal maximum and other past hyperthermal events, Clim. Past, 7(3), 831–846, doi:10.5194/cp-7-831-2011.
Dickens, G. R., J. R. O’Neil, D. K. Rea, and R. M. Owen (1995), Dissociation of oceanic methane hydrate as a cause of the carbon isotope
excursion at the end of the Paleocene, Paleoceanography, 10, 965–971, doi:10.1029/95PA02087.
Dickens, G. R., M. M. Castillo, and J. C. G. Walker (1997), A blast of gas in the latest Paleocene: Simulating first-order effects of massive dissociation of oceanic methane hydrate, Geology, 25(3), 259–262, doi:10.1130/0091-7613.
Dickens, G. R., T. Fewless, E. Thomas, and T. J. Bralower (2003), Excess barite accumulation during the Paleocene-Eocene thermal maximum:
Massive input of dissolved barium from seafloor gas hydrate reservoirs, Spec. Pap. Geol. Soc. Am., 369, 11–23, doi:10.1130/0-8137-23698.11.
Dickson, A. J., A. S. Cohen, and A. L. Coe (2012), Seawater oxygenation during the Paleocene-Eocene thermal maximum, Geology, 40(7),
639–642, doi:10.1130/g32977.1.
Ellis, R., J. Pine, and J. M. Gieskes (1979), Interstitial Water Studies, Leg 48, Initial Reports of the Deep Sea Drilling Project, 48, 297–303, doi:
10.2973/dsdp.proc.48.110.1979.
Erbacher, J., D. C. Mosher, and M. J. Malone (2004), Proceedings of the Ocean Driling Program, Initial Rep., vol. 207, Ocean Drill. Program, College Station, Tex., doi:10.2973/odp.proc.ir.207.2004.
Farley, K. A., and S. F. Eltgroth (2003), An alternative age model for the Paleocene-Eocene thermal maximum using extraterrestrial 3He,
Earth Planet. Sci. Lett., 208(3–4), 135–148, doi:10.1016/S0012-821X(03)00017-7.
Faul, K. L., and A. Paytan (2005), Phosphorus and barite concentrations and geochemistry in Site 1221 Paleocene/Eocene boundary sediments, in Proceedings of the Ocean Drilling Program, Sci. Results, edited by P. A. Wilson, M. Lyle, and J. V. Firth, pp. 1–23, Ocean Drill. Program, College Station, Tex.
Froelich, P. N., G. P. Klinkhammer, M. L. Bender, N. A. Luedtke, G. R. Heath, D. Cullen, P. Dauphin, D. Hammond, B. Hartman, and V. Maynard
(1979), Early oxidation of organic matter in pelagic sediments of the eastern equatorial Atlantic: Suboxic diagenesis, Geochim. Cosmochim. Acta, 43(7), 1075–1090, doi:10.1016/0016-7037(79)90095-4.
Gibbs, S. J., T. J. Bralower, P. R. Bown, J. C. Zachos, and L. M. Bybell (2006), Shelf and open-ocean calcareous phytoplankton assemblages
across the Paleocene-Eocene thermal maximum: Implications for global productivity gradients, Geology, 34(4), 233–236, doi:10.1130/
G22381.1.
Gingele, F. X., and S. Kasten (1994), Solid-phase manganese in Southeast Atlantic sediments: Implications for the paleoenvironment, Mar.
Geol., 121(3–4), 317–332, doi:10.1016/0025-3227(94)90037-X.
Higgins, J. A., and D. P. Schrag (2006), Beyond methane: Towards a theory for the Paleocene-Eocene thermal maximum, Earth Planet. Sci.
Lett., 245(3–4), 523–537, doi:10.1016/j.epsl.2006.03.009.
Horita, J., H. Zimmermann, and H. D. Holland (2002), Chemical evolution of seawater during the Phanerozoic: Implications from the record
of marine evaporites, Geochim. Cosmochim. Acta, 66(21), 3733–3756, doi:10.1016/S0016-7037(01)00884-5.
Kaiho, K., K. Takeda, M. R. Petrizzo, and J. C. Zachos (2006), Anomalous shifts in tropical Pacific planktonic and benthic foraminiferal test
size during the Paleocene-Eocene thermal maximum, Palaeogeography, Palaeoclimatology, Palaeoecology, 237(2–4), 456–464, doi:
10.1016/j.palaeo.2005.12.017.
Kelly, D. C. (2002), Response of Antarctic (ODP Site 690) planktonic foraminifera to the Paleocene-Eocene thermal maximum: Faunal evidence for ocean/climate change, Palaeoceanography, 17 (4), 1071, doi:10.1029/2002PA000761.
Kelly, D. C., J. C. Zachos, T. J. Bralower, and S. A. Schellenberg (2005), Enhanced terrestrial weathering/runoff and surface ocean carbonate
production during the recovery stages of the Paleocene-Eocene thermal maximum, Paleoceanography, 20, PA4023, doi:10.1029/
2005PA001163.
Kelly, D. C., T. M. J. Nielsen, H. K. McCarren, J. C. Zachos, and U. R€
ohl (2010), Spatiotemporal patterns of carbonate sedimentation in the
South Atlantic: Implications for carbon cycling during the Paleocene-Eocene thermal maximum, Palaeogeogr. Palaeoclimatol. Palaeoecol., 293(1–2), 30–40, doi:10.1016/j.palaeo.2010.04.027.
Kennett, J. P., and L. D. Stott (1991), Abrupt deep-sea warming, palaeoceanographic changes and benthic extinctions at the end of the
Palaeocene, Nature, 353, 225–229, doi:10.1038/353225a0.
Kurtz, A. C., L. R. Kump, M. A. Arthur, J. C. Zachos, and A. Paytan (2003), Early Cenozoic decoupling of the global carbon and sulfur cycles,
Paleoceanography, 18(4), 1090, doi:10.1029/2003PA000908.
Larrasoa~
na, J. C., A. P. Roberts, L. Chang, S. A. Schellenberg, J. D. Fitz Gerald, R. D. Norris, and J. C. Zachos (2012), Magnetotactic bacterial
response to Antarctic dust supply during the Palaeocene-Eocene thermal maximum, Earth Planet. Sci. Lett., 333–334, 122–133, doi:
10.1016/j.epsl.2012.04.003.
Lourens, L. J., A. Sluijs, D. Kroon, J. C. Zachos, E. Thomas, U. Rohl, J. Bowles, and I. Raffi (2005), Astronomical pacing of late Palaeocene to
early Eocene global warming events, Nature, 435(7045), 1083–1087, doi:10.1038/nature03814.
Lunt, D. J., P. J. Valdes, T. D. Jones, A. Ridgwell, A. M. Haywood, D. N. Schmidt, R. Marsh, and M. Maslin (2010), CO2-driven ocean circulation
changes as an amplifier of Paleocene-Eocene thermal maximum hydrate destabilization, Geology, 38(10), 875–878, doi:10.1130/
G31184.1.
Lunt, D. J., A. Ridgwell, A. Sluijs, J. Zachos, S. Hunter, and A. Haywood (2011), A model for orbital pacing of methane hydrate destabilization
during the Palaeogene, Nat. Geosci., 4(11), 775–778, doi:10.1038/ngeo1266. [Available at http://www.nature.com/ngeo/journal/v4/n11/
abs/ngeo1266.html.].
Lyle, M., et al. (2002), Proceedings of the Ocean Drilling Program, Initial Rep., vol. 199, Ocean Drill. Program, College Station, Tex., doi:
10.2973/odp.proc.ir.199.2002.
Mangini, A., M. Jung, and S. Laukenmann (2001), What do we learn from peaks of uranium and of manganese in deep sea sediments?,
Mar. Geol., 177(1–2), 63–78, doi:10.1016/S0025-3227(01)00124-4.
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1051
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
McCarren, H., E. Thomas, T. Hasegawa, U. R€
ohl, and J. C. Zachos (2008), Depth dependency of the Paleocene-Eocene carbon isotope excursion: Paired benthic and terrestrial biomarker records (Ocean Drilling Program Leg 208, Walvis Ridge), Geochem. Geophys. Geosyst., 9,
Q10008, doi:10.1029/2008GC002116.
McDuff, R. E. (1981), Major cation gradients in DSDP interstitial waters: The role of diffusive exchange between seawater and upper oceanic
crust, Geochim. Cosmochim. Acta, 45(10), 1705–1713, doi:10.1016/0016-7037(81)90005-3.
Montadert, L., and D. G. Roberts (1979), Site 401. Init. Rept. DSDP, 48, 73–123.
Morford, J. L., and S. Emerson (1999), The geochemistry of redox sensitive trace metals in sediments, Geochim. Cosmochim. Acta, 63(11–12),
1735–1750, doi:10.1016/S0016-7037(99)00126-X.
Murphy, B. M., M. Lyle, and A. Olivarez Lyle (2006), Biogenic Burial Across the Paleocene/Eocene Boundary: Ocean Drilling Program Leg 199
Site 1221, 1–12 pp., Ocean Drilling Program, College Station, TX.
Nicolo, M. J., G. R. Dickens, and C. J. Hollis (2010), South Pacific intermediate water oxygen depletion at the onset of the Paleocene-Eocene
thermal maximum as depicted in New Zealand margin sections, Paleoceanography, 25, PA4210, doi:10.1029/2009PA001904.
Nunes, F., and R. D. Norris (2006), Abrupt reversal in ocean overturning during the Palaeocene/Eocene warm period, Nature, 439(7072), 60–
63, doi:10.1038/nature04386.
Panchuk, K., A. Ridgwell, and L. R. Kump (2008), Sedimentary response to Paleocene-Eocene thermal maximum carbon release: A modeldata comparison, Geology, 36(4), 315–318, doi:10.1130/G24474A.1.
Pancost, R. D., D. S. Steart, L. Handley, M. E. Collinson, J. J. Hooker, A. C. Scott, N. V. Grassineau, and I. J. Glasspool (2007), Increased terrestrial
methane cycling at the Palaeocene-Eocene thermal maximum, Nature, 449(7160), 332–335, doi:10.1038/nature06012.
Pardo, A., G. Keller, E. Molina, and J. Canudo (1997), Planktic foraminiferal turnover across the Paleocene-Eocene transition at DSDP Site
401, Bay of Biscay, North Atlantic, Mar. Micropaleontol., 29(2), 129–158, doi:10.1016/S0377-8398(96)00035-7.
Paytan, A., S. Mearon, K. Cobb, and M. Kastner (2002), Origin of marine barite deposits: Sr and S isotope characterization, Geology, 30(8),
747–750, doi:10.1130/0091-7613(2002)030<0747:OOMBDS>2.0.CO;2.
Paytan, A., K. Averyt, K. Faul, E. Gray, and E. Thomas (2007), Barite accumulation, ocean productivity, and Sr/Ba in barite across the Paleocene-Eocene Thermal Maximum, Geology, 35(12), 1139–1142, doi: 10.1130/G24162A.1.
Pedersen, T. F., and N. B. Price (1982), The geochemistry of manganese carbonate in Panama Basin sediments, Geochim. Cosmochim. Acta,
46(1), 59–68, doi:10.1016/0016-7037(82)90290-3.
R€
ohl, U., and L. J. Abrams (2000), High-resolution, downhole, and nondestructive core measurements from Sites 999 and 1001 in the Caribbean Sea: Application to the Late Paleocene thermal maximum, in Proceedings of the Ocean Drilling Program, Sci. Results, edited by R. M.
Leckie et al., pp. 191–203, Ocean Drill. Program, College Station, Tex.
R€
ohl, U., T. J. Bralower, R. D. Norris, and G. Wefer (2000), New chronology for the late Paleocene thermal maximum and its environmental
implications, Geology, 28(10), 927–930, doi:10.1130/0091-7613(2000)28<927:NCFTLP>2.0.CO;2.
R€
ohl, U., T. Westerhold, T. J. Bralower, and J. C. Zachos (2007), On the duration of the Paleocene-Eocene thermal maximum (PETM), Geochem. Geophys. Geosyst., 8, Q12002, doi:10.1029/2007GC001784.
Rudnick, R. L., and S. Gao (2003), Composition of the continental crust, in Treatise on Geochemistry, edited by H. D. Holland and K. K. Turekian, pp. 1–64, Pergamon, Oxford, U. K.
Sexton, P. F., P. A. Wilson, and R. D. Norris (2006), Testing the Cenozoic multi-site composite d18O and d13C curves: new Eocene monospecific records from a single locality, Demerara Rise (ODP Leg 207), Paleoceanography, 21, PA2019, doi: 10.1029/2005PA001253.
Sluijs, A., et al. (2006), Subtropical Arctic Ocean temperatures during the Palaeocene/Eocene thermal maximum, Nature, 441(7093), 610–
613, doi:10.1038/nature04668.
Stap, L., A. Sluijs, E. Thomas, and L. Lourens (2009), Patterns and magnitude of deep sea carbonate dissolution during Eocene thermal maximum 2 and H2, Walvis Ridge, southeastern Atlantic Ocean, Paleoceanography, 24, PA1211, doi:10.1029/2008PA001655.
Stoll, H. M., and S. Bains (2003), Coccolith Sr/Ca records of productivity during the Paleocene-Eocene thermal maximum from the Weddell
Sea, Paleoceanography, 18(2), 1049, doi:10.1029/2002PA000875.
Stoll, H. M., N. Shimizu, D. Archer, and P. Ziveri (2007), Coccolithophore productivity response to greenhouse event of the PaleoceneEocene thermal maximum, Earth Planet. Sci. Lett., 258(1–2), 192–206, doi:10.1016/j.epsl.2007.03.037.
Svensen, H., S. Planke, A. Malthe-Sorenssen, B. Jamtveit, R. Myklebust, T. Rasmussen Eidem, and S. S. Rey (2004), Release of methane from a
volcanic basin as a mechanism for initial Eocene global warming, Nature, 429(6991), 542–545, doi:10.1038/nature02566.
Takeda, K., and K. Kaiho (2007), Faunal turnovers in central Pacific benthic foraminifera during the Paleocene-Eocene thermal maximum,
Palaeogeogr. Palaeoclimatol. Palaeoecol., 251(2), 175–197, doi:10.1016/j.palaeo.2007.02.026.
Thomas, D. J., J. C. Zachos, T. J. Bralower, E. Thomas, and S. Bohaty (2002), Warming the fuel for the fire: Evidence for the thermal dissociation of methane hydrate during the Paleocene-Eocene thermal maximum, Geology, 30(12), 1067–1070, doi:10.1130/0091-7613(2002).
Thomas, D. J., T. J. Bralower, and C. E. Jones (2003), Neodymium isotopic reconstruction of late Paleocene-early Eocene thermohaline circulation, Earth Planet. Sci. Lett., 209(3–4), 309–322, doi:10.1016/S0012-821X(03)00096-7.
Thomas, E. (Ed.) (1998), Biogeography of the Late Paleocene Benthic Foraminiferal Extinction, pp. 214–243, Columbia Univ. Press, New York.
Thomas, E. (2003), Extinction and food at the seafloor: A high-resolution benthic foraminiferal record across the initial Eocene thermal maximum, Southern Ocean site 690, Spec. Pap. Geol. Soc. Am., 369, 319–332, doi:10.1130/0-8137-2369-8.319.
Thomas, E. (2007), Cenozoic mass extinctions in the deep sea: What perturbs the largest habitat on Earth?, Spec. Pap. Geol. Soc. Am., 424, 1–
23, doi:10.1130/2007.2424(01).
Torfstein, A., G. Winckler, and A. Tripati (2010), Productivity feedback did not terminate the Paleocene-Eocene Thermal Maximum (PETM),
Clim. Past, 6(2), 265–272, doi:10.5194/cp-6-265-2010.
Tribovillard, N., T. J. Algeo, T. Lyons, and A. Riboulleau (2006), Trace metals as paleoredox and paleoproductivity proxies: An update, Chem.
Geol., 232(1–2), 12–32, doi:10.1016/j.chemgeo.2006.02.012.
Westerhold, T., U. R€
ohl, J. Laskar, I. Raffi, J. Bowles, L. J. Lourens, and J. C. Zachos (2007), On the duration of magnetochrons C24r and C25n
and the timing of early Eocene global warming events: Implications from the Ocean Drilling Program Leg 208 Walvis Ridge depth transect, Paleoceanography, 22, PA2201, doi:10.1029/2006pa001322.
Winguth, A., C. Shellito, C. Shields, and C. Winguth (2010), Climate response at the Paleocene-Eocene thermal maximum to greenhouse
gas forcing—A model study with CCSM3, J. Clim., 23(10), 2562–2584, doi:10.1175/2009JCLI3113.1.
Winguth, A. M. E., E. Thomas, and C. Winguth (2012), Global decline in ocean ventilation, oxygenation, and productivity during the
Paleocene-Eocene thermal maximum: Implications for the benthic extinction, Geology, 40(3), 263–266, doi:10.1130/g32529.1.
Yamaguchi, T., and R. D. Norris (2012), Deep-sea ostracode turnovers through the Paleocene-Eocene thermal maximum in DSDP Site 401,
Bay of Biscay, North Atlantic, Mar. Micropaleontol., 86–87, 32–44, doi:10.1016/j.marmicro.2012.02.003.
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1052
Geochemistry, Geophysics, Geosystems
10.1002/2013GC005074
Zachos, J. C., M. W. Wara, S. Bohaty, M. L. Delaney, M. R. Petrizzo, A. Brill, T. J. Bralower, and I. Premoli-Silva (2003), A transient rise in tropical
sea surface temperature during the Paleocene-Eocene thermal maximum, Science, 302(5650), 1551–1554, doi:10.1126/science.1090110.
Zachos, J. C., et al. (2004), Proceedings of the Ocean Drilling Program, Initial Rep., vol. 208, pp. 1–112, Ocean Drill. Program, College Station,
Tex.
Zachos, J. C., et al. (2005), Rapid acidification of the ocean during the Paleocene-Eocene thermal maximum, Science, 308(5728), 1611–1615,
doi:10.1126/science.1109004.
Zachos, J. C., H. McCarren, B. Murphy, U. R€
ohl, and T. Westerhold (2010), Tempo and scale of late Paleocene and early Eocene carbon isotope cycles: Implications for the origin of hyperthermals, Earth Planet. Sci. Lett., 299(1–2), 242–249, doi:10.1016/j.epsl.2010.09.004.
Zeebe, R. E., and J. C. Zachos (2007), Reversed deep-sea carbonate ion basin gradient during Paleocene-Eocene thermal maximum, Paleoceanography, 22, PA3201, doi:10.1029/2006PA001395.
€
PALIKE
ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
1053