The Pennsylvania State University
The Graduate School
College of Earth and Mineral Sciences
WEATHERING RESPONSES TO RAPID CLIMATE CHANGE:
ANALYSIS OF A HIGH RESOLUTION OSMIUM ISOTOPE RECORD AT
THE PALEOCENE-EOCENE THERMAL MAXIMUM
A Thesis in
Geosciences
by
Rebecca A. Wieczorek
© 2012 Rebecca A. Wieczorek
Submitted in Partial Fulfillment
of the Requirements
for the Degree of
Master of Science
May 2012
The thesis of Rebecca A. Wieczorek was reviewed and approved* by the following:
Matthew S. Fantle
Assistant Professor of Geosciences
Thesis Co-Adviser
Lee R. Kump
Professor of Geosciences
Thesis Co-Advisor
Michael A. Arthur
Professor of Geosciences
Christopher J. Marone
Professor of Geosciences
Program Chair of the Department of Geosciences
*Signatures are on file in the Graduate School.
ii
ABSTRACT
The Paleocene-Eocene Thermal Maximum (PETM, ~55.8 million years ago; Ma) was an abrupt
~170 thousand year (ky) global warming event in which a large amount of carbon was added to
the atmosphere-ocean system over a few thousand years. According to our current
understanding, such perturbations of the carbon cycle are regulated over geologic timescales by
weathering processes, as atmospheric carbon is removed through the process of silicate
weathering. Evidence of past weathering (“paleoweathering”) is determined through the use of
isotope proxies; the marine osmium isotope record is one proxy commonly utilized for its ability
to record geologically rapid events due to osmium’s short (10 ky) residence time in the ocean.
Here I present an osmium isotope record (187Os/188Os) from Spitsbergen, Svalbard (core BH
9/05) spanning the PETM interval. The core is a marine grey-shale sequence with the PETMdefining carbon isotope excursion spanning ~50 m, making it ideal for high-resolution analysis.
For this study, forty-four bulk rock samples were digested using Carius tube procedures and/or
NiS fire assay bulk fusions and their rhenium (Re) and Os concentrations and Os isotopic
compositions measured using inductively-coupled plasma mass spectrometry (ICP-MS).
Bulk rock Os concentrations range from 111 to 207 pg/g while Re concentrations range from 0.7
to 13.1 ppb. Both Re and Os concentrations increase over the PETM interval, coincident with an
interval of laminated shales and elevated pyrite content, suggesting regional euxinic conditions.
187
Os/188Os ratios range from 0.3557 to 0.8238, with a definitive maximum ~6 ky after the onset
of the carbon isotope excursion (CIE), concurrent with a large increase in the proportion of
kaolinite. These increases in 187Os/188Os ratios and kaolinite content are interpreted as an
increase in continental weathering through increased riverine Os input rather than lithologic
source change. Modeling also supports this conclusion, though influence from a source change
cannot be absolutely ruled out.
Bulk elemental oxide concentrations were also measured in shales spanning the PETM in order
to provide additional constraints on both timing of events and potential weathering sources.
Bulk oxides are used to calculate the chemical index of alteration (CIA), a common weathering
intensity index. CIA measurements indicate source material deposited during the onset of the
PETM was intensely weathered. Additionally, this independent CIA record is nearly identical to
the 187Os/188Os record, with coincident maxima and minima, supporting the hypothesis that
variations in both proxies are linked to either a change in source or variation in the local
environment.
Though BH 9/05 reflects the shape of contemporaneous Os isotope records, the measured
187
Os/188Os ratios in BH 9/05 are nearly double those of open ocean seawater (~0.3-0.45) before
and after the PETM. This leads to the inference that BH 9/05 represents a restricted environment
rather than the open ocean, most likely a basin environment with limited ocean contact and
periodic episodes of anoxia. Modeling yields a global river increase of 40-50% while smaller
rivers may have tripled their Os output over this interval. Overall, though the restricted basin
setting of BH 9/05 may limit conclusions about global weathering patterns at the PETM, it
simultaneously provides an important look at the paleo-Arctic response to rapid warming and
may give further insight into climate change in the polar regions.
iii
TABLE OF CONTENTS
LIST OF FIGURES…………………………………………………………………………… vii
LIST OF TABLES……………………………………………………………………………. xi
ACKNOWLEDGEMENTS………………………………………………………………….. xii
Chapter 1 Introduction…………………………………………………………………………1
Chapter 2 Background……………………………………………………………………….... 3
2.1 Climate……………………………………………………………………………. 3
2.1.1 Early Cenozoic Climate………………………………………………… 3
2.1.2 The Paleocene-Eocene Thermal Maximum (PETM)…………...…….... 5
2.2 Weathering………………………………………………………………………... 6
2.3 Osmium Isotopes and the Os Cycle………………………………………………. 7
Chapter 3 Regional Setting…………………………………………………………………….11
3.1 Tectonic and Lithological Setting………………………………………………… 11
3.1.1 Tectonic History of the Central Basin………………………………...… 12
3.1.2 Lithology of the Van Mijenfjorden Group…………………….......……. 14
3.2 Core Description: Lithology……………………………………………………… 15
Chapter 4 Analytical methods……………………………………………………………….... 18
4.1 Osmium and Rhenium Geochemistry…………………………………………….. 18
4.1.1 Osmium and Rhenium Analysis………………………………………… 18
4.1.2 Preparation: Bulk Digestion and Carius Tube Comparison……………. 18
4.1.3 NiS Fire Assay Bulk Digestions………………………………………… 20
4.1.4 Bulk Rhenium Acid Digestions…………………………………………. 21
4.1.5 Carius Tube Digestion…………………………………………………... 22
4.1.6 Rhenium Anion Exchange Column Chemistry…………………………. 23
4.1.7 Mass Spectrometers……………………………………………………... 24
4.1.8 Blank Corrections……………………………………………………….. 25
4.2 Bulk Element Analysis……………………………………………………………. 25
4.3 Numerical Model Description…………………………………………………….. 26
Chapter 5 Results………………………………………………………………………………29
5.1 Osmium and Rhenium Concentrations and 187Os/188Os Measurements………….. 29
5.2 Bulk Oxide Analysis……………………………………………………………… 33
iv
Chapter 6 Discussion………………………………………………………………………….. 37
6.1 BH 9/05 as a Basin Signal………………………………………………………… 37
6.2 Evidence for Enhanced Chemical Weathering over the PETM…………………... 41
6.3 The 187Os/188Os and CIA Maxima: Weathering vs. Lithologic Change…………..43
6.4 The Potential Influence of Diagenesis……………………………………………..49
6.5 The 187Os/188Os Minimum: Evidence of Seawater Intrusion…………………….. 52
6.5 Proposed Change in Depositional Environment over the PETM…………………. 58
6.5.1 Effects of Terrestrial Runoff and Rising Sea Level…………………….. 58
6.5.2 Evidence for local hypoxia and a freshwater cap………………………. 59
6.5.3 Summary of depositional environment of BH 9/05…………………….. 64
Chapter 7 Conclusions…………………………………………………………………………65
References…………………………………………………………………………………….. 66
Appendices……………………………………………………………………………………. 73
Appendix A: Osmium and Rhenium Measurements…………………………………. 73
Appendix B: Model Setup and Initial Parameters……………………………………. 76
1. Global model (1-box)……………………………………………………... 76
1.1 Parameters and assumptions…………………………………... 76
1.2. Schematics……………………………………………………. 79
1.2.1 Fr Model Setup……………………………………… 79
1.2.2 Rr Model Setup……………………………………... 80
2. Basin-ocean combined model (2-box)…………………………………….. 81
2.1 Matching C-data from BH 9/05 and DSDP 549—Case 1
and Case 2………………………………………………………... 81
2.2 Models………………………………………………………… 83
2.3 Schematics…………………………………………………….. 84
2.3.1 Fr Model Setup……………………………………… 84
2.3.2 Rr Model Setup……………………………………... 85
2.3.3. Fin Model Setup……………………………………. 86
2.4 Combined Fin-Fr model………………………………………. 86
2.4.1 Parameters and assumptions………………………… 87
Appendix C: Model Output…………………………………………………………... 91
1. Global model………………………………………………………………. 91
1.1 Fr model………………………………………………………. 91
1.2 Rr model………………………………………………………. 92
2. Basin-ocean combined model……………………………………………... 93
2.1 Fr model………………………………………………………. 93
2.2 Rr models……………………………………………………... 95
v
2.3 Fin model……………………………………………………… 98
2.4 Fin model with river increase…………………………………. 99
vi
LIST OF FIGURES
Fig. 1: Geochemical record of the early Cenozoic showing the PETM excursion in carbon and oxygen
isotopes, modified from Zachos et al. (2001). Measurements are reported relative to the VPDB
(Vienna Pee Dee Belemnite) standard. Data are 5-pt averages of carbon and oxygen isotopes
measured in deep-sea foraminifera (sp. Cibicidoides and Nuttallides) at a variety of DSDP and
ODP sites. ............................................................................................................................................................................ 4
Fig. 2: Schematic of the basic Os and Re budgets, compiled from the literature. (Ravizza, 1993;
Levasseur et al., 1999; Peucker-Ehrenbrink and Ravizza, 2000; Chen et al., 2006; Colodner et al.,
1993; Hauri and Hart, 1997; Yamashita et al., 2007) ....................................................................................... 8
Fig. 3: Plate orientation during the opening of the Norwegian-Greenland Sea (modified from
Talwani and Eldholm (1977), showing Svalbard’s paleolocation and relative proximity to
Greenland and subsequent counterclockwise motion of Greenland relative to Svalbard. The red
and green colors are used to denote opposite sides of the spreading ridge which became the
north end of the Mid-Ocean Ridge system. ......................................................................................................... 12
Fig. 4: Geologic map of Spitsbergen, showing location of Core BH 9/05 in red. Modified from Hjelle
(1993) and Riber (2009). Core BH 9/05 is part of the Van Mijenfjorden group, a Tertiary
deposit covering the Central Basin. During the uplift of the West Spitsbergen Orogeny, the
Central Basin became a foreland-basin environment (Müller and Spielhagen, 1990). .................... 13
Fig. 5: (left) Stratigraphic column of BH 9/05, modified from Riber (2009). The most negative δ13C
values occur between 536 and 519.7 m (Cui et al, 2011), followed by a gradual recovery. An
interval of high siderite and pyrite is present prior to and during the δ13C minimum (Riber,
2009)................................................................................................................................................................................... 16
Fig. 6: a) The carbon isotope curve from BH 9/05 (Cui et al., 2011) b) Age-corrected osmium
isotope ratios (187Os/188Os) over the PETM, c) Osmium concentration over the PETM, in parts
per trillion (ppt). d) Rhenium concentration over the PETM, in parts per billion (ppb). Blue
diamonds indicate bulk measurements, while the red circles represent Carius tube
measurements. The gray line indicates the depth of the 187Os/188Os minimum while the dashed
line marks the depth of the 187Os/188Os maximum. Note that age-corrected 187Os/188Os ratios
are considerably less variable than measured 187Os/188Os ratios (Appendix A) and that initial
187Os/188Os ratios determined by different methods typically agree within <0.03 187Os/188Os
units..................................................................................................................................................................................... 31
Fig. 7: a) Comparisons of measured bulk vs. Carius tube 187Os/188Os ratios. Agreement in
187Os/188Os ratios appears to be the best overall, with a best-fit slope of 0.988. b) Comparisons
of bulk vs. Carius tube measurements for osmium concentration. Agreement in this panel is the
lowest, with a correlation value of 0.59. c) Comparisons of bulk vs. Carius tube measurements
for Re concentration. Re concentrations have a relatively good correlation of 0.8649 and an R2
value of 0.94. In each panel a 1:1 line (blue dashed line) is included for reference. Differences
in Os and Re concentrations are attributed to sample powder inhomogeneity. Note that in order
to make a direct comparison of Os values, Os measurements are not age corrected, so that
differences in Re will not affect the result. See Appendix A for table of values. ................................. 32
Fig. 8: Record of increased weathering over the PETM. a) (top): CIA measurements for BH 9/05,
including both bulk CIA and residual CIA (CIA of the remainder of the material sans kaolinite),
with the 187Os/188Os record for comparison. Dashed line = depth 533.07 m. Note that this is
where the maximum occurs in 187Os/188Os ratios and kaolinite ratios. Residual CIA values were
calculated using a mass balance equation and interpolated at depths of existing kaolinite
vii
measurements from Riber (2009). b) (bottom): kaolinite ratios in BH 9/05 for comparison,
from Riber (2009). Kaolinite ratios, the CIA, and 187Os/188Os ratios are weathering indicators.36
Fig. 9: Results from the global ocean model for changes in the river Os flux. This model was made
to see if both core records could be caused by nearly the same input fluxes. However, since the
required Os kg/yr input for BH 9/05 is more than twice that of DSDP 549, the conclusion is that
they cannot both be recording the same signal (since Os is well-mixed in the ocean (Ravizza and
Zachos, 2003), the two signals would not be influenced by different sources, the other
possibility), and so they cannot both be an ocean Os signal. ....................................................................... 39
Fig. 10: Results from the global ocean model for changes in the river 187Os/188Os ratio. Changes in
Rr causing BH 9/05 would require a global isotope ratio higher than 2. While this is not
impossible, it would require a higher crustal abundance of radiogenic material exposed at the
surface. The question arises whether such large-scale lithology exposure changes can occur
over such a short geological time. Thus the conclusion of this study is that the observed
187Os/188Os records were most likely not caused by a change in the global Rr alone. ...................... 40
Fig. 11: Results from the basin-ocean model for changes in the river Os flux. Note that in this
scenario, the Arctic river Os shows a much greater proportional increase than the global river
average, fitting with both a greater temperature change in response to stimuli (i.e. Sluijs et al.,
2006) and increased precipitation at the poles (Pagani et al., 2006), both of which would serve
to increase the total amount of Os reaching the basin from the river. .................................................... 45
Fig. 12: Results from the basin-ocean model for changes in the basin river 187Os/188Os ratio. Note
that this scenario would require a maximum basin river Rr value near 3. While this cannot be
completely ruled out, it would require changes to much more radiogenic lithology at the surface
and back within the span of the PETM, and the timing of the increase and maximum 187Os/188Os
ratio aligning with the increase and maximum kaolinite (another indicator of weathering) to be
coincidental. Therefore, while changes in Rr influencing the basin Os record cannot be ruled
out entirely, a sole Rr cause is unlikely. ............................................................................................................... 46
Fig. 13: Results from the basin-ocean model for changes in both the Arctic River and the global
average 187Os/188Os ratio. To achieve a global river average 187Os/188Os ratio near 2, the
exposed lithology would have to be much more radiogenic than today. While I cannot eliminate
changes in Rr as a possible influence on the Os record, it is questionable whether such largescale changes could have occurred over such a short geological time, and therefore unlikely that
Rr changes were responsible for the entire change in the Os record. ..................................................... 47
Fig. 14: Hypothetical scenario of the case of rhenium loss. Figure a) [left] is the actual Re record
from BH 9/05, while b) is a close-up of the circled region with hypothetical pre-loss estimated
Re values. .......................................................................................................................................................................... 50
Fig. 15: Changes in oceanic water flux to the basin. a) top: the 187Os/188Os record from BH 9/05. b)
middle: Results from the basin-ocean model for changes in the Os exchange flux with the ocean.
c) bottom: Results from the basin-ocean model for changes in the water exchange flux with the
ocean. These data were calculated based on the molar Os input flux from the ocean to the basin
(Fin) divided by the ocean concentration at time (t) and converted to kg H2O. ................................. 55
Fig. 16: Results from the basin-ocean model for the most realistic scenario in which river Os input
continued to increase prior to the PETM but was overridden by a seawater flooding event from
4 to 19.4 ky after the model leaves steady state at 20 ky. a) top panel shows the hypothesized
river input to the basin (the output from the Fr basin-ocean model) with the dashed line
representing the ‘swamping’ of the 187Os/188Os signal after a seawater excursion. b) bottom
viii
panel shows the required increase in seawater input to the basin necessary to override the river
signal during that time to achieve the recorded 187Os/188Os ratio. ........................................................... 56
Fig. 17: Hypothesized bottom-water environment over the course of the PETM. Reducing or
suboxic conditions are assumed to be correlated with increases in Re concentrations and pyrite
contents. ............................................................................................................................................................................ 59
Fig. 18: schematic of the global model setup for the Fr case (driving the observed marine
187Os/188Os ratio by changes in the riverine Os flux). ..................................................................................... 79
Fig. 19: schematic of the global model setup for the Rr case (driving the observed marine
187Os/188Os ratio by changes in the average global river 187Os/188Os ratio).......................................... 80
Fig. 20: top: Comparison of the carbon isotope records of BH 9/05 and DSDP 549 assuming Case 1
and 2 matching. Carbon isotope data from Ravizza et al. (2001) is measured from bulk
carbonate, while carbon isotope data from BH 9/05 is total organic carbon (Cui et al., 2011).
bottom: Comparison of the corresponding 187Os/188Os records of BH 9/05 and DSDP 549 for
Case 1 and 2 data matching. ...................................................................................................................................... 82
Fig. 21: Schematic of the basic Fr model (change in river Os flux) for the combined basin-ocean
system. For these models, BH 9/05 is assumed to be the Os record for the basin, while DSDP
549 is assumed to be the global ocean Os record. ........................................................................................... 84
Fig. 22: Schematic of the basic Rr model (change in local and global river 187Os/188Os ratio) for the
combined basin-ocean system. ................................................................................................................................ 85
Fig. 23: Schematic of the basic Fin model (change in the exchange flux between the basin and
ocean). Note that while in this case the basin signal is assumed to be caused by changes in the
exchange flux (Fin and Fout), the ocean signal remains a product of changes in the global river
Os flux (global Fr). ......................................................................................................................................................... 86
Fig. 24: Results from the global ocean model for changes in the river Os flux. This model was made
to see if both core records could be caused by nearly the same input fluxes. However, since the
required Os kg/yr input for BH 9/05 is more than twice that of DSDP 549, the conclusion is that
they cannot both be recording the same signal (since Os is well-mixed in the ocean (Ravizza and
Zachos, 2003), the two signals would not be influenced by different sources, the other
possibility), and so they cannot both be an ocean Os signal........................................................................ 91
Fig. 25: Results from the global ocean model for changes in the river 187Os/188Os ratio. Changes in
Rr causing BH 9/05 would require a global isotope ratio higher than 2. While this is not
impossible, it would require a higher crustal abundance of radiogenic material exposed at the
surface. The question arises whether such large-scale lithology exposure changes can occur
over such a short geological time. Thus the conclusion of this study is that the observed
187Os/188Os records were most likely not caused by a change in the global Rr alone. ...................... 92
Fig. 26: Results from the basin-ocean model for changes in the river Os flux. Note that in this
scenario, the Arctic river Os shows a much greater proportional increase than the global river
average, fitting with both a greater temperature change in response to stimuli (i.e. Sluijs et al.,
2006) and increased precipitation at the poles (Pagani et al., 2006), both of which would serve
to increase the total amount of Os reaching the basin from the river. .................................................... 94
Fig. 27: Results from the basin-ocean model for changes in the basin river 187Os/188Os ratio. Note
that this scenario would require a maximum basin river Rr value near 3. While this cannot be
completely ruled out, it would require changes to much more radiogenic lithology at the surface
and back within the span of the PETM, and the timing of the increase and maximum 187Os/188Os
ratio aligning with the increase and maximum kaolinite (another indicator of weathering) to be
ix
coincidental. Therefore, while changes in Rr influencing the basin Os record cannot be ruled
out entirely, a sole Rr cause is unlikely. ............................................................................................................... 95
Fig. 28: Results from the basin-ocean model for changes in both the Arctic River and the global
average 187Os/188Os ratio. To achieve a global river average 187Os/188Os ratio near 2, the
exposed lithology would have to be much more radiogenic than today. While I cannot eliminate
changes in Rr as a possible influence on the Os record, it is questionable whether such largescale changes could have occurred over such a short geological time, and therefore unlikely that
Rr changes were responsible for the entire change in the Os record. ..................................................... 97
Fig. 29: Results from the basin-ocean model for changes in the Os water exchange flux with the
ocean. These data were calculated based on the molar Os input flux from the ocean to the basin
(Fin) divided by the ocean concentration at time (t) and converted to kg H2O. ................................. 98
Fig. 30: Results from the basin-ocean model for the most realistic scenario in which river Os input
continued to increase prior to the PETM but was overrided by a seawater flooding event from 4
to 19.4 KY after the model leaves steady state at 20 KY. The top figure shows the hypothesized
river input to the basin (the output from the Fr basin-ocean model) with the dashed line
representing the ‘swamping’ of the 187Os/188Os signal after a seawater excursion. The middle
figure shows the required increase in water input to the basin from the ocean necessary to
override the river signal during that time to achieve the recorded 187Os/188Os ratio. .................. 100
x
LIST OF TABLES
Table 1: Major oxides reported as weight percents, from bulk oxide analysis. Average oxide
measurements for BH 9/05 appear to be closer to an average shale than to a typical gray shale,
exceptions being lower than average levels of CaO and SiO2. ..................................................... 34
Table 2: Osmium and Rhenium Analyses from BH 9/05. (*) indicates age-corrected values. 2
values for Re are < 4% of the recorded value. .............................................................................. 73
Table 3: Comparison of measured Os concentration (top left) and 187Os/188Os ratios (top right)
from Carius tubes and from bulk samples. 187Os/188Os ratios are direct measurements (pre-age
correction) so that differences in Re values have no influence on the result. Re measurements
are included separately (bottom). Note that the samples with the largest difference in 187Os/188Os
also have large differences between measured Os and Re concentrations, suggesting that sample
splits are not compositionally homogenous. ................................................................................. 75
Table 4: Initial parameters for the global model. ......................................................................... 78
Table 5: Different model simulations used in this study. ............................................................ 83
Table 6: Initial parameters for the Basin-Ocean combined model. ............................................. 89
Table 7: Summary of salinity results for the different models ................................................... 101
xi
ACKNOWLEDGMENTS
First, I would like to thank my committee (Lee Kump, Mike Arthur, and Matt Fantle) for the
invaluable aid and advice they have given during the course of the project. I would like to thank
them for always being available to talk and their willingness to help with any issue that arose. I
especially would like to thank them for their guidance, suggestions for future endeavors, and
constant encouragement over the course of the project.
I am also indebted to Greg Ravizza, Denys Vonderhaar, and Francois Paquay for their
kindness and freedom of time, information, and use of lab space and materials. I very much
appreciate all of the time they spent instructing me in laboratory techniques, making sure my stay
was enjoyable, and all of the effort and personal oversight they have put in since in making sure
my project went smoothly. Thank you so very much.
I would also like to thank the staff at both Penn State and the University of Hawaii at Manoa
for their helpfulness and aid both in helping with paperwork and in making my stay comfortable.
A big thank you also goes out to Nichole Wonderling and Henry Gong for their time,
explanations, and help with analysis.
Lastly, I would like to thank all of the wonderful supportive people in my life who have
made this thesis possible through their encouragement and love and words of wisdom, including
my parents, Robby, the wonderful group at PSCG, and my officemates.
xii
Chapter 1
Introduction
The aim of this master’s thesis is to characterize weathering responses to rapid changes in
climate by focusing on a Cenozoic hyperthermal event (the Paleocene-Eocene Thermal
Maximum, PETM). Negative isotopic excursions in carbon (-2 to -5‰) and oxygen (~ -1.5‰)
(Zachos et al., 2001) during the PETM imply a temperature increase of ~5°C and release of
petagrams (Pg) of methane or CO2 to the ocean-atmosphere system (e.g. Zachos et al, 2001;
Sluijs et al., 2006; Cui et al., 2011). The PETM is also characterized by shifts in precipitation
patterns (Pagani et al., 2006), widespread deep-sea sedimentary carbonate dissolution (Zachos et
al., 2005), sea level rise (Sluijs et al., 2006), and benthic foraminiferal extinction (Kennett and
Stott, 1991).
In order to maintain a stable climate system, processes must exist to remove excess CO2 from
the atmosphere. In the geologic carbon cycle, silicate weathering reactions are one potential
mechanism for removal (e.g., Walker et al., 1981). Accordingly, understanding the response of
silicate weathering to rapid temperature change is critical to the parameterization of climate
models and, hence, to obtaining accurate projections of future climate change. This is especially
relevant today, as such knowledge may prove essential to correctly anticipate the long-term
climatic consequences of anthropogenic CO2 emissions.
Geochemical proxy records can provide well-resolved snapshots of physical and chemical
weathering through time. Silicate weathering rates have been determined from ion
concentrations in rivers (used to measure both physical erosion (e.g. Jacobson and Blum, 2003)
and chemical weathering (e.g. Hren et al., 2007; Bluth and Kump, 1994)), laboratory
1
measurements (e.g. White et al., 1999), soil profiles (e.g. Stonestrom et al., 1998), and porewater
measurements (e.g. Murphy et al., 1998). In riverine and marine environments, studies of
strontium and osmium isotopes have been used to explore changes in silicate weathering over
time (Raymo et al., 1988; Ravizza, 1993; Peucker-Ehrenbrink and Ravizza, 1996). Marine
records of Sr and Os are of special interest because they are believed to provide a globally
integrated record of weathering-induced changes in oceanic inputs. Weathering changes across
the PETM interval are particularly important to reconstruct in order to answer key questions
about the magnitude of weathering responses to elevated levels of greenhouse gases resulting in
specific increases in temperature and precipitation. In this study I use geochemical analyses of
osmium and rhenium of bulk sediment samples, coupled with bulk elemental oxide composition,
to constrain terrestrial weathering during the PETM. This thesis presents a new high-resolution
weathering record at the PETM, interprets the record using both multiple proxies and numerical
modeling, and discusses implications for global and/or regional weathering changes during the
PETM.
2
Chapter 2
Background
2.1 Climate
2.1.1. Early Cenozoic Climate
During the Cenozoic (65.5 Ma to present), a significant long-term shift in global climate
occurred as the Earth transitioned from a warm, nearly ice free “greenhouse” world to a cold,
glacial “icehouse.” Paleogene climate was warm and wet (Tripati et al., 2001; Kalgutkar and
McIntyre, 1991), and characterized by warm sea-surface temperatures (Tripati et al., 2003) and
subtropical temperatures in the polar regions with little or no ice on land (Jahren, 2007; Sluijs et
al., 2006). The Paleogene Arctic in particular was temperate enough to support mixed coniferangiosperm forests and warm-climate taxa (i.e. crocodilians, tortoises) (Jahren, 2007; Estes and
Hutchison, 1980). This climate was likely due to a combination of factors including high
concentrations of atmospheric greenhouse gases (Paleocene estimates vary from 2 (Royer, 2006)
to 8 times modern atmospheric carbon dioxide levels (Pearson and Palmer, 2000)) and greater
volcanic emissions than the present (Zachos et al., 2008). Climate warmed throughout the late
Paleocene and early Eocene, reached a Cenozoic maximum at the Early Eocene Climatic
Optimum (52 Ma), and generally cooled through the rest of the Cenozoic (Zachos et al., 2001;
Zachos et al., 1993).
Superimposed upon the late Paleocene-early Eocene warming trend are several short
instances of rapid warming (so-called “hyperthermals”), which lasted several tens to a few
hundred thousand years each. These include the Paleocene-Eocene Thermal Maximum (PETM)
3
and the Eocene Thermal Maximum 2, among others (Sluijs et al., 2008). Because these
geologically brief warming events are primarily indicated by large excursions in the carbon
isotope record, and are often accompanied by extinctions or faunal turnover in the marine fossil
record (Sluijs et al., 2008; Zachos et al., 1993), these hyperthermals are the focus of intense
study in the geosciences.
Fig. 1: Geochemical record of the early Cenozoic showing the PETM excursion in carbon and
oxygen isotopes, modified from Zachos et al. (2001). Measurements are reported relative to the
VPDB (Vienna Pee Dee Belemnite) standard. Data are 5-pt averages of carbon and oxygen
isotopes measured in deep-sea foraminifera (sp. Cibicidoides and Nuttallides) at a variety of
DSDP and ODP sites.
4
2.1.2. The Paleocene-Eocene Thermal Maximum (PETM)
The first and largest of these early Cenozoic hyperthermals (known as the Paleocene-Eocene
Thermal Maximum, or PETM) occurred ~ 55.8 Ma (Charles et al., 2011) at the PaleoceneEocene boundary. The PETM occurred abruptly and lasted only 170 ky (Röhl et al., 2007) but is
preserved in a variety of sites worldwide covering a geographical range from Antarctica to the
Arctic (Riber, 2009; Kennett and Stott, 1991; Sluijs et al., 2008; Sluijs et al., 2006). During the
PETM, marine and terrestrial realms record large negative excursions (Fig. 1) in carbon (~3‰)
and oxygen isotopes (~1.5‰) (i.e. Kennett and Stott, 1991; Ravizza and Zachos, 2003;
McCarren et al., 2008). The abruptness and magnitude of the carbon isotope excursion (CIE),
together with large-scale marine carbonate dissolution (Zachos et al., 2005), implies a massive
release of carbon-bearing greenhouse gas to the atmosphere in a relatively short period of time
(~20 ky (Cui et al., 2011; Zachos et al., 2005)).
Other effects of the PETM included an average global temperature increase of 5°C (Thomas
et al., 2002; Zachos et al., 2001), mass extinctions of benthic foraminifera (Kennett and Stott,
1991; Thomas and Shackleton, 1996), regional ocean anoxia (Thomas and Shackleton, 1996), a
strengthening of the global hydrologic cycle and increased precipitation transport to the poles
(Pagani et al., 2006), floral changes on land (Wing et al., 2005), and eustatic sea level variations
during the event (Sluijs et al., 2008).
The source of the isotopically light carbon that triggered the PETM is debated. Possible
sources include: (1) methane from the dissolution of methane hydrates due to bottom-water
warming in the ocean (e.g. Dickens et al., 1995) or mechanical slope failure (Katz et al., 2001),
(2) contact metamorphism-released CH4 during the rifting of the North Atlantic (Svensen et al.,
2004), or (3) the release of carbon dioxide or methane during massive burning of peat due to
5
wildfires (Kurtz et al., 2003). Depending on the carbon species and their carbon isotope
composition, recent estimates place the overall amount of carbon released from 2500 (CH4) to
13000 (CO2) Pg (Cui et al., 2011).
A massive increase in atmospheric carbon should effect a change in silicate weathering
through increased temperatures and precipitation (e.g. Berner et al., 1983). During the PETM,
increased amounts of the clay mineral kaolinite (Al2Si2O5OH4) appear in many oceanic records
locations worldwide (Gibson et al., 2000; Clechenko et al., 2007). Traditionally, kaolinite is
taken as an indicator of intense chemical weathering, as it is the end product of silicate
weathering. Because kaolinite forms in warm, humid climates in the modern system, its presence
in the geologic record suggests warm, humid conditions during the PETM (i.e. Clechenko et al.,
2007). Some have questioned whether the PETM was long enough for kaolinite to have formed
via weathering on the continents, instead suggesting that kaolinite in pre-existing deposits was
simply mobilized during the PETM (Thiry, 2000).
2.2 Weathering
The climate system is controlled by a series of positive and negative feedbacks in which the
main negative feedback over geologic time scales is thought to be silicate weathering (Walker et
al., 1981). Silicate weathering can account for long-term atmospheric CO2 removal according to
the reactions of silicate minerals with atmospheric carbon dioxide and water to form calcium
ions, bicarbonate, and silica (Eq. 1), and then calcium and bicarbonate reaction to form calcium
carbonate, water, and carbon dioxide (Eq. 2) (Berner et al., 1983):
6
CaSiO3 + 2CO2 + H2O  Ca2+ + 2HCO3- + SiO2
Ca2+ + 2HCO3-  CaCO3 + H2O + CO2
(Eq. 1)
(Eq. 2)
which, when combined, gives:
CaSiO3 + CO2  CaCO3 + SiO2
(Eq. 3)
The strength of this weathering “pump” is thought to be dependent primarily on temperature and
precipitation (White et al., 1999), though the exact degree of climatic influence is uncertain
(Riebe et al., 2001; Edmond and Huh, 1997). Extent of weathering can be estimated by using the
ratio of the major oxides, known as the Chemical Index of Alteration (CIA) (Nesbitt and Young,
1982).
Isotopic proxies can also be useful in estimating paleoweathering because they provide
independent data sets for evaluating weathering. For this study, osmium was chosen as a
geochemical proxy because of its short residence time/responsiveness to short-term changes in
fluxes. Osmium has been used previously in a variety of paleoweathering studies (see Ravizza
and Zachos (2003) for a comprehensive review).
2.3 Osmium Isotopes and the Os Cycle
Osmium (Os) has seven naturally occurring nuclides (184, 186, 187, 188, 189, 190, 192), two
of which are products of radioactive decay:
187
Os (from the decay of 187Re, 1/2 = 4.35·1010 a)
and 186Os (from the decay of 190Pt, 1/2 = 6.5·1011 a). 186Os is technically radioactive (1/2 =
2.0·1015 a), decaying to 182W, but is treated as a stable nuclide over relatively short million-year
time scales (Dickin, 2005; Dąbek and Halas, 2007).
7
Fig. 2: Schematic of the basic Os and Re budgets, compiled from the literature. (Ravizza, 1993;
Levasseur et al., 1999; Peucker-Ehrenbrink and Ravizza, 2000; Chen et al., 2006; Colodner et
al., 1993; Hauri and Hart, 1997; Yamashita et al., 2007)
Osmium isotopes can be diagnostic for silicate sources of weathering, as they are sensitive to
both weathering intensity and changes in lithologic source material. Osmium in seawater is a
mixture of Os obtained from rivers (from the weathering of continental rocks), seafloor basalts,
and cosmic dust (Chen et al., 2006; Peucker-Ehrenbrink, 1996; Peucker-Ehrenbrink and Ravizza,
2000). Fig. 2 summarizes what is known about the modern Os cycle. Average continental crust
has a high (i.e., “radiogenic”) 187Os/188Os ratio of ~1.26-1.4 (Sharma et al., 1999) and a high Os
concentration (0.03-0.05 ppb) (Ravizza, 1993; McDaniel et al., 2004). The most radiogenic
lithologies are granites and black shales. Granites typically have a low Os concentration of ~10
pg/g but can have high 187Os/188Os ratios between 1 and >2.5 (Sharma et al., 1999; Chen et al.,
8
2006), due to preferential Re partitioning in the melt and subsequent decay to 187Os (Ravizza,
1993). Organic-rich shales have very high Os ratios, some of the highest reported in the
literature, with 187Os/188Os ratios as high as 14.5 and Os concentrations between 40 ppt (parts per
trillion) and 13.5 ppb (Jaffe et al., 2002; Pierson-Wickmann et al., 2001; Ravizza and Turekian,
1989). Basalt and extraterrestrial materials have unradiogenic Os isotopic compositions, and
contribute Os with low 187Os/188Os ratios when they dissolve (Ravizza et al., 2001; PeuckerEhrenbrink, 1996). Meteorites have typical 187Os/188Os ratios between 0.12 and 0.13 and Os
concentrations of approximately 500 ng/g (Peucker-Ehrenbrink and Ravizza, 2000). Inputs to the
ocean from basalt comes from two sources: high-temperature weathering at spreading ridges and
low-temperature alteration away from ridges; such input has typical 187Os/188Os of 0.12 to 0.13
and Os concentrations of 2.8 to 98 fg/g (Peucker-Ehrenbrink and Ravizza, 2000; Esser and
Turekian, 1993).
Modern seawater is assumed to be homogenous with respect to Os, with a 187Os/188Os ratio of
1.06 and an Os concentration of ~10 pgOs/kgH2O (Levasseur et al., 1998; Pegram et al., 1992;
Ravizza, 1993). While the residence time of Os is debated, most estimates place it on the order
of 104 years, longer than the oceanic mixing time, meaning that Os should be homogenous in its
concentration and isotopic composition in the ocean, which has a mixing time of 1200 years, yet
responsive to ~200 ka events such as the PETM (Ravizza and Zachos, 2003; Peucker-Ehrenbrink
et al., 1995).
Osmium in the water column is assumed to be removed with no isotopic fractionation, which
makes it ideal as a geochemical proxy. Osmium is removed from the water column during
precipitation of ferromanganese crusts (Klemm et al., 2005) or sorbed onto organic matter (OM)
(Koide et al., 1991; Pierson-Wickmann et al. 1999). Ravizza and Turekian (1992) suggested that
9
OM-associated Os is dominantly hydrogenous and thereby records a seawater Os signal. In
general, the Os found in the sediment is a mix of detrital Os from particulate continental
material, hydrogenous (seawater-derived Os), and extraterrestrial dust. It has been estimated that
70-80% of the hydrogenous osmium is derived from the continents, while the remaining 20-30%
is comprised of unradiogenic Os from hydrothermal sources and extraterrestrial dust (Sharma et
al., 1997; Levasseur et al., 1999). Extraterrestrial dust is usually ignored because it is assumed to
be only a minor part of the Os in seawater at any given time (Peucker-Ehrenbrink and Ravizza,
2000).
An understanding of rhenium (Re) is also important to the use of Os as a geochemical proxy
because rhenium concentrations are used to age-correct 187Os/188Os ratios to original in situ
values. Rhenium is conservative in the ocean and typically concentrates in deposited organic
matter (Jaffe et al., 2002; Colodner et al., 1993); Re is taken up in the sediments, at the anoxic
boundary just slightly below the depth of U but before the depth of Mo reduction (Crusius et al.
1996). A study by Yamashita et al. (2007) exploring general changes in partitioning in Os and
Re in the sediment with changing redox state found that Re was much more sensitive than Os to
redox state and would only be removed under reducing conditions, while Os could be removed
under both reducing and oxidizing conditions. However, more work needs to be done in this
area to quantify the Re and Os relationships within specific redox conditions, since changes in
partitioning affect the interpretation of bulk sediment Os and Re concentration records.
10
Chapter 3
Regional Setting
3.1 Tectonic and Lithological Setting
Svalbard is an island group (whose major islands are Spitsbergen (largest) followed by
Nordaustlandet, Edgeøya, and Barentsøya) in the Arctic Ocean, located east of Greenland and
north of Norway. Prior to the Cenozoic, Svalbard was part of a subcontinent that included
Greenland, Canada, the British Isles, Scandinavia, and the eastern United States (Appalachia).
The majority of the area’s tectonic history has been reconstructed by Harland (1969). The oldest
exposed sediments are late-Precambrian/early Paleozoic in age and display signs of compression
that occurred as a result of plate collision during the formation of the Caledonian orogeny,
accompanied by several episodes of volcanism and large-scale metamorphism during the
Devonian. One major orogenic episode occurred on Svalbard during the Silurian and early
Devonian, with a later minor episode of tectonic uplift and faulting during the Carbonaceous
(Harland, 1969). During the early part of the Cenozoic, the North Atlantic continental land mass
began to break apart with the opening of the mid-ocean ridge in the North Atlantic (Fig. 3)
(Harland, 1969).
Mid-ocean ridge volcanism during the early Cenozoic breakup of the North Atlantic region is
invoked as one possible trigger for the PETM (Jolley et al. 2002; Svensen et al., 2004). Jolley et
al. (2002) and Svensen et al. (2004) argued that North Atlantic volcanism caused the observed
carbon event either directly or indirectly, through magmatic intrusion into carbonaceous
sediments and subsequent release of methane from clathrates rather than through direct release of
carbon dioxide. Additionally, the large North Atlantic Igneous Province was emplaced during
11
the early Cenozoic (Storey et al., 2007). The PETM falls in the middle of the period of North
Atlantic igneous activity, which began around 61 MA and continued episodically until
approximately 50 MA (Storey et al., 2007). Fig. 3 shows plate motion during the opening of the
Norwegian Sea. Though it is clear that North Atlantic volcanism did occur at this time, its exact
role in events surrounding the PETM is ambiguous.
3.1.1. Tectonic History of the Central Basin
Fig. 3: Plate orientation during the opening of the Norwegian-Greenland Sea (modified from
Talwani and Eldholm (1977), showing Svalbard’s paleolocation and relative proximity to
Greenland and subsequent counterclockwise motion of Greenland relative to Svalbard. The red
and green colors are used to denote opposite sides of the spreading ridge which became the
north end of the Mid-Ocean Ridge system.
On Svalbard’s largest island of Spitsbergen (~39,000 km2), the majority of exposed strata
consists of late Paleozoic-Mesozoic, Devonian, or pre-Silurian deposits (Fig. 4). The largest
Cenozoic sedimentary deposit is the Van Mijenfjorden group which occurs in the central
depositional basin (the Central Basin). The Central Basin was likely opened during the initial
stages of the opening of the North Atlantic (Müller and Spielhagen, 1990) and became a foreland
12
basin environment during the formation of the West Spitsbergen Orogeny between the Late
Paleocene and the Miocene (Harland, 1969; Bruhn and Steel, 2003).
Fig. 4: Geologic map of Spitsbergen, showing location of Core BH
9/05 in red. Modified from Hjelle (1993) and Riber (2009). Core
BH 9/05 is part of the Van Mijenfjorden group, a Tertiary deposit
covering the Central Basin. During the uplift of the West
Spitsbergen Orogeny, the Central Basin became a foreland-basin
environment (Müller and Spielhagen, 1990).
13
3.1.2. Lithology of the Van Mijenfjorden Group
Core BH 9/05 is part of the Van Mijenfjorden group, which covers an area of 200 by 600
km2 in the center of the island (Riber, 2009; Nagy, 2005). The Van Mijenfjorden group is a
thick clastic sequence which spans the Paleocene and early Eocene (Nagy, 2005). The Paleocene
succession is approximately 700 m thick and comprises the Firkanten Formation, Basilika
Formation, Grumantbyen Formation, and the lower part of the Frysjaodden Formation (Bruhn
and Steel, 2003). The Paleocene formations are predominantly shales, siltstones, and sandstones,
with infrequent conglomerate and coal deposits. The Paleocene succession is thought to record
an overall transgressive trend from near-shore shallow water through a prograding delta
environment to an outer shelf environment (Riber, 2009; Sætre, 2011). The Eocene succession is
roughly 1400 m thick and is composed of the remainder of the Frysjaodden Formation, Batfjellet
Formation, and Aspelintoppen Formation (Bruhn and Steel, 2003). The Eocene sequence shows
an overall trend of regression (Bruhn and Steel, 2003).
The focus in this study is the Frysjaodden Formation, which in BH 9/05 is further divided
into two subunits: the Marstranderbreen Member and the Gilsonryggen Member. Western
deposits contain a third unit (Hollendardalen Fm) between the other two members, which
decreases in thickness eastward and eventually is no longer present. The Gilsonryggen Member
contains the PETM and is a predominantly shale unit interpreted to be a deep-water outer shelf
deposit (described in more detail in the following section) (Riber, 2009; Dypvik et al., 2011).
14
3.2 Core Description: Lithology
Core BH 9/05 was collected by the Store Norske Spitsbergen Kulkompani and provided to
the pACE group (paleo-Arctic Climates and Environments), part of the Worldwide Universities
Network (WUN) that aims to characterize Paleogene climate change. Core BH 9/05 is from
central Spitsbergen, near the Van Mijenfjorden inlet (Fig. 4), and samples a marine sedimentary
sequence spanning the upper Grumantbyen Formation (late Paleocene), Frysjaodden Formation,
and the lower Batfjellet Formation (lower-Eocene) (Riber, 2009; Cui et al., 2011). This study
focuses on the interval between 545 and 480 mbs (meters below surface) (Gilsonryggen
Member), which contains the onset and recovery of the PETM as identified by the total organic
carbon isotopic composition of the bulk rock (Cui et al., 2011).
15
Fig. 5: (left) Stratigraphic column of BH
9/05, modified from Riber (2009). The
most negative δ13C values occur between
536 and 519.7 m (Cui et al, 2011),
followed by a gradual recovery. An
interval of high siderite and pyrite is
present prior to and during the δ13C
minimum (Riber, 2009).
Core BH 9/05 likely represents a distal shelf depositional environment (Riber, 2009). The
lithology of the PETM interval (Fig. 5) consists mainly of organic-rich grey shales and siltstones,
with a conglomerate unit below 548.40 m. Between 548.40 and 535.00 meters there is a
moderately sorted silty clay unit containing agglutinated foraminifera and moderate bioturbation
(Riber, 2009). The section from 532.00 m to 498.00 m contains laminated shales and contains
16
the most negative values of the carbon isotope excursion (onset at 536 m) (Cui et al., 2011).
Bedded and concretionary siderite (FeCO3) is found between 535.00 and 532.00 m and above
498 m (Riber, 2009). Pyrite (FeS2) is not common in the core (<5%), but increases between
529.75 and 517.08 meters (Riber, 2009).
Dating of the core has been done by Charles et al. (2011) using variations in Fe and Mn
concentrations of bulk rock assumed to vary on an orbital scale, matched to a similar section
from Longyearbyen (characterized in Harding et al. (2011)) and scaled to two pre-existing agemodels (ODP Site 1263 (Röhl et al., 2007); ODP site 1266 (Murphy et al., 2010)). These ages
were then corrected using the measured age of a bentonite from the Longyearbyen section.
Sedimentation rates for the PETM interval are estimated at 21.5 to 27.4 cm/ka depending on
which age model is used, so that the CIE is recorded over roughly 49 meters. Other PETM
sections have much smaller recorded CIE sections (i.e. ~9.5 m (ACEX core; Sluijs et al., 2006),
~1.5 m (ODP 1263, 1265, 1262; McCarren et al., 2008), ~3 m (DSDP 549; Ravizza et al., 2001)
making this one of the highest resolution PETM records available.
17
Chapter 4
Analytical Methods
4.1 Osmium and Rhenium Geochemistry
4.1.1. Osmium and Rhenium Analysis
In order to obtain a Paleogene Os isotopic record, I analyzed NiS fire assay fusions and
Carius tube digestions to determine bulk rock 187Os/188Os ratios and Os and rhenium (Re)
concentrations (described below). The Os isotopic record was derived from these measurements,
after correction for radioactive in-growth of 187Os from 187Re. All age corrections assumed
closed system dynamics (i.e. that no Re or Os had left or entered the system after deposition),
which is crucial in obtaining an accurate estimate of the amount of 187Re decay and subsequent
calculations of 187Os.
In this study, I report all osmium (Os) isotopic ratios as 187Os/188Os. Both 187Os/188Os and
187
Os/186Os ratios are reported in the literature, and one can convert between the two using the
formula 187Os/188Os = 0.120343*(187Os/186Os) (Luck and Turekian, 1983).
187
Os/186Os is no
longer used because of 186Os’s potential for in-growth due to radioactive decay, even though it
both decays and is added over very long time scales, such that effects are negligible over millionyear time scales.
4.1.2. Preparation: Bulk Digestion and Carius Tube Comparison
Carius tubes are thick-walled glass tubes in which reactants are sealed, heated, and
subsequently digested. Carius tubes have been employed for Re-Os analysis in many studies
18
(e.g. Shirey and Walker 1995; Shen et al., 1996; Cohen and Waters, 1996; Selby and Creaser
2003). Compared to bulk digestions, the Carius tube method has the advantages of increased
ease of preparation through decreased preparation time and equipment and lower procedural
blank measurements (Meisel et al., 2003; Shirey and Walker, 1995). Unlike NiS bulk fusions,
Carius tubes also allow Re and Os to be measured on the same sample split, reducing the risk of
differences caused by sample heterogeneity. Comparisons between different bulk digestion
methods and Carius tube methods have found that Carius tube digestions of 187Os/188Os generally
agree with bulk measurements. Shirey and Walker (1995) reported <15% difference between
bulk and Carius tube measurements for 80% of samples, while Meisel et al. 2003 showed <10%
difference in reproducibility, indicating that the Carius tube method is nearly comparable to bulk
fusions in terms of reliability.
Previous Carius tube studies have used a variety of very oxidizing acid combinations as the
acting reagent (i.e. aqua regia, HF-HBr, H2SO4-CrO3, HF-BF3). Using aqua regia, Shirey and
Walker (1995) found that nearly all Re and Os contained within a silicate matrix were released.
Selby and Creaser (2003) compared the use of aqua regia and H2SO4-CrO3 Carius Tubes and
concluded that aqua regia was more likely to dissolve silicate minerals and therefore liberate
non-hydrogenous Os, while H2SO4-CrO3 only liberated hydrogenous Os. Since this study is
focused on the hydrogenous Os signal, a nitric acid — hydrogen peroxide mixture was used as
the reagent, being a strong oxidizer but less likely to dissolve silicate material since the nitric
acid-peroxide combination is not as strong as HF or Cr(VI). The principle advantage of this
digestion mixture is that OsO4 can be easily volatized from this solution, a feature that can be
exploited to facilitate rapid Os isotope analyses.
19
Carius tubes have also been used for isolated rhenium measurements. In Shirey and Walker
(1995), rhenium digestion in Carius tubes was compared to the established digestion method
using Teflon bombs. The authors concluded that there was no appreciable difference in the
measured Re concentrations for concentrations over 100 ppt (parts per trillion), which was
reinforced in Meisel et al. (2003). Therefore for the present study, in which all Re
concentrations were greater than 100 ppt, separate digestions were performed in Teflon bombs
only on a small subset of samples, while most digestions were done in Carius tubes.
4.1.3. NiS Fire Assay Bulk Digestions
Osmium geochemical analyses were conducted in a clean room on fusions of bulk rocks,
following the NiS fire assay method (Ravizza and Pyle, 1997). Prior to the fusions, five grams
of bulk rock were ground in an agate mortar and pestle to uniform consistency. The sample was
weighed, transferred to a Coors porcelain crucible, and spiked with an enriched 190Os solution
(sample spike: [Os] = 4.6 ng/g, 190Os/188Os = 186.7891). This sample-spike mixture dried
overnight and was covered with filter paper to prevent contamination from dust.
Fusions entailed combining 10 g sodium tetraborate, 0.2 g elemental sulfur, and 0.3 g nickel
powder flux with the spike-sample mixture. The flux mixture was prepared in clean containers
that had not previously come in contact with sample and shaken to mix. Upon homogenizing the
sample-spike mixture by hand, the flux was added to the sample-spike mixture and mixed
thoroughly. The covered crucibles (three to four at a time) were then placed into an oven at 1000
°C and heated for 75 minutes. Blanks were prepared along with all samples, combining 10g of
sodium tetraborate with a flux mixture consisting of 20 g sodium tetraborate, 0.6 g elemental
sulfur, and 0.9 g nickel powder (blank spike: [Os] = 0.96 ng/g, 190Os/188Os = 186.7891).
20
Upon removal from the oven, samples cooled for 60 minutes, and then were crushed. All NiS
beads were then picked from the crushed glass and small NiS beads were recovered. These
small sulfide nuggets are the result of an immiscible sulfide phase that forms during fusion. Os
is quantitatively extracted from the borosilicate melt into the sulfide phase during the fusion.
The NiS beads were dissolved in 100 mL 6 N ultrapure HCl at 100 °C for 24 hours. Upon
dissolution, solutions were filtered through 0.45 µm Millipore cellulose filter paper (25 mm
diameter); the filtrate retained the sample Os as insoluble particles and was saved for processing
preceding spectrometer analysis. The filtered solution was discarded.
4.1.4. Bulk Rhenium Acid Digestions
Rhenium geochemical analyses were conducted on a small subset of samples using acid
digestions of bulk powdered rocks. In this case, 0.5 g of sample powder was weighed in a
Teflon vial and spiked with an enriched 185Re solution ([Re] = 70.7 ng/g, 187Re/185Re = 0.05), to
which “inverse” aqua regia was added (5 mL of 6.2 N distilled HCL and 5 mL 8 N HNO3). The
vial was capped and digested on a hot plate overnight at 100°C.
Following digestion, the solution was centrifuged and the supernatant decanted and saved.
The pelletized residue was recombined with the remaining undigested residue in the Teflon vial.
The centrifuge tube was rinsed twice with 2 mL of 8 N HNO3, and the rinse combined with the
residue in the Teflon vial. The residue was further digested in 1 mL of H2O2 overnight at 100 °C.
This procedure was repeated once more; the remaining residue was finally digested overnight in
2 mL of concentrated HF, 1 mL of 8N HNO3, and 2 mL of deionized water at 100 °C. The final
digestate was combined with the existing solution and dried down in preparation for anion
exchange purification.
21
4.1.5. Carius Tube Digestion
Two Carius tube methods were used for this study. The first (used for measuring Re and Os
together) used a nitric acid-hydrogen peroxide combination as the reagent. Carius tube
digestions were conducted on approximately 0.5 grams of powdered rock sample. All Carius
tubes were cleaned in 110°C aqua regia (2:1 16N HNO3 and 12N HCl) and rinsed with high
purity 18 MΩ deionized water. Sample aliquots were spiked with enriched 190Os spike ([Os] =
4.6 ng/g, 190Os/188Os = 186.7891) and a 185Re spike ([Re] = 70.7 ng/g, 187Re/185Re is 0.05) and
combined in the Carius tube with 2.5 mL 4 N ultrapure HCl , 1 mL concentrated H2O2, and 25
mL 16 N ultrapure brand HNO3. After adding aqua regia, the Carius tube was frozen in a dry-ice
ethanol slurry to avoid production of CO2 and sealed with an acetylene torch. After sealing,
Carius tubes were placed in steel sleeves with screw-on threaded metal ends and heated over ~30
min to 250 °C; the digestions were subsequently conducted at 250 °C for ~18 hours. Following
digestion, samples cooled to room temperature over ~6 hours.
The second method (used to obtain isolated bulk Re measurements) was based on the
methods of Shirey and Walker (1995) and Selby and Creaser (2003), with aqua regia as the
reagent. Aqua regia Carius tubes are widely used and so were used in the current study for
comparison purposes. The proportions of HNO3 and HCl were those used in Selby and Creaser
(2003) (2:1 16N HNO3 and 12N HCl, with 9 mL used overall).
Approximately 0.5 g of sample was added to a Carius tube that had been previously rinsed
with MilliQ 18MΩ water and cleaned with boiling aqua regia. The sample powder was spiked
with isotope spike enriched in 185Re.
A series of freezings were applied as in Shirey and Walker (1995) to keep the nitric and
hydrochloric acids from reacting with each other and with the sample during the process. The
22
freezing was also necessary since aqua regia produces vapor whose pressure would make the
tube difficult to seal. First the previously spiked sample powder was frozen in a slurry of dry ice
and ethanol. Then 3 mL of reagent-grade 12N HCl was added and frozen immediately. Once
the HCl had frozen, 6 mL of reagent-grade 16N HNO3 was added and once again frozen. At this
point the tube was sealed and further treated like the osmium Carius tubes, in the manner
described above.
Once opened, the contents of the Carius tubes (which were only used for Re analysis) were
transferred to Teflon 15 mL Savillex beakers and dried down on a hot plate overnight in
preparation for Re anion exchange column chemistry. Two procedural blanks were prepared by
added 2.5 mL of 4N HNO3 and 1 drop of low level Re spike enriched in 185Re.
4.1.6. Re Anion Exchange Column Chemistry
The column chemistry procedure used is similar to previously published methods (Ravizza
and Paquay, 2008). For each ion exchange column, 1 mL of anion exchange resin (Biorad AG1X8 100-200 mesh) was used. The resin was first cleaned by running 10 mL 8 N HNO3 and 2 mL
of 4 N HNO3 through each column. The resin was then conditioned with 10 mL 0.5 N HNO3.
Dried samples were resuspended in 8 mL 0.5 N HNO3 and added to the column. The column was
washed three times with 5 mL 0.5 N HNO3 (waste discarded) and once with 2 mL 4 N HNO3
(waste discarded). Rhenium was eluted in 10 mL 4 N HNO3 and the elutant dried down on a hot
plate overnight.
23
4.1.7. Mass Spectrometry
Introducing osmium or rhenium into the mass spectrometer via a gas carrier (‘sparging’) is
done following the procedures of Hassler et al. (2000). Sparging involves oxidizing Os to its
highest valence state and carrying it from solution via a gas stream to the mass spectrometer’s
plasma torch. One of sparging’s main advantages over traditional nebulizers is its reduction of
Os loss during introduction (Hassler et al., 2000).
In this study, in preparation for sparging, all samples were transferred to Teflon beakers.
Carius tubes were opened and their liquid contents transferred directly, while 1000 µL of HNO3
and 100 µL of H2O2 were added to the bulk samples. All Teflon beakers were sealed and placed
in an oven at 100 °C for 45 minutes. After being removed from the oven, the Teflon beakers
were placed on ice in order to condense vapors. When ready to sparge, the beakers were fitted
with sparge caps and diluted with MilliQ water if initial volumes were low, to ensure good
contact between the sparge-cap stem and solution and to minimize acid vapor flow into the
machine.
Subsequently, samples were sparged directly into an Element2 high-resolution magnetic
sector inductively-coupled plasma mass spectrometer (ICP-MS). Argon gas was bubbled
through each sample solution, carrying osmium directly to the plasma as OsO4. Measurements
were collected for the seven Os nuclides via peak hopping (changing between peaks of specified
isotopes). Sampling was done over 30 runs, monitoring 10 points across each peak. An in-house
standard (LOsST) and a gas blank were run at the beginning and end of every run and after every
6 samples to ensure that there was no carryover between samples. During the time of this study,
measurement of the LOsST solution yielded 187Os/188Os ratios that were typically within 1% of
the accepted value. Re was measured by conventional solution nebulization using a Meinhardt
24
nebulizer and Scott-type spray chamber. Data were subsequently corrected for instrumental
mass bias, 187Os/188Os ratios were corrected for influence of minor isotopes in the spike solution,
and concentrations determined through isotope dilution.
4.1.8. Blank Corrections
Blank Os concentrations were typically on the order of ~1pg or less. Blank corrections for
both bulk fusions and Carius tubes were very small, amounting to <0.4% of the uncorrected Os
concentration value and 0.02% of the 187Os/188Os ratio. Blank corrections for Re were similarly
small, ~1% for acid digestion in Teflon bombs, <3% for nitric Carius tubes, and 5-9% for aqua
regia Carius tubes.
4.2 Bulk Element Analysis
Samples were analyzed for their major oxide compositions using bulk element
spectrochemical analysis. Samples were prepared via lithium metaborate (Li2BO2) fusion, in
which 0.25 g powdered samples were combined with 1.25 g lithium metaborate flux and heated
at 1050°C in an ultrapure graphite crucible for 20 minutes (Totland et al., 1992; Van Loon and
Parissis, 1969). The melt was dissolved in 0.8 M nitric acid, stirred for 30 min, and diluted to
250mL prior to analysis by inductively coupled plasma emission spectroscopy (ICP-AES).
Certified rock standards, prepared in the same manner as the samples, were run with each sample
set three times. Standard accepted values, measured values, averages, and standard deviation for
each oxide is listed in Table 1.
25
4.3 Numerical Model Description
A time-dependent numerical model was created to simulate mass and isotopic fluxes within
the global and regional Os cycles. The model contained two main components: a global ocean
(including riverine, hydrothermal, and extraterrestrial inputs and a sedimentation output), and a
local semi-restricted basin (including riverine and global ocean inputs and a sedimentation
output). Three model simulations were used to examine the Os isotope data over the PETM. The
first recreates the recorded 187Os/188Os record through changes in the Os river flux. The second
recreates the recorded 187Os/188Os record through changes in the average river 187Os/188Os ratio.
The third recreates the recorded 187Os/188Os record through changes in the Os oceanic exchange
flux. These simulations were used to answer specific questions: (1) What input sources could be
responsible for the observed changes in the 187Os/188Os ratios and Os concentrations? and (2)
What are the required magnitudes of any source changes?
To describe the system, an equation for the change in the reservoir isotopic system based on
input and output fluxes was derived, similar to the one used for the strontium isotope system
(Kump, 1989):
(
(
)
)
(∑
∑
(∑
∑
)
)
(
(
)
)
(∑
∑
(Eq. 4), where
26
(∑
∑
)
)
(
)
(∑
∑
)
∑
And similarly for the basin:
(
)
(
)
(∑
∑
(∑
∑
)
(
)
27
)
(∑
∑
)
Where
As in Kump (1989), and unlike the simplified equation typically used for calculating values, the
(∑
)
∑
term is necessary since 188Os is not a dominant nuclide (which an
absence of the correction term implies). For the Sr system, total Sr concentration is treated as
directly proportional to each of the Sr isotopes since Note that since Os leaves each reservoir
with the same isotope ratio as the water (no fractionation upon deposition), in each case the
deposition term (Fdep, which would normally be subtracted) becomes zero.
I then used the system dynamics box-modeling program STELLA (v.6.1) as a tool to solve
the equations above in a time-dependent manner. Initial fluxes, Os concentrations, and
187
Os/188Os ratios of sources were assumed to be equivalent to their modern counterparts where
reasonable (see Appendix B). The 187Os/188Os data from cores BH 9/05 and DSDP 549 were
interpolated using piecewise linear functions in order to create evenly spaced 187Os/188Os and
time derivative curves for each core. The interpolated data were then assumed to represent the
global ocean 187Os/188Os (DSDP 549) and the local basin 187Os/188Os (BH 9/05). The data
thereby served as the primary constraint in evaluating changes in the Os cycle over the PETM.
28
Chapter 5
Results
5.1 Osmium and Rhenium Concentrations and 187Os/188Os
Measurements
Osmium and rhenium data are listed in Appendix A. The general features of the 187Os/188Os
record over the depth interval sampled in this study are: (1) an initial 187Os/188Os value ~0.6, (2)
a sharp drop to 0.4 at ~537 meters (~55.8 Ma) prior to the onset of the CIE, (3) a subsequent
peak value of 0.8238 at 533.07 meters during the CIE onset, and (4) a gradual decline back to 0.6
(Fig. 6b). Replicate measurements, as well as measurements of multiple samples near 537
meters, verify the presence of the 187Os/188Os minimum.
Measured Os concentrations in the solid phase, which range from 113 parts per trillion (ppt)
to 218 ppt, generally rise over the depth interval measured. The maximum Os concentration
occurs ~516.6 mbs, during the recovery of carbon isotopic composition to pre-CIE values.
Rhenium concentrations in the solid phase increase from 2 to 5 ppb during the onset of the CIE
and increase further to a first maximum near 13 ppb at ~520 m, during the interval of the most
negative carbon isotope values. A second maximum of ~17ppb exists at 514.75 m during the
initial stages of the CIE recovery, but the robustness of this point is unclear, as it represents a
60% increase from prior values. The Re concentration subsequently decreases and reaches
values around 2 ppb at depths < 510 meters. In the beginning of the record, there is some fine
structure in Re concentration that coincides with changes in both Os concentrations and
187
Os/188Os. Between 540 and 534 mbs, Re concentrations (Carius tube digestions) rise from ~2
29
to 5 ppb; over this same interval, Os concentrations (Carius tube digestions) rise from ~130 to
170 ppt and 187Os/188Os falls from ~0.65 to ~0.4. Interestingly, the consistency of 187Os/188Os
measured in five samples at the 187Os/188Os minimum (0.4 – 0.5) over a range of Re
concentrations (2 – 5 ppb) suggests that the age corrections made are accurate.
Because two digestion methods were employed, it is important to examine any systematic
differences in data that may arise due to methodology. Accordingly, I compare the Os and Re
geochemical data derived from the Carius tube and bulk fusion digestions (Fig. 7). Results show
that osmium isotopic composition is largely unaffected by the choice of digestion method, as the
slope of the fit is essentially 1. Though the non-zero intercept (0.025) suggests a small offset
between the two techniques, it is much smaller than the size of the signal that I am interpreting.
The agreement between the digestion methods is also fairly good for Os and Re concentrations.
(regression slopes are 0.78 and 0.86, respectively), with variations mainly due to powder
heterogeneity (i.e. because the abundances of Re and Os are so low, sample splits may not be
exactly identical). Despite slight systematic methodological discrepancies, all three datasets
appear strong in light of the signals I am interpreting.
30
Fig. 6: a) The carbon isotope curve from BH 9/05 (Cui et al., 2011) b) Age-corrected osmium
isotope ratios (187Os/188Os) over the PETM, c) Osmium concentration over the PETM, in parts
per trillion (ppt). d) Rhenium concentration over the PETM, in parts per billion (ppb). Blue
diamonds indicate bulk measurements, while the red circles represent Carius tube measurements.
The gray line indicates the depth of the 187Os/188Os minimum while the dashed line marks the
depth of the 187Os/188Os maximum. Note that age-corrected 187Os/188Os ratios are considerably
less variable than measured 187Os/188Os ratios (Appendix A) and that initial 187Os/188Os ratios
determined by different methods typically agree within <0.03 187Os/188Os units.
31
Fig. 7: a) Comparisons of measured bulk vs. Carius tube 187Os/188Os ratios. Agreement in
187
Os/188Os ratios appears to be the best overall, with a best-fit slope of 0.988. b) Comparisons
of bulk vs. Carius tube measurements for osmium concentration. Agreement in this panel is the
lowest, with a correlation value of 0.59. c) Comparisons of bulk vs. Carius tube measurements
for Re concentration. Re concentrations have a relatively good correlation of 0.8649 and an R2
value of 0.94. In each panel a 1:1 line (blue dashed line) is included for reference. Differences
in Os and Re concentrations are attributed to sample powder inhomogeneity. Note that in order
to make a direct comparison of Os values, Os measurements are not age corrected, so that
differences in Re will not affect the result. See Appendix A for table of values.
32
5.2 Bulk Oxide Analysis
The elemental compositions of bulk rocks over the interval 544.09 to 469.6 mbs are shown in
Table 1. The rocks analyzed have major element chemistries that are comparable to a typical
shale, with average Al2O3 (17.1%), Na2O (1.1%), CaO (0.4%), K2O (2.9%), SiO2 (57.5%), and
MgO (1.5%) compositions that vary little downcore. There is some variability in the Fe2O3
concentration of the rocks, but this variability does not appear to be systematically related to
depth or to pyrite abundance.
33
Table 1: Major oxides reported as weight percents, from bulk oxide analysis. Average oxide
measurements for BH 9/05 appear to be closer to an average shale than to a typical gray shale,
exceptions being lower than average levels of CaO and SiO2.
Sample
Depth
Al2O3
(%)
BaO
(%)
CaO
(%)
Fe2O3
(%)
K 2O
(%)
MgO
(%)
MnO
(%)
Na2O
(%)
P2O5
(%)
SiO2
(%)
SrO
(%)
TiO2
(%)
469.58
15.1
14.1
15.9
16.3
17.3
17.2
16.3
16.9
16.1
17.4
18.7
18.3
18.1
14.5
17.9
18.0
17.2
17.4
18.1
18.5
17.4
19.9
17.4
15.5
15.0
18.8
19.1
17.1
0.05
0.04
0.05
0.05
0.05
0.06
0.05
0.05
0.05
0.06
0.06
0.06
0.05
0.04
0.05
0.05
0.05
0.06
0.06
0.05
0.05
0.06
0.05
0.04
0.05
0.06
0.07
0.1
0.68
0.62
0.29
0.25
0.31
0.49
0.28
0.25
0.55
0.45
0.39
0.46
0.62
0.36
0.43
0.24
0.27
0.24
0.25
0.20
0.37
0.42
0.29
0.40
0.44
0.42
0.36
0.4
10.6
7.65
5.44
5.38
7.21
6.29
5.41
5.51
7.70
6.70
5.73
7.23
7.83
14.5
5.61
6.66
6.63
7.36
7.46
8.11
9.32
6.99
5.67
6.99
12.9
6.70
7.40
7.4
2.50
2.22
2.54
2.80
2.73
3.08
2.78
2.83
2.71
3.02
3.36
3.23
2.93
2.20
2.62
3.01
2.68
2.88
2.94
3.25
3.03
3.19
2.66
2.31
2.69
3.31
3.70
2.9
1.56
1.47
1.44
1.42
1.46
1.60
1.46
1.51
1.55
1.49
1.56
1.59
1.65
1.24
1.55
1.49
1.45
1.47
1.42
1.45
1.52
1.58
1.57
1.49
1.53
1.61
1.75
1.5
0.26
0.18
0.04
0.03
0.04
0.06
0.04
0.04
0.12
0.12
0.05
0.09
0.16
0.05
0.07
0.02
0.05
0.06
0.07
0.09
0.14
0.06
0.02
0.03
0.04
0.04
0.04
0.1
1.01
1.06
1.10
1.09
1.08
1.20
1.06
1.07
1.10
1.08
1.15
1.09
1.02
0.94
1.06
1.13
1.16
1.05
1.06
1.01
0.97
0.96
1.41
1.47
0.98
1.02
1.02
1.1
0.34
0.34
0.17
0.16
0.19
0.33
0.17
0.15
0.29
0.25
0.26
0.27
0.28
0.20
0.24
0.16
0.16
0.15
0.15
0.14
0.13
0.21
0.14
0.16
0.17
0.21
0.17
0.2
55.5
60.7
63.3
62.8
57.3
59.3
63.0
61.8
58.3
59.2
58.5
56.3
55.1
48.9
57.6
58.1
59.6
56.7
55.3
55.3
54.5
54.7
58.8
58.0
52.0
55.9
55.3
57.5
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0
0.0
0.7
0.66
0.74
0.72
0.79
0.79
0.72
0.77
0.75
0.77
0.83
0.82
0.85
0.76
0.92
0.84
0.82
0.79
0.79
0.74
0.72
0.92
0.91
0.88
0.73
0.94
0.89
0.8
Al2O3
(%)
BaO
(%)
CaO
(%)
Fe2O3
(%)
K 2O
(%)
MgO
(%)
MnO
(%)
Na2O
(%)
P2O5
(%)
SiO2
(%)
SrO
(%)
TiO2
(%)
15.45 0.02 10.86 10.83
0.63
6.37
0.17
2.20 0.14
53
473.51
475.05
478.05
481.05
484.05
487.05
490.05
492
495.18
498.45
501.4
511.05
504.05
507.75
513.05
516.05
519.05
521.45
524.45
527.45
530.99
534.05
535.76
538.72
541.7
544.09
Average
Quality
Control
Std
W-2
Ref.
Actual
Value
34
0 1.06
Total
88.3
89.1
91.0
91.0
88.5
90.4
91.3
90.9
89.2
90.6
90.6
89.5
88.6
83.7
88.1
89.7
90.1
88.2
87.6
88.9
88.2
89.0
88.9
87.3
86.6
89.0
89.8
99
W-2 (1)
W-2 (2)
W-2 (3)
Average
15.5
15.5
15.7
15.57
0.02 10.8 10.9 0.6
0.02
11 10.7 0.6
0.02 11.3 11.2 0.6
0.02 11.03 10.93 0.63
6.56
6.4
6.72
6.56
0.17
0.17
0.18
0.17
2.25 0.1
2.2 0.2
2.21 0.1
2.22 0.14
53
52
54
53
0 1.1
0 1.08
0 1.1
0 1.09
Stand.
Dev.
0.12
0.00
0.25
0.25
0.02
0.16
0.01
0.03
0.01
1.2
0
0.01
Al2O3
(%)
BaO
(%)
CaO
(%)
Fe2O3
(%)
K 2O
(%)
MgO
(%)
MnO
(%)
Na2O
(%)
P2O5
(%)
SiO2
(%)
SrO
(%)
TiO2
(%)
18.78
1.29
7.18
3.68
2.19
1.19
0.16
62
10.7
0.69
5.13
2.1
1.12
0.39
0.2
Average
Shale
Average
grey
Shale
65
100
98
102
100
0.99
0.61
The chemical index of alteration (CIA), defined as (Nesbitt and Young, 1982):
[Al2O3/(Al2O3+K2O+Na2O+CaO)]
(Eqn. 5)
was calculated using the oxide weight percent data, converted into molal units (moles/kg
rock). The CIA therefore compares the “immobile” alumina content of a rock to the total major
cation inventory. Values of CIA close to one are classified as highly weathered, while CIA
values less than one are less weathered. Calculated values of CIA between 544.09 to 469.6 mbs
are shown in Fig. 8 for two cases: the bulk rock and the bulk rock from which the kaolinite
contribution has been mathematically removed (kaolinite contents obtained from Riber (2009),
based on XRD analysis). The most significant features of both records are the decrease, and
subsequent increase, in CIA that occurs contemporaneously with the fall and rise of 187Os/188Os
(Fig. 8).
35
Fig. 8: Record of increased weathering over the PETM. a) (top): CIA measurements for BH
9/05, including both bulk CIA and residual CIA (CIA of the remainder of the material sans
kaolinite), with the 187Os/188Os record for comparison. Dashed line = depth 533.07 m. Note that
this is where the maximum occurs in 187Os/188Os ratios and kaolinite ratios. Residual CIA values
were calculated using a mass balance equation and interpolated at depths of existing kaolinite
measurements from Riber (2009). b) (bottom): kaolinite ratios in BH 9/05 for comparison, from
Riber (2009). Kaolinite ratios, the CIA, and 187Os/188Os ratios are weathering indicators.
36
Chapter 6
Discussion
6.1 BH 9/05 as a Basin Signal
In contrast to other cores that have evaluated the Os record during the PETM, I hypothesize
that Core BH 9/05 represents a basin signal, influenced by both the global ocean response to the
PETM as well as more local factors such as riverine input. Though the observed 187Os/188Os
trend is similar to records obtained at other sites over the PETM, the 187Os/188Os values obtained
at BH 9/05 prior to the PETM are much higher than “open ocean” 187Os/188Os ratios. Previouslymeasured open-ocean 187Os/188Os records come from DSDP Sites 549 and 213 (located in the N.
Atlantic and the Indian Ocean, respectively) (Ravizza et al., 2001) and sites from Dababiya
Egypt and Zumaya Spain (Schmitz et al., 2004). However, the site from Dababiya was affected
by Re and/or Os mobility during diagenesis, and the record from Zumaya showed no clear
evidence of a peak near the PETM, possibly due to influences from North Atlantic volcanism.
Site 213 does not capture the onset of the PETM due to depositional hiatus. Site 549 is the most
complete core; therefore it is used as the “open or global ocean” 187Os/188Os record.
Comparison of the two records shows that BH 9/05 187Os/188Os values are nearly twice the
values of those found at Site 549 (Ravizza et al., 2001); the latter’s steady state, pre-event
187
Os/188Os ratios were 0.3633 and increased to 0.4387 at the maximum during the PETM. In
contrast, the initial pre-PETM 187Os/188Os ratio at BH 9/05 began at ~0.6 and rose to a maximum
of 0.8238. The difference between the two records is not consistent with the hypothesis that the
37
ocean 187Os/188Os is homogenous, assuming that core BH-9/05 records global seawater (Ravizza,
1993).
There are several possible explanations for the difference between the 187Os/188Os records of
BH 9/05 and DSDP 549. First, since BH 9/05 does have a much higher estimated sedimentation
rate and thus greater temporal resolution, it is possible that Site 549 does not capture the full
magnitude of the 187Os/188Os ratio increase. However, that would not explain the differences in
the values of the pre-event steady state 187Os/188Os ratios between the two sites. It is also
possible that Site 549 and/or BH 9/05 do not record a primarily hydrogenous Os signal and are
influenced by detrital material (average continental crust 187Os/188Os ratio ~1.2-1.4 (Esser and
Turekian, 1992)). Site 549 sediments are comprised of nannofossil oozes (Ravizza et al., 2001),
while core BH 9/05 contains clay-rich, near-shore sediments. Though I cannot definitively rule
out a detrital component in BH 9/05, the similarity between the two records implies that either
the detrital fraction is only a small portion of the total sample, or that the detrital and
hydrogenous portions have similar 187Os/188Os ratios (which is unlikely, since the detrital portion
is typically much more radiogenic than the hydrogenous signal). Therefore, it seems most likely
that the detrital fraction does not contribute much Os so that the Os measured is predominantly
hydrogenous.
Since the Os isotope record at Site 549 is also thought to be a hydrogenous signal (Ravizza et
al., 2001), another possibility is that both sites record hydrogenous signals but one or both is not
a global seawater signal. In this case, it seems most likely that Site 549 (a deep ocean site far
removed from terrestrial influences) records the global ocean; 187Os/188Os ratios at Site 549 are
consistent with seawater 187Os/188Os in the Paleogene ocean (Peucker-Ehrenbrink et al., 1995).
By contrast, the nearshore BH 9/05 is likely to be affected by terrestrial influences, such as
38
riverine input. I propose that BH 9/05 records a signal consistent with a semi-restricted deep
water sedimentary basin.
Fig. 9: Results from the global ocean model for changes in the river Os flux. This model was
made to see if both core records could be caused by nearly the same input fluxes. However,
since the required Os kg/yr input for BH 9/05 is more than twice that of DSDP 549, the
conclusion is that they cannot both be recording the same signal (since Os is well-mixed in the
ocean (Ravizza and Zachos, 2003), the two signals would not be influenced by different sources,
the other possibility), and so they cannot both be an ocean Os signal.
39
Fig. 10: Results from the global ocean model for changes in the river 187Os/188Os ratio. Changes
in Rr causing BH 9/05 would require a global isotope ratio higher than 2. While this is not
impossible, it would require a higher crustal abundance of radiogenic material exposed at the
surface. The question arises whether such large-scale lithology exposure changes can occur over
such a short geological time. Thus the conclusion of this study is that the observed 187Os/188Os
records were most likely not caused by a change in the global Rr alone.
Numerical simulations of the Os cycle (Fig. 10, Fig. 11; summarized in Appendix C)
suggest that substantially different Os inputs are required to explain 187Os/188Os ratios at Site 549
and BH 9/05. For instance, modeling the change in the 187Os/188Os record as a change in river Os
flux requires a riverine Os flux increase of 41% for DSDP 549 compared to a 116% increase for
BH 9/05. Similarly, modeling a change in global 187Os/188Os ratio requires a much greater
riverine 187Os/188Os increase for BH 9/05 (55%) compared to DSDP 549 (33%). Both of these
scenarios require significantly a different global environment for DSDP 549 than those implied
at BH 9/05. My modeling results indicate that 187Os/188Os ratios at BH 9/05 cannot be explained
by the same Os cycle as Site 549. I therefore contend that BH 9/05 is strongly affected by local
processes.
40
6.2 Evidence for Enhanced Chemical Weathering over the PETM
The Os isotope record derived from BH 9/05 provides considerable evidence for enhanced
chemical weathering over the PETM. The Os isotope record derived from BH 9/05 exhibits a
large 187Os/188Os increase leading to a maximum value during the onset of the CIE and a
subsequent decrease during the recovery period. The substantial increase in 187Os/188Os ratios is
indicative of either a relative increase in the mass flux of radiogenic Os to the basin (compared to
unradiogenic sources), or an increase in the 187Os/188Os ratio of the weathering flux to the basin.
These options are consistent with a shift in terrestrial patterns during the PETM.
Other indices of weathering also indicate a potential increase in weathering. An increase in
kaolinite occurs prior to and during the onset of the CIE in many marine sites worldwide (i.e.
Clechenko et al., 2007; Gibson et al., 2000; Robert and Kennett, 1994). In BH 9/05, the
maximum 187Os/188Os coincides with both the kaolinite/(kaolinite+chlorite) and the
kaolinite/(kaolinite+illite) maxima at 533.07 m during the onset of the CIE (Fig. 8) (Riber,
2009). Though kaolinite can be formed under both detrital and diagenetic conditions, Riber
(2009) and Dypvik (2011) concluded that the majority of the kaolinite in BH 9/05 was detrital in
origin.
Another proxy for weathering, the chemical index of alteration (CIA), quantifies the degree
to which a material is weathered based on the ratio of immobile (Al2O3) to total cations (Nesbitt
and Young, 1982). In BH 9/05, the CIA over the PETM interval (calculated from bulk oxide
molar percentages) has values ranging from 0.73 to 0.78, indicating that material deposited
during the PETM was intensely weathered. Additionally the CIA has its maximum value of 0.78
at 532.48 m during the onset of the CIE, just after the 187Os/188Os and kaolinite maxima
(533.07m). Though the CIA is indicative of weathering intensity not weathering rate, and does
41
not distinguish between present and past episodes of intense weathering, the nearly coincident
timing of highly weathered material associated with the CIA increase and the onset of the CIE,
187
Os/188Os increase, and kaolinite increase suggests that the three proxies are closely linked,
either through an increase in the flux of detrital material to the basin due to increased terrestrial
weathering or through similar mechanistic effects from changes in the local environment.
It is useful to decouple the influence of kaolinite from the bulk CIA signal in order to control
for variations in kaolinite content, since kaolinite has the maximum CIA value of 1 (Nesbitt and
Young, 1982) and therefore can strongly influence the CIA. To remove the influence of
kaolinite, a mass balance calculation was performed for the material to assess the CIA of the
non-kaolinite portion, using kaolinite contents obtained from XRD analysis (Riber, 2009). The
resulting curve, termed the “residual” CIA, is shown in Fig. 8. The curve exhibits a strong
similarity to the 187Os/188Os curve and displays a more pronounced peak and subsequent
drawdown than the initial measured CIA values. This suggests the curve is robust and may be
indicative of an even stronger effect than the initial bulk CIA curve would indicate.
.
42
6.3 The 187Os/188Os and CIA Maxima: Weathering vs. Lithologic
Change
The Os isotopic composition of the ocean, or any other water body receiving Os influx, is
sensitive to a change in the composition of the source. This is also true of the CIA, suggesting
another possibility: that the maxima in 187Os/188Os and CIA reflect a shift in the source
weathering material rather than enhanced weathering conditions. Over the PETM, this scenario is
not likely feasible on a global scale for several reasons, as addressed in Ravizza (2001). First,
according to our current understanding of the weathering process, causative factors of increased
physical and chemical weathering (i.e. increased precipitation and increased temperatures) are
present during the PETM (Zachos et al., 2001; Pagani et al., 2006), implying weathering should
have also increased. Second, a predominant change in source as opposed to increased
weathering would require a large shift in exposed lithology to a larger proportion of radiogenic
sources, which implies a huge amount of weathering and/or tectonic activity over a very short
time scale. While not impossible, it is questionable whether such a change on the global level
could occur over the brief hundred-thousand year time scale of the PETM (Röhl et al., 2007).
However, such changes might be likely over the smaller local scale of the basin.
To assess the possible contribution of a source change, modeling was performed (Appendix
C). In this simulation, I coupled a terrestrially influenced basin (~800 km3) to the global ocean,
whose Os cycle is constrained by the data from Site 549. Carbon isotope data from Site 549 and
BH 9/05 were used to correlate the records under two possible scenarios. Initial steady-state
187
Os/188Os values prior to the PETM were generated by an assumed lower river Os flux,
assuming lower continental weathering before the uplift of the Tibetan Plateau as proposed by
Pegram et al. (1992). There is no evidence for reduction in hydrothermal or extraterrestrial
43
fluxes (the unradiogenic end-members) at this time, so they are assumed to remain constant over
the interval.
The resulting “basin-global ocean” model was used to examine the following scenarios to
explain the BH 9/05 187Os/188Os record:
1. Increased river Os input
2. Increased river 187Os/188Os ratio (from a change in source)
3. Changes in the oceanic exchange flux as drivers for the 187Os/188Os record.
Modeling results (Fig. 11) show that in the case of increased river Os input, global riverine
Os flux would have had to increase by 45-50% (to 54.6-56.3 kgOs/yr), while a small river
entering the basin would have to triple its pre-PETM amount (to 0.47 kgOs/yr). For the case of
increased river 187Os/188Os ratio, two scenarios were examined: one in which both the ocean and
basin were driven by changes in riverine 187Os/188Os, and another in which only the basin was
driven by riverine 187Os/188Os (examining the case of local source changes) (Fig. 12, Fig. 13).
For all models, Os residence time in the basin is assumed to be much shorter than that of the
ocean (~11 yrs). Results for the first scenario showed that global average riverine 187Os/188Os
would have to increase to ~2, while basin riverine 187Os/188Os would have to increase to ~2.8.
Results for the second case (the case of local source change) were similar. The last option
examined the case where basin 187Os/188Os ratios were explained by a decrease in the exchange
flux with the ocean (since open ocean 187Os/188Os at ~0.3-0.44 is less than the basin 187Os/188Os
over this interval). Results for this scenario showed that the exchange flux would have to
decrease by a factor of ~3 to produce the required basin 187Os/188Os maximum.
44
Fig. 11: Results from the basin-ocean model for changes in the river Os flux. Note that in this
scenario, the Arctic river Os shows a much greater proportional increase than the global river
average, fitting with both a greater temperature change in response to stimuli (i.e. Sluijs et al.,
2006) and increased precipitation at the poles (Pagani et al., 2006), both of which would serve to
increase the total amount of Os reaching the basin from the river.
45
Fig. 12: Results from the basin-ocean model for changes in the basin river 187Os/188Os ratio.
Note that this scenario would require a maximum basin river Rr value near 3. While this cannot
be completely ruled out, it would require changes to much more radiogenic lithology at the
surface and back within the span of the PETM, and the timing of the increase and maximum
187
Os/188Os ratio aligning with the increase and maximum kaolinite (another indicator of
weathering) to be coincidental. Therefore, while changes in Rr influencing the basin Os record
cannot be ruled out entirely, a sole Rr cause is unlikely.
46
Fig. 13: Results from the basin-ocean model for changes in both the Arctic River and the global
average 187Os/188Os ratio. To achieve a global river average 187Os/188Os ratio near 2, the exposed
lithology would have to be much more radiogenic than today. While I cannot eliminate changes
in Rr as a possible influence on the Os record, it is questionable whether such large-scale
changes could have occurred over such a short geological time, and therefore unlikely that Rr
changes were responsible for the entire change in the Os record.
47
These modeling results imply that the increase in 187Os/188Os recorded in BH 9/05 over the
PETM is most likely due to an increase in river Os flux instead of a change in riverine
187
Os/188Os or a decrease in the exchange flux. In considering the exchange flux possibility,
there is no evidence for a decrease in oceanic input to the basin during the onset of the CIE;
instead, the basin shows evidence of a seawater increase (Riber, 2009; Harding et al., 2011). So
this option is not as likely. Second, an overall increase in global river flux seems the most likely
since the magnitude of river flux increase required remains well below the current riverine Os
flux (350 kg Os/yr; Levasseur et al, 1999). In contrast, a global riverine 187Os/188Os would have
to increase by 38% while basin riverine 187Os/188Os would have to double its initial value to 2.8
(initial 187Os/188Os values are assumed to be average continental crust ~1.4) (Ravizza and
Zachos, 2003). Most rock types have 187Os/188Os values near or below 1; some granites and
black shales can have values over 2 (Chen et al., 2006 Ravizza 1993), but they only cover a
certain proportion of the exposed lithology at any time (Peucker-Ehrenbrink and Ravizza, 2000).
A global riverine 187Os/188Os value near 2 would require a greater proportion of exposed cratonic
shields and/or black shale units. Again, this change to more radiogenic riverine input implies
large-scale tectonics that are questionable on the global level over the brevity of the PETM
interval, thus making a change to more radiogenic sources unlikely on the global scale.
However, we must consider that the global and basin levels may not be driven by the same
processes (i.e. the global increase in 187Os/188Os ratios may be a result of increased total riverine
Os flux, while the basin 187Os/188Os signal may be a result of a change in source to a more
radiogenic unit). A change in the basin riverine 187Os/188Os to values near ~2.8 again requires a
change in source lithology but this scenario may be more plausible for the basin than for the
global scale, as underlying geological units are largely shale (which can be very radiogenic) and
48
so are potentially better able to contribute to an increased riverine 187Os/188Os ratio (Harland,
1969). However, basin riverine 187Os/188Os values are still very radiogenic compared to current
crustal averages, and would require the weathering of a very radiogenic source. While modeling
results do not definitively rule out a source change as a possibility, a change in the total riverine
Os seems more likely.
6.4 The Potential Influence of Diagenesis
Diagenesis (i.e., post-depositional alteration) can change the initial isotopic composition of
the sediment. In our system, we are most concerned about Re mobility, as both the loss and gain
of Re in sediments will affect the age correction applied to the Os isotope data. A viable
hypothesis, therefore, is that the 187Os/188Os peak is an artifact of poor age correction due to Re
mobility during diagenesis. There are two likely processes to consider here; the first is Re loss
from sediment and the second is asynchronous Re addition (i.e., addition of Re to a particular
sediment that occurs at a different time than Os deposition).
As mentioned earlier, the peak in 187Os/188Os ratios corresponds with a Re minimum. If I
hypothesize that the Re data in this interval are affected by post diagenetic loss, then I must
assume that my 187Os/188Os data are under-corrected. The carbon isotope curve does not display
any features coincident with the Re record, suggesting that any diagenetic effect is exclusive to
rhenium. Rhenium loss can occur as oxidation state changes in the surrounding environment
(Yamashita et al., 2007). Rhenium is readily partitioned into sediments under suboxic and
anoxic conditions but has a smaller partition coefficient in oxic environments (Crusius et al.,
1996). According to this hypothesis, then, Re loss requires a shift to oxic conditions (Fig. 14).
49
Fig. 14: Hypothetical scenario of the case of rhenium loss. Figure a) [left] is the actual Re
record from BH 9/05, while b) is a close-up of the circled region with hypothetical pre-loss
estimated Re values.
Calculations demonstrate that a large environmental Eh change is required, using the
hypothetical ‘pre-loss’ Re values (Fig. 14b, open diamonds) and using the partitioning-Eh
relationship (
( )
(
)
(
)) between 187Re/188Os and Eh determined by
Yamashita et al. (2007). An Eh change of +0.2 – 0.3 V would be required to produce the desired
effect on Re. Yet there are a few lines of evidence that argue against such a shift to oxic
conditions. First, abundant siderite nodules/laminations are present between 535 and 532 mbs
(Riber et al., 2009). Siderite typically forms under reducing conditions, under which rhenium is
concentrated in sediments. Second, between 535 and 532 mbs, the δ13C record decreases,
implying an increasing amount of organic carbon input to the sediment, which would also make
the sediment more reducing. So a shift to more oxic conditions is unlikely over this interval.
To quantify the degree to which Re loss on the order of 80-90% could affect the 187Os/188Os
record, I recalculated the Os age corrections to reflect the ‘lost’ Re’s contribution to 187Os. In
this scenario, I assume that Re was initially deposited with concentrations increasing steadily
50
from 5.09-8.89 ppb and liberated 16.4 ka later during diagenesis as the basin became more oxic.
The calculations show that such as process would change the 187Os/188Os curve by less than 1%.
I therefore contend that the age-corrected 187Os/188Os peak is not a result of diagenetic Re loss.
If the Re minimum was incurred by a change to oxic conditions at an even later time (where the
higher Re would have longer to decay, and thus contribute more 187Os unaccounted for in the age
correction), this becomes a possibility. But a later liberation of Re seems unlikely because 1)
only a small interval of the Re record would have been affected, and thus it is questionable why
the rest of the Re record remained intact, and 2) sedimentation rates are high over this interval
(Charles et al., 2011), making it increasingly difficult for changes in the overlying water column
to affect the sediment.
A second possibility to consider is that rhenium was added to the sediment after Os was
deposited, which would result in an overcorrected 187Os/188Os record. Rhenium is typically
assumed to have been emplaced at or shortly after time of deposition; however, rhenium can also
be incorporated in diagenetic minerals formed under reducing conditions (McCandless et al.,
1993). Two of the main diagenetically-formed minerals of concern are molybdenite (MoS2) and
pyrite (FeS2). Molybdenite is widely acknowledged as a Re sink, in which Re is concentrated
(which would affect Os in the age corrections). However, while molybdenite has the potential to
affect bulk sediment Re concentrations and, therefore, 187Os/188Os, molybdenite has not been
found in significant quantities BH 9/05. In bulk elemental analyses of the core (not presented in
this study), molybdenum is below detection limits for the entire PETM interval, implying that
total Mo amounts are very low and thus molybdenite is not likely to be present (or makes up only
a very small percentage of the overall content).
51
Pyrite, whose low concentrations in BH 9/05 vary over the PETM interval (Riber, 2009), also
has the potential to influence bulk sediment Re concentrations. In this scenario, pyrite is simply a
means of capturing mobile Re in the sedimentary column. However, concentrations of Re in
pyrite are relatively low, because Re does not fit very well into the pyrite lattice (G. Ravizza,
pers. comm.). Pyrite neither contains enough Re, nor occurs at high enough levels in the
sediment, to affect the 187Os/188Os record. Overall, then, diagenesis does not seem to be a likely
explanation for the Re or Os geochemistry of core BH 9/05.
My conclusion, therefore, is that the 187Os/188Os peak at 533.07 m is an environmental signal,
and not the product of Re mobility during diagenesis.
6.5 The 187Os/188Os Minimum: Evidence of Seawater Intrusion
The record from BH 9/05 implies a likely seawater intrusion to the basin prior to the CIE,
when prominent coincident decreases occur in 187Os/188Os ratios and CIA values. Os isotope
ratios in the basin reflect a balance of Os from radiogenic sources and unradiogenic sources, so
that a substantial decrease in 187Os/188Os values implies either decreased influence from the
radiogenic endmember (river), or increased influence from the unradiogenic endmember
(seawater). There is no evidence that terrestrial Os flux decreased at this time, so a seawater
intrusion seems more likely. Over this time, basin 187Os/188Os ratios decrease from~0.6 to 0.4,
with minimum values nearly identical to synchronous open ocean seawater 187Os/188Os values
(0.36, Ravizza et al. 2001), suggesting a large addition of seawater to the basin. Additionally,
sea level rise has been recorded previously in a variety of sites worldwide during the PETM
(Sluijs et al., 2006; Sluijs et al., 2008; Iakovleva et al., 2001). The cause of the transgression has
52
been attributed to thermal expansion (~5 m rise (Sluijs et al., 2006)), melting of small ice
reserves on land, or displacement due to the emplacement of the North Atlantic Igneous Province
(NAIP) (Sluijs et al., 2008). Based on this evidence, it seems that sea level transgression did
occur over this time interval and could explain the decrease in 187Os/188Os ratios in BH 9/05.
There is also evidence river input remained high at this time as well. A decrease is
observed in CIA values, reflecting either a change to less-weathered source material or a greater
proportion of less-weathered material in the sample. The CIA decrease is concurrent with an
interval of increased grain size (Zr/Rb data from BH 9/05; Riber, 2009), implying that coarsergrained, less weathered particles began to dominate the composition, lowering the CIA value.
Since seawater input is assumed to be high at this time, the source of these grains could be from
increased river output. This suggests that river input continued to increase on top of a seawater
transgression to the basin and could explain the concurrent decreases in 187Os/188Os ratios and
CIA values in BH 9/05.
Additional evidence of sea level rise in the Central Basin is recorded in shifts in dinocyst
assemblages (Harding et al., 2011). A nearby core from Longyearbyen spanning the PETM
shows an initial dominance of terrestrial phytoclasts and low abundance of marine dinocysts in
the basin prior to the CIE. Subsequent increases in low salinity dinocyst assemblages and
abundance of terrestrial fern spores indicate that river input remained high throughout the onset
of the CIE. However, mineralogical clay evidence of an increasingly distal terrestrial source and
the appearance of the dinocyst genus Apectodinium prior to the CIE led Harding et al. (2011) to
conclude that marine transgression and local environmental changes began prior to the onset of
the CIE. Another record from Sluijs et al. (2008) suggests that sea level rise off the New Jersey
53
coast began 20-200 ka prior to the PETM, in which case the record from BH 9/05 may further
constrain timing of the transgression.
Numerical modeling performed in this study constrains the potential amount of water
added using Os concentrations, according to several different scenarios of seawater influence
(Fig. 15, Fig. 16). The first depicts an endmember case in which river input to the basin does not
change prior to the CIE. In this scenario, river [Os] is set to the average river [Os] value of 7.9
pgOs/kgH2O (Levasseur et al., 1999), while ocean [Os] is calculated by the model, and from this
water exchange is derived. This scenario requires a seawater water flux input of ~7x the initial
water flux to the basin. However, when river input steadily increases prior to the CIE (likely a
more realistic scenario, due to increased temperatures and precipitation, with ocean [Os] once
more calculated from the model and river water calculated from a river [Os] of 7.9
pgOs/kgH2O), a greater seawater Os contribution is required to overcome the river Os
contribution, requiring a seawater flux of ~18x the initial seawater flux to the basin. Modeled
maximum salinities for these scenarios are roughly 33 ppt (parts per thousand) and 34 ppt
respectively, nearly identical to modern marine salinity (35 ppt), providing evidence that the
basin would have to have become nearly fully marine at this time in order to produce the
observed 187Os/188Os record.
54
Fig. 15: Changes in oceanic water flux to the basin. a) top: the 187Os/188Os record from BH
9/05. b) middle: Results from the basin-ocean model for changes in the Os exchange flux with
the ocean. c) bottom: Results from the basin-ocean model for changes in the water exchange
flux with the ocean. These data were calculated based on the molar Os input flux from the ocean
to the basin (Fin) divided by the ocean concentration at time (t) and converted to kg H2O.
55
Fig. 16: Results from the basin-ocean model for the most realistic scenario in which river Os
input continued to increase prior to the PETM but was overridden by a seawater flooding event
from 4 to 19.4 ky after the model leaves steady state at 20 ky. a) top panel shows the
hypothesized river input to the basin (the output from the Fr basin-ocean model) with the dashed
line representing the ‘swamping’ of the 187Os/188Os signal after a seawater excursion. b) bottom
panel shows the required increase in seawater input to the basin necessary to override the river
signal during that time to achieve the recorded 187Os/188Os ratio.
56
Though abundant evidence supports a marine transgression prior to the CIE, another
potential explanation for the decrease in 187Os/188Os ratios and CIA values prior to the onset of
the CIE could be increased input from another unradiogenic Os source: volcanism. Increased
contributions from basalt are possible, since large-scale rifting occurred in the North Atlantic
throughout the early Cenozoic. Increased basaltic and/or hydrothermal influence at the onset of
the PETM could be used to suggest a volcanic trigger for the PETM, as hypothesized by Jolley et
al. (2002) and Svensen et al., (2004). The 187Os/188Os record from Zumaya (Schmitz et al., 2004)
also seems to display a general decrease across the onset of the PETM, which could be additional
evidence for volcanic influence, though further analyses are needed to resolve the timing and
magnitude of the excursion. However, there is no evidence that rifting occurred directly within
the Central Basin at this time, and DSDP 549 does not record a decrease in seawater 187Os/188Os
values prior to the 187Os/188Os maximum. Additionally, elemental records from BH 9/05 (not
presented in this study) show no clear peaks in elements typically concentrated in basalts (e.g.
Mg, Fe, Si) or heavy metals coinciding with the 187Os/188Os decrease, making volcanism an
unlikely cause for the 187Os/188Os minimum.
A last explanation for the decrease in 187Os/188Os values could potentially be contribution
from an extraterrestrial source. Kent et al. (2003) suggest that a comet impact may be the main
trigger for the PETM (though re-examination of the Ir data by Schmitz et al. (2004) argues
strongly against this explanation) and so it is possible that the decrease in 187Os/188Os values
prior to the CIE is evidence of that impact. However, several lines of reasoning argue against
this. First, while chondritic impactors typically have unradiogenic Os values (0.12-0.13)
(Paquay et al., 2008) so could potentially cause the observed 187Os/188Os decrease, they also have
high Os concentrations ~500,000 ppt (Peucker-Ehrenbrink, 1996). If such an impact were the
57
case, one would expect a simultaneous surge in Os concentration. However, there is no
appreciable change in Os concentration in BH 9/05 at the time of the 187Os/188Os decrease.
Second, there is a lag of >12 ka between the 187Os/188Os decrease and the decrease in 13C
values. If a comet impact was responsible for the light carbon addition evidenced by the CIE as
proposed by Kent et al. (2003), the two events should be simultaneous. However, they are not.
Lastly, Paquay et al. (2008) examine the effects of several large impacts on the marine
187
Os/188Os record; for each the duration of the 187Os/188Os decrease is much longer than that
observed in BH 9/05. This evidence implies that the source of the 187Os/188Os decrease was not a
comet impactor. Thus, at this time, it appears that the most likely source of the decrease in
187
Os/188Os ratios was rising sea level associated with global marine transgression.
6.6 Proposed Change in Depositional Environment over the PETM
6.6.1. Effects of Terrestrial Runoff and Rising Sea Level
The environment at BH 9/05 appears to be a deep water basin characterized by two
competing trends: increasing river output and sea level transgression during the PETM interval.
These trends could lead to two very different basin environments: a stratified basin with marine
input covered by a lens of freshwater, or a mixed-water brackish environment. These
possibilities are considered in further detail below.
58
Fig. 17: Hypothesized bottom-water environment over the course of the PETM. Reducing or
suboxic conditions are assumed to be correlated with increases in Re concentrations and pyrite
contents.
6.6.2. Evidence for Local Hypoxia and a Freshwater Cap
During the onset and recovery period of the CIE, there is evidence that the depositional
environment of BH 9/05 became hypoxic. Rhenium concentrations increase from ~0.7 to 17 ppb
from 529.05 to 514.75 mbs, coincident with a sequence of laminated shales and lack of
bioturbation (Fig. 17), the lowest 13C values in BH 9/05, elevated levels of pyrite (10-23%), and
foraminiferal and biomarker data which suggest anoxic conditions in the basin (Riber, 2009;
Harding et al., 2011; Cui et al., 2011). Rhenium is sensitive to redox conditions and is greatly
concentrated in reducing conditions (Yamashita et al., 2007; Crusius et al. 1996). Osmium
concentrations also display a slight peak during this interval (~200 pg/g, an increase of 33% from
initial steady state values), which is consistent as Os is only mildly affected by local redox
59
conditions (Yamashita et al., 2007). Over this interval the Th/U ratio also decreases as U
increases (Riber, 2009), further indicating reducing conditions as U is enriched in reducing
sediments (i.e. Calvert and Pedersen, 1993). Therefore the large increase in rhenium and
moderate increase in Os concentration is interpreted as a change in redox state to reducing
conditions of suboxic or anoxic levels. This transition to anoxic or suboxic conditions over the
PETM is consistent with observations of anoxia over the PETM in a variety of sites worldwide
(Thomas and Shackleton, 1996; Sluijs et al., 2006).
Anoxia in the Central Basin over the PETM could have been caused by high productivity
(Cui et al., 2011) and/or by stratification of the water column (Harding et al., 2011). Riber
(2009) places the anoxic interval within the maximum flooding zone during the carbon isotope
recovery, indicating sea levels were high in the basin during this time. River input was also
theorized to be high during the PETM interval, evidenced by co-dominance of low-salinity and
marine dinocyst assemblages (Harding et al., 2011). The mixing of less-saline waters with
surface marine waters could create a less-dense brackish top layer freshwater ‘cap’, limiting the
mixing of deeper water with the water above and thereby allowing for oxygen depletion of
bottom waters. Harding et al. (2011) suggest there may have been periodic inputs of fresh water
during this anoxic interval, consistent with occasional pulses of increased river input, which
would serve to keep the surface water brackish and the water column stratified.
Tracing the Os cycle shows that a stratified basin with a freshwater lens is plausible. In this
scenario, the majority of Os in the deposited sediments would be taken up with organic matter in
the river-derived freshwater cap, thus recording predominantly a river Os signal (more
radiogenic than open seawater). As organic matter settled through the water column, the
60
associated Os would be deposited in the sediments. This could explain why 187Os/188Os ratios
remain consistently more radiogenic than seawater throughout the CIE and recovery.
A second possibility is that during the PETM, the basin instead became a brackish
environment. In this case there would be no freshwater lens, and the basin would be
characterized by a reduction in water mixing as opposed to full stratification. This case is
supported by the BH 9/05 pyrite record, as frequent seawater contact is necessary to renew the
supply of sulfur species in the basin used in the production of pyrite, assuming the pyrite formed
at the time of deposition (cf. Fig. 17), though the low Mo levels in BH 9/05 limit the amount of
free H2S in the water column. In this scenario, Os could be derived from anywhere in the water
column, as radiogenic Os from the surface would be mixed with unradiogenic Os from open
seawater, creating a more radiogenic Os signal in the deposited sediments than pure seawater. In
this scenario, anoxia would derive from over-productivity rather than from lack of contact with
oxygenated surface waters. However, this scenario doesn’t adequately explain the large presence
of low-salinity dinocysts in the basin throughout the PETM. One possibility is that the lowsalinity dinocysts were typically located near the river mouths and were washed out into the
basin with increased river water flux during the onset of the CIE. However, that wouldn’t
explain their continued high abundances throughout the main body of the CIE and early recovery
(Harding et al., 2011). Therefore, I conclude that this case is less likely to have characterized the
Central Basin over the PETM.
Comparisons between BH 9/05 and modern restricted basins show that a restricted basin
environment is plausible but emphasize that the Os record of the basin is greatly determined by
the amount of exchange with the open ocean. For instance, in the restricted euxinic Black Sea
(considered one of the best modern analogues to past anoxic environments), a significant portion
61
of the total Os is associated with—and thought to be removed by—organic matter (Ravizza et
al., 1991), similar to the first basin environment proposed above.
187
Os/188Os ratios of sediments
in the Black Sea range from 0.68-0.85, while Os and Re concentrations are 230-690 ppt and 2185 ppb respectively (Ravizza et al., 1991). These concentrations are much higher than those
observed in BH 9/05, providing strong evidence for greater concentration of both Os and Re in
more restrictive and reducing environments. Water flow to the Black Sea is ~7E11 kg/yr, two
orders of magnitude smaller than the amount predicted for the BH 9/05 basin; accordingly the Os
burial flux is nearly 50 times greater in the Black Sea (Ravizza et al., 1991), supporting a
consistency between the numerical modeling of this study and the relationship between degree of
restriction and Os uptake to sediments.
In contrast, the Black Sea’s 187Os/188Os ratios appear to show a trend opposite to that
observed at BH 9/05. While similar in value to those of BH 9/05, the Black Sea’s 187Os/188Os
ratios are much lower than those of the modern ocean (~1.06; Levasseur et al., 1998), unlike BH
9/05’s 187Os/188Os values which are hypothesized to be more radiogenic than contemporaneous
seawater. Paradoxically, while the effects on the Os record are opposite those of BH 9/05, the
cause is likely the same—increased influence from terrestrial Os. In BH 9/05, the underlying
lithology is largely granite and shale (Harland 1969), which would contribute radiogenic Os to
the system, thus making the basin more radiogenic than seawater. In contrast, the lithogenic
units surrounding the Black Sea are likely unradiogenic (Ravizza et al., 1991), making the basin
less radiogenic than modern seawater, with occasional seawater intrusions having large effects
on the Os record (Ravizza et al., 1991), similar to the scenario hypothesized for BH 9/05. The
Black Sea thus provides a critical example of a terrestrially-dominated restricted basin and
62
illustrates the importance of the surrounding environment and the degree of restriction on the Os
and Re records.
Os evidence from the Cariaco Basin in Venezuela provides another modern example for
comparison. The basin is semi-restricted today, with an anoxic zone at depth and oceanic flow
exchange occurring through a series of surficial sills. Unlike BH 9/05 or the Black Sea,
187
Os/188Os ratios in the Cariaco Basin are nearly identical to synchronous seawater (0.9-1.05;
Oxburgh et al., 2007), suggesting that water flow is less restricted, similar to the second basin
environment. This lack of restriction is supported by oxygen isotopic data, showing a marked
similarity to global 18O records through time (Oxburgh et al., 2007). In this less-restricted
environment, basin 187Os/188Os appear to be influenced more by the balance between terrestrial
and marine 187Os/188Os values and less by redox concentration effects on geochemistry.
Determining a basin’s restriction through time is challenging yet can have important
implications for geochemical records. The degree of restriction in reducing basins can be
estimated by plotting a regression line of Mo to total organic carbon (TOC) (McArthur et al.,
2008), with lower numbers indicating more severe restriction. Molybdenum data have not been
obtained beyond low-resolution general levels for BH 9/05, and future studies are needed to
investigate this, along with further examination of partitioning constraints on Mo, Re, and Os in
reducing sediments, in order to further characterize the environment of the BH 9/05 basin.
63
6.6.3. Summary of Depositional Environment at BH 9/05
The depositional environment at BH 9/05 appears to have been a deepwater basin setting,
(Riber, 2009; Dypvik, 2011) near enough to land to show terrestrial influences (Cui et al., 2011).
Elevated levels of terrestrial weathering are hypothesized to have occurred prior to the onset of
the CIE at the PETM, supported by physical and modeling evidence.
Over the PETM interval, the water column likely became stratified, supporting a freshwater
lens as the top layer, leading to anoxia in the bottom waters. This shift to reducing conditions
caused changes in partitioning of several key redox-sensitive elements in the sediment, including
Re, Os, and U (Riber, 2009). Consideration of the basin as a brackish environment does not
appear to be consistent with dinocyst data from Harding et al. (2011). Lastly, Harding et al.
(2011) propose periodic pulses of fresh water input to the basin over the course of the PETM
anoxic interval. This could also potentially explain some of the variance in the 187Os/188Os
record during the recovery of the CIE.
64
Chapter 7
Conclusions
Core BH 9/05 provides one of the highest resolutions of the PETM interval to date. Over the
PETM interval, 187Os/188Os and CIA values show a sharp decrease prior to and during the onset
of the CIE, followed by an increase to maximum values before gradually returning to steady state
values during the CIE recovery. Pre-PETM steady state and maximum 187Os/188Os values are
~0.6, nearly twice those of the open ocean, suggesting that BH 9/05 most likely records a basin
signal influenced by terrestrial sources. Increased 187Os/188Os ratios, high CIA values, and
previously published kaolinite values (Riber, 2009) all peak in the interval from ~530-533 m
(during the CIE onset), providing three independent lines for increased terrestrial weathering.
Modeling results support this idea, pointing to an increased river molar Os flux due to increased
weathering rather than a change in source lithology. A decrease in 187Os/188Os ratios prior to the
PETM is hypothesized to be a large seawater intrusion into the basin as part of a global marine
transgression, superimposed upon a general trend of increased terrestrial runoff and increasing
weathering. Os and Re concentrations peak during the CIE recovery, during an interval of
laminated shales, low Th/U ratios, and high pyrite concretions, providing evidence for the
occurrence of regional euxinia (Riber, 2009; Cui et al., 2011). These changes imply that
increased weathering did occur during the hyperthermal conditions of the PETM and reinforce a
link between weathering responses and rapidly changing climate.
65
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APPENDIX A: Osmium and Rhenium Geochemical Data
Table 2: Osmium and Rhenium Analyses from BH 9/05. (*) indicates age-corrected values. 2
values for Re are < 4% of the recorded value.
( ^ ) = One rhenium outlier was discarded; the Re used to age correct that sample was instead an
average of other Re measurements at that depth.
measured
187
Os/188Os
0.6885
0.7645
0.7898
0.8098
0.708
0.527
0.4339
0.4942
0.5626
0.6111
0.5791
0.8478
0.8162
Re
(ppb)
2.93
3.30
3.26
2.49
2.12
3.12
3.12
2.58
3.45
4.53
5.09
0.68
1.03
187Re/
188Os
116.35
104.51
109.25
74.963
83.144
86.701
84.022
88.764
109.84
142.28
146.34
25.754
31.395
correction
0.108
0.097
0.102
0.070
0.077
0.081
0.078
0.083
0.102
0.132
0.136
0.024
0.029
agecorrected
187
Os/188Os*
0.58
0.67
0.69
0.74
0.63
0.45
0.36
0.41
0.46
0.48
0.44
0.82
0.79
Age
(MA)
55.8443
55.8427
55.8427
55.8391
55.8357
55.8327
55.8327
55.8310
55.8300
55.8288
55.8271
55.8248
55.8232
Depth
544.09
543.16
543.16
541.70
539.05
537.36
537.36
536.36
535.76
535.05
534.05
533.07
532.38
Type
Carius
Bulk
Carius
Carius
Carius
Bulk
Bulk
Carius
Carius
Carius
Carius
Bulk
Bulk
Os
(ppt)
130.12
164.59
156.37
174.11
132.34
182.16
185.8
146.62
160.04
163.06
177.37
138.17
171.55
532.38
531.68
531.35
530.99
529.05
527.45
527.05
526.45
525.90
520.75
520.75
516.60
Carius
Bulk
Bulk
Carius
Carius
Bulk
Bulk
Bulk
Bulk
Bulk
Carius
Bulk
165.76
147
170.82
165.52
171.56
166.46
172.15
172
185.17
200.68
181.43
206.77
0.7944
0.839
0.8372
0.7754
0.8728
0.9173
0.8609
0.938
0.9096
0.9876
1.0058
0.8128
0.96
1.24
2.21
1.45
5.57
8.89
8.34
10.51
9.31
13.13
11.63
11.04
30.369
44.315
68.18
45.821
171.57
283.9
255.83
325.69
266.98
350.8
344.33
280.23
0.028
0.041
0.063
0.043
0.160
0.264
0.238
0.303
0.248
0.326
0.320
0.260
0.77
0.80
0.77
0.73
0.71
0.65
0.62
0.64
0.66
0.66
0.69
0.55
55.8232
55.8214
55.8206
55.8197
55.8147
55.8107
55.8097
55.8082
55.8068
55.7938
55.7938
55.7833
516.60
514.75
512.58
510.05
507.75
507.75
506.75
Carius
Carius
Carius
Carius
Carius
Bulk
Bulk
198.88
179.88
218.2
184.41
213.81
161.85
122
0.8346
1.0795
0.7916
0.6892
0.7473
0.7001
0.752
10.80
17.30
1.88
1.85
4.19
1.70
1.23
285.86
520.99
45.069
51.859
102.15
54.388
52.372
0.266
0.484
0.042
0.048
0.095
0.051
0.049
0.57
0.60
0.75
0.64
0.65
0.65
0.70
55.7833
55.7779
55.7772
55.7580
55.7482
55.7482
55.7440
73
504.40
502.05
502.05^
502.05
498.45
498.45
494.97
491.03
490.05
483.05
483.05
Carius
Bulk
Carius
Carius
Carius
Bulk
Bulk
Bulk
Carius
Bulk
Carius
135.93
123.79
118.73
147.11
119.37
131.64
114.31
116
110.57
113
116.4
0.7517
0.6311
0.6533
0.6133
0.6852
0.6466
0.626
0.66
0.6628
0.692
0.7048
1.25
2.10
2.10
2.19
1.63
1.51
0.83
1.29
1.54
3.04
2.02
47.859
87.137
91.098
76.311
70.677
58.891
37.149
57.241
71.717
139.19
89.951
74
0.044
0.081
0.085
0.071
0.066
0.055
0.034
0.053
0.067
0.129
0.083
0.71
0.55
0.57
0.54
0.62
0.59
0.59
0.61
0.60
0.56
0.62
55.7340
55.7215
55.7215
55.7215
55.7052
55.7052
55.6901
55.6744
55.6705
55.6243
55.6243
Table 3: Comparison of measured Os concentration (top left) and 187Os/188Os ratios (top right)
from Carius tubes and from bulk samples. 187Os/188Os ratios are direct measurements (pre-age
correction) so that differences in Re values have no influence on the result. Re measurements
are included separately (bottom). Note that the samples with the largest difference in 187Os/188Os
also have large differences between measured Os and Re concentrations, suggesting that sample
splits are not compositionally homogenous.
187
Os/188Os
Os (ppt)
Sample
Depth
543.16
532.38
520.75
516.60
507.75
502.05
498.45
483.05
Carius
156.37
165.76
181.43
198.88
213.81
147.11
119.37
116.4
Bulk
164.59
171.55
200.68
206.77
161.85
123.79
131.64
113
Sample
Depth
543.16
532.38
520.75
516.60
507.75
502.05
498.45
483.05
% diff
-5
3
-10
-4
32
19
-9
3
Carius
0.7898
0.7944
1.0058
0.8346
0.7473
0.6133
0.6852
0.7048
Re (ppb)
Sample
Depth
543.16
Carius
3.26
Bulk
3.30
% diff
-1
532.38
0.96
1.03
7
520.75
11.63
13.13
-11
516.60
10.80
11.04
-2
507.75
4.19
1.70
147
498.45
1.63
1.51
8
483.05
2.02
3.04
-34
75
Bulk
0.7645
0.8162
0.9876
0.8128
0.7001
0.6311
0.6466
0.692
% diff
3
3
2
3
7
-3
6
2
APPENDIX B: Model Setup and Initial Parameters
1. GLOBAL MODEL (1-Box)
Two global models were set up: one model for a case solving for changes in the global riverine
Os flux (Fr Model) based on the observed oceanic Os record, and another model solving for
changes in the global riverine isotope ratio (Rr Model). The models were run for two cases: one
using BH 9/05 Os data as the global ocean Os signal, and the second using DSDP 549 as the
global Os signal.
BH 9/05 data was prepared in the following way: data was first age corrected using recorded Re
values, then assigned ages based on Cui et al. (2011) (using the age model from Charles et al.
(2011)). Samples of combined depths were assigned the average age. Since STELLA only
allows single data entries, multiple Os values at a single depth were averaged.
1.1 Parameters and Assumptions
The constants and equations used in the model constructions are as follows. As a general rule,
fluxes are notated as Fx and isotope ratios as Rx, where ‘x’ stands for the first letter of the input
or output source or reservoir in question. For example, Fr means the molar Os flux from rivers.
Oceanic Os Equation:
(
)
(∑
∑
)
(
)
(Equation 4), where
76
(∑
∑
)
(
)
(∑
∑
)
∑
Below is a list of the initial parameters for the global ocean model.
77
Table 4: Initial parameters for the global model.
Parameter
Fluxes
Fh
FET
Fr
Value
Units
100 kgOs/yr
36 kgOs/yr
calculated from Fr Model
89.336623 kgOs/yr
37.342307 kgOs/yr
Deposition Flux
(
Os Isotope Ratios (187Os/188Os)
Rh
RET
R
) (
Reason
estimated modern-day valuea
estimated modern-day valueb
using initial Ro from BH 9/05 data
using initial Ro from DSDP 549
data
()
(
(
)
)
)
estimated modern-day valuec
estimated modern-day valuec
modern average river 187Os/188Osd
0.12
0.1
1.4
Other Parameters
Ocean Os Conc, init
10.86 pgOs/kgH2O estimated modern-day valuee
Mols Os, init
calculated: (Ocean Vol* Init Ocean [Os])/(Os At.Wt.)
modern volume plus addition of
amt of water in ice today (assumed
to be no ice at PETM)f
Timestep
not necessary, used only to make conversions in one step
Sum of the stable Os isotopes
normalized to 188Os (not including
187
Sum Os ratios
7.41092
Os/188Os)
a
McDaniel et al. (2004) based on estimate of proportion of river input
b
Peucker-Ehrenbrink (1996)
c
Peucker-Ehrenbrink and Ravizza (2000)
d
Ravizza and Zachos (2003)
e
Levasseur et al. (2000)
f
Trenberth et al. (2007)
Ocean Volume
1361390000 km^3
78
1.2 Schematics
1.2.1 Fr Model Setup
Fig. 18: schematic of the global model setup for the Fr case (driving the observed marine
187
Os/188Os ratio by changes in the riverine Os flux).
79
1.2.2 Rr Model Setup
Fig. 19: schematic of the global model setup for the Rr case (driving the observed marine
187
Os/188Os ratio by changes in the average global river 187Os/188Os ratio).
80
2. BASIN-OCEAN COMBINED MODEL (2-Box)
2.1. Matching C-data from BH 9/05 and DSDP 549—Case 1 and Case 2
In the following models, BH 9/05 Os data is assumed to be the PETM record of the
Svalbard Central Basin, while DSDP 549 Os data is assumed to be the PETM global ocean
record. This requires that the two datasets be temporally matched. Matching was done using the
carbon isotope records from each core, as the carbon isotope data for each core is higherresolution than the Os data.
Carbon data from BH 9/05 was used from Cui et al. (2011). Carbon data from DSDP 549
was combined from studies Nunes and Norris (2006) and Stott et al. (1996). These two studies
use different species of foraminifera, so for the current study carbon isotope data was used from
Stott et al. (1996), for two reasons: because the carbon data was more extensive, and because it
had been previously matched with the Os record in DSDP 549 (Ravizza et al., 2001). The
carbon isotope record from Nunes and Norris (2006) was used to correlate age of samples and
depth at Site 549. Ages in Nunes and Norris (2006) are based on the age-model of Norris and
Röhl (1999), using tie-points between the carbon isotope records from DSDP 549 and ODP 690.
Where possible, depths and carbon data from Stott et al. (1996) were matched with a
direct age from Nunes and Norris (2006). Linear sedimentation rates were calculated between
each two reported depths and assumed to be constant in the interim and near the ends. If direct
correlation between depth and age was not possible (for depths that were not reported in Nunes
and Norris (2006)), the calculated sedimentation rate was used to calculate approximate ages. In
such a way over the PETM interval each of the carbon isotope data points reported in Stott et al.
(1996) was assigned an age.
To correlate the carbon isotope records from DSDP 549 and BH 9/05, the onsets of the
two carbon excursions were matched, under the assumption that the carbon isotope excursion
was synchronous worldwide. In core BH 9/05, the onset of the PETM begins at 536 m (Cui et
al., 2011), and ends at ~487 m (excursion length of 170 ky). For DSDP 549, depth 339.97m was
picked as the onset of the C-excursion, as δ13C decreases coincident with the increase in
187
Os/188Os (Ravizza et al., 2001).
Two possible endpoints exist for the DSDP 549 excursion. At depth 336.68 m, both the
Os record and the C-record from Stott et al. (1996) appear to have reached a new steady-state
value, making 336.68 m one possible endpoint of the excursion. The case could also be made for
depth 337.47 m as the end of the excursion, as it represents a major inflection point and possibly
the return to new steady-state conditions, as subsequent 187Os/188Os values change less than 2%.
Thus for each of the following models two cases were run: one with an excursion endpoint at
337.47 m (Case 1) and another with an endpoint at 336.68 m (Case 2). In both cases the
excursion lengths were set to 170 ky.
81
Fig. 20: top: Comparison of the carbon isotope records of BH 9/05 and DSDP 549 assuming
Case 1 and 2 matching. Carbon isotope data from Ravizza et al. (2001) is measured from bulk
carbonate, while carbon isotope data from BH 9/05 is total organic carbon (Cui et al., 2011).
bottom: Comparison of the corresponding 187Os/188Os records of BH 9/05 and DSDP 549 for
Case 1 and 2 data matching.
82
2.2 Models
Several variations were run for the combined basin-ocean model: Fr models (assuming
that the 187Os/188Os change was caused by a change primarily in riverine Os flux), Rr models
(assuming that the 187Os/188Os change was caused by a change primarily in the local or global
average riverine 187Os/188Os ratio), and Fin models (assuming as an endmember case that the
change in the basin 187Os/188Os change was caused by a change primarily in the basin-ocean
water exchange flux). Each model was run twice, using first Case 1 and then Case 2 as the
global ocean 187Os/188Os record. Below is a table of the different models used in this study
(Table 5).
Table 5: Different model simulations used in this study.
Model
Fr Model
Rr Model
Fin Model
Combined Models:
Fin model
Number
1a
2a
2b
3
Changed (dependent calculated) parameter
basin Fr, global Fr (basin Fr change in river [Os])
basin Rr, global Fr
basin Rr, global Rr
basin Fin, global Fr
4a basin Fin, global Fr, with basin increase in Fr
basin Fin with basin increase in Fr, global Fr (basin Fr
increase in water discharge, basin volume stays in
4b steady state)
83
2.3 Schematics
2.3.1 Fr Model Setup
Fig. 21: Schematic of the basic Fr model (change in river Os flux) for the combined basin-ocean
system. For these models, BH 9/05 is assumed to be the Os record for the basin, while DSDP
549 is assumed to be the global ocean Os record.
84
2.3.2 Rr Model Setup
Fig. 22: Schematic of the basic Rr model (change in local and global river 187Os/188Os ratio) for
the combined basin-ocean system.
85
2.3.3 Fin Model Setup
Fig. 23: Schematic of the basic Fin model (change in the exchange flux between the basin and
ocean). Note that while in this case the basin signal is assumed to be caused by changes in the
exchange flux (Fin and Fout), the ocean signal remains a product of changes in the global river
Os flux (global Fr).
2.4 Combined Fin-Fr Model:
This model tests the potential influence of a seawater ‘swamping’ of the riverine Os
signal between 539 m and 533.01 m, during which the 187Os/188Os values decrease to near
marine values. This model assumes an increasing local river molar Os flux is the dominant
factor driving the basin Os signal except between 539 m and 533.01 m, where the model solves
for the amount of seawater input (Fin) required to overprint the river Os signal. During this
interval, the conservative estimate of river Os flux increase as a straight line between the points
is used as the river signal, while Fin is calculated.
86
2.4.1 Parameters and Assumptions
The ocean box of this model is set up using the following equations, assuming input
fluxes from rivers, hydrothermal mixing, extraterrestrial matter, and a basin exchange flux
(inbound flux Fout from basin and outbound flux Fin to basin), and an additional outbound
deposition flux. Note that water leaving the ocean will have the same 187Os/188Os as the ocean
(Ro), and water leaving the basin will have the basin’s 187Os/188Os (Rb).
The equations used are as follows:
Oceanic Os Equation:
(
)
(
(∑
∑
(∑
)
∑
)
(
)
(∑
∑
)
)
(Equation 4), where
∑
87
(
)
(∑
∑
)
And similarly for the basin:
Basin Os Equation:
(
)
(∑
∑
)
(
)
Where
88
(∑
∑
)
Table 6: Initial parameters for the Basin-Ocean combined model.
Ocean:
Parameter
Fluxes
Fh
FET
Fr
Value
Units
100 kgOs/yr
36 kgOs/yr
calculated from Fr Model
195.570001270 kgOs/yr
Reason
estimated modern-day valuea
estimated modern-day valueb
using initial Ro from DSDP 549
data
37.342307 kgOs/yr
Fin(Os) to basin
(flux out of ocean to
basin)
Deposition Flux
calculated: Fin (water)*Ocean_Conc/(1E12pg/g*Os At. Wt.)
(
Os Isotope Ratios (187Os/188Os)
Rh
RET
Rr
) (
()
(
(
)
)
)
estimated modern-day valuec
estimated modern-day valuec
modern average river 187Os/188Osd
0.12
0.12
1.4
Other Parameters
Ocean Os Conc, init
Mols Os, init
Ocean Volume
Timestep
Sum Os ratios
10.86 pgOs/kgH2O
estimated modern-day valuee
calculated: (Ocean Vol* Init Ocean [Os])/(Os At.Wt.)
modern volume plus addition of
amt of water in ice today
(assumed to be no ice at PETM)f
not necessary, used only to make conversions in one step
Sum of the stable Os isotopes
normalized to 188Os (not including
187
7.41092
Os/188Os)
1361390000 km^3
Water Cycle
Basin Volume
Arctic River Water
Discharge
Fin (water) to basin
Fout (water)
estimated from paleogeographical
continental locationsg
average of several small Arctic
19.92 km^3/yr
rivers todayh
calculated from init values of Fin model, based on Rr1 of 1.4 and
Arctic River Conc of 7.9 pgOs/kgH2O
set to equal Arctic River water + Fin (water) to keep Basin Volume
constant
~8E14 kg H2O
89
Basin:
Fluxes
Fr Basin (Arctic
River into basin)
Fout of basin (Os
flux)
Deposition Flux
calculated based on Arctic River Discharge and River Os Conc
calculated: (Basin_Conc/(1E12*190.2))*Fout_water_from_basin
(
Os Isotope Ratios (187Os/188Os)
Rr1
) (
()
(
(
)
)
)
modern average river 187Os/188Osd
1.4
Other Parameters
calculated assuming partitioning same in basin as ocean, using
averaged DSDP 549 data
9.20338983050847
pgOs/kg
Basin River Os Conc
7.9 pgOs/kgH2O
modern average river [Os]i
a
McDaniel et al. (2004) based on estimate of proportion of river input
b
Peucker-Ehrenbrink (1996)
c
Peucker-Ehrenbrink and Ravizza (2000)
d
Ravizza and Zachos (2003)
e
Levasseur et al. (2000)
f
Trenberth et al. (2007)
g
Talwani and Eldholm (1977)
h
Milliman and Syvitski (1992)
i
Levasseur et al. (1999)
Basin Os Conc, init
90
APPENDIX C: Model Output
1. GLOBAL MODEL
1.1 Fr Model (Endmember case that all observed change is caused by changes in the Os
river flux)
Global River Os Mass Flux
BH 9/05 DSDP 549
Initial value
89.3
37.3
Min
44.9
37.3
Max
193.2
52.7
94
39.9
Ending value
Fig. 24: Results from the global ocean model for changes in the river Os flux. This model was
made to see if both core records could be caused by nearly the same input fluxes. However,
since the required Os kg/yr input for BH 9/05 is more than twice that of DSDP 549, the
conclusion is that they cannot both be recording the same signal (since Os is well-mixed in the
ocean (Ravizza and Zachos, 2003), the two signals would not be influenced by different sources,
the other possibility), and so they cannot both be an ocean Os signal.
91
1.2 Rr Model
Global River 187Os/188Os
BH 9/05
DSDP 549
Initial value
1.4
1.4
Min
0.87
1.4
Max
2.18
1.87
Ending value
1.45
1.48
Fig. 25: Results from the global ocean model for changes in the river 187Os/188Os ratio. Changes
in Rr causing BH 9/05 would require a global isotope ratio higher than 2. While this is not
impossible, it would require a higher crustal abundance of radiogenic material exposed at the
surface. The question arises whether such large-scale lithology exposure changes can occur over
such a short geological time. Thus the conclusion of this study is that the observed 187Os/188Os
records were most likely not caused by a change in the global Rr alone.
92
2.
BASIN-OCEAN COMBINED MODEL
2.1 Fr Model (assuming both rivers (local Basin River and global rivers) are the entire
cause of the changes in the Os records, solving for Basin River Fr and Global Fr)
93
Basin River Os Mass Flux
Case 1
Initial value
Global River Os Mass Flux
Case 2
0.157
0.157
Min
0.0225
Max
Ending value
Case 1
Case 2
Initial value
37.2
37.2
0.0225
Min
37.0
37.0
0.472
0.473
Max
56.3
54.7
0.161
0.161
Ending value
39.7
39.7
Basin River Os Concentration Change
Case 1
River Conc initial
value
River Conc Min
River Conc Max
7.9
1.13
23.7
Case 2
7.9
1.13
23.8
Fig. 26: Results from the basin-ocean model for changes in the river Os flux. Note that in this
scenario, the Arctic river Os shows a much greater proportional increase than the global river
average, fitting with both a greater temperature change in response to stimuli (i.e. Sluijs et al.,
2006) and increased precipitation at the poles (Pagani et al., 2006), both of which would serve to
increase the total amount of Os reaching the basin from the river.
94
2.2 a) Rr Model (assuming the case where the basin Os record is formed entirely from a
change in the local basin river 187Os/188Os ratio. The global ocean remains driven by
global river Os.))
Basin River Rr
Case 1
Initial value
Global River Os Mass Flux
Case 2
Case 1
1.4
1.4
Min
0.53
Max
Ending value
Case 2
Initial value
37.2
37.2
0.53
Min
37.1
37.1
2.83
2.84
Max
56.4
54.7
1.42
1.42
Ending value
39.7
39.7
Fig. 27: Results from the basin-ocean model for changes in the basin river 187Os/188Os ratio.
Note that this scenario would require a maximum basin river Rr value near 3. While this cannot
be completely ruled out, it would require changes to much more radiogenic lithology at the
surface and back within the span of the PETM, and the timing of the increase and maximum
187
Os/188Os ratio aligning with the increase and maximum kaolinite (another indicator of
weathering) to be coincidental. Therefore, while changes in Rr influencing the basin Os record
cannot be ruled out entirely, a sole Rr cause is unlikely.
95
2.2 b) Rr Model (assuming the case where both the global ocean and the basin are
driven by a change in river Rr)
96
Basin River Rr
Case 1
Initial value
Global River Rr
Case 2
Case 1
1.4
1.4
Min
0.53
Max
Ending value
Case 2
Initial value
1.4
1.4
0.53
Min
1.4
1.4
2.83
2.84
Max
1.99
1.94
1.41
1.41
Ending value
1.48
1.48
Fig. 28: Results from the basin-ocean model for changes in both the Arctic River and the global
average 187Os/188Os ratio. To achieve a global river average 187Os/188Os ratio near 2, the exposed
lithology would have to be much more radiogenic than today. While I cannot eliminate changes
in Rr as a possible influence on the Os record, it is questionable whether such large-scale
changes could have occurred over such a short geological time, and therefore unlikely that Rr
changes were responsible for the entire change in the Os record.
97
2.3 Fin Model (assuming the case where the basin Os record is formed entirely from a
change in the Os exchange flux with the ocean. The global ocean remains driven by
global river Os.)
Water Exchange Flux Proportions
Case 1
Case 2
1
1
0.34
0.34
7.0
7.0
0.98
0.97
Initial value
Min
Max
Ending value
Fig. 29: Results from the basin-ocean model for changes in the Os water exchange flux with the
ocean. These data were calculated based on the molar Os input flux from the ocean to the basin
(Fin) divided by the ocean concentration at time (t) and converted to kg H2O.
98
2.4 Fin Model (assuming the case where the basin 187Os/188Os increase is a product of
increased river flux while the pre-PETM decrease is due to a sudden flooding
(increased Os exchange flux with the ocean). The global ocean remains driven by
global river Os.)
99
Water Exchange Flux Increase (proportional)
Case 1
Case 2
1
1
Initial value
Max, 1st peak
18
18
Max, 2nd peak
9.3
9.3
1
1
Ending value
Fig. 30: Results from the basin-ocean model for the most realistic scenario in which river Os
input continued to increase prior to the PETM but was overridden by a seawater flooding event
from 4 to 19.4 KY after the model leaves steady state at 20 KY. The top figure shows the
hypothesized river input to the basin (the output from the Fr basin-ocean model) with the dashed
line representing the ‘swamping’ of the 187Os/188Os signal after a seawater excursion. The
middle figure shows the required increase in water input to the basin from the ocean necessary to
override the river signal during that time to achieve the recorded 187Os/188Os ratio.
100
Table 7: Summary of salinity results for the different models
Case 2
data
Case 1
data
Model
Increase basin Fr,
increase global Fr,
no change in
water
Increase basin Rr
Changes in basin
driven solely by
changes in ocean
exchange flux
Ocean pulse on
top of increasing
Fr in basin (Fr
increase is mol
change in Os)
Ocean pulse on
top of increasing
Fr in basin (Fr
increase is due to
water increase)
Basin
Max
Salinity
Basin
Min
Salinity
Model
Basin
Max
Salinity
Basin
Min
Salinity
24.8
24.8
24.8
24.8
Increase basin
Fr, increase
global Fr, no
change in water
24.8
24.8
Increase basin Rr
24.8
24.8
15.8
Changes in basin
driven solely by
changes in ocean
exchange flux
33.1
15.7
24.7
Ocean pulse on
top of increasing
Fr in basin (Fr
increase is mol
change in Os)
34.2
24.7
15.8
Ocean pulse on
top of increasing
Fr in basin (Fr
increase is due
to water
increase)
33.1
15.7
33.1
34.2
33.1
101