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Paleoceanography of the South Atlantic Ocean from the Middle Eocene through the Oligocene

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Paleoceanography of the South Atlantic Ocean from the Middle Eocene through the Oligocene
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Mead, Gregory A
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xiv, 315 leaves : ill. ; 29 cm.

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Carbonates ( jstor )
Deep water ( jstor )
Drilling ( jstor )
Geomagnetic polarity time scale ( jstor )
Glaciation ( jstor )
Oceans ( jstor )
Oxygen ( jstor )
Salinity ( jstor )
Sea water ( jstor )
Seas ( jstor )
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bibliography ( marcgt )
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Thesis:
Thesis (Ph. D.)--University of Florida, 1995.
Bibliography:
Includes bibliographical references (leaves 294-314).
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Typescript.
General Note:
Vita.
Statement of Responsibility:
by Gregory A. Mead.

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PALEOCEANOGRAPHY OF THE SOUTH ATLANTIC OCEAN
FROM THE MIDDLE EOCENE THROUGH THE OLIGOCENE










By

GREGORY A. MEAD


A DISSERTATION PRESENTED TO THE GRADUATE SCHOOL
OF THE UNIVERSITY OF FLORIDA IN PARTIAL FULFILLMENT
OF THE REQUIREMENTS FOR THE DEGREE OF
DOCTOR OF PHILOSOPHY

UNIVERSITY OF FLORIDA


1995




PALEOCEANOGRAPHY OF THE SOUTH ATLANTIC OCEAN
FROM THE MIDDLE EOCENE THROUGH THE OLIGOCENE
By
GREGORY A. MEAD
A DISSERTATION PRESENTED TO THE GRADUATE SCHOOL
OF THE UNIVERSITY OF FLORIDA IN PARTIAL FULFILLMENT
OF THE REQUIREMENTS FOR THE DEGREE OF
DOCTOR OF PHILOSOPHY
UNIVERSITY OF FLORIDA
1995


This Dissertation is dedicated to the memories of my
Grandmother Minnie Morris and my Mother Shirley Lopez,
Father Daniel Weiss, and Step-father Bruce Mead.


ACKNOWLEDGMENTS
I would like to thank first my advisor and friend, Dr.
David A. Hodell, and the rest of my committee, Drs. Paul F.
Ciesielski, Douglas S. Jones, Neil D. Opdyke, and Claire L.
Schelske, for (at various times) advice, encouragement, and
financial support. Drs. G. W. Brass, J. P. Kennett, D. V.
Kent, K. G. Miller, J. S. Oslick, E. Thomas, J. D. Wright,
and J. C. Zachos contributed valuable comments on the
chapters. The idea for the Braarudosphaera study was
initiated in a class led by Dr. E. L. (Jerry) Winterer. Dr.
Kennett also supplied the washed samples from ODP Hole 689B,
and Dr. Zachos allowed me to use unpublished data from ODP
Hole 748B. It is impossible to thank Ray Thomas enough for
keeping the equipment running. Rich Cooke, Jason Curtis,
Dr. Ann Heatherington, Vicky Mejia, and Dr. Paul Mueller
also helped materially in the laboratory. Richard Crockett
was the Scanning Electron Microscope technician. Hallie
Smith produced the drafted figures.
My fellow Micropaleo lab students, Alex Amigo, Jennifer
Callahan, Jason C., Jose Garrido, Rick Greenfield, Anon
Vayavananda, and Kathy Venz, were invaluable sources of
advice, commentary, brain-storming, and entertainment. I
thank them, along with all those who worked in the lab,
Soraya Balmes, Gail Beals, Belinda Seidel Breidenbach, Gil
Butler, Rich C., Charlotte Davison, Sabine Eppler, Paul
Gremillion, Binhe Gu, Alexandra Isern, Vicky M., Danny
Mller, Mark Mulligan, Tamara Mullis, Sarah Palmer, Mark


Schult, Laura Stanley, Loni Terry, Jang Vayavananda, Ana
Vrba, and Doug Wilder, for their friendship and for making
the lab a pleasant (if at times distracting) place to work
and study.
In addition, I especially thank Gail B., Kim D'Arcy and
Paul Meyers, Vic Divenere (also for teaching me recorder),
Ken Gilland, Lori Guthrie, Jack Holt, George Houston, Paul
(the Mac Master) Kirk, Nancy Mullins (also for the loan of
her printer), Mark S., Ana V., and Mick and Julie Whitelaw
for their companionship, understanding, and friendship over
the years I've been here.
I want also to thank all my family, and my extended
families, the Abrams and the Beals, for their support and
love during this (at times) seemingly endless endeavour.
I also thank the Gainesville Cycling Club, the
Gainesville Area Fiddle Jam and Fast Molasses, and the J. J.
Finley Ultimate players for helping keep me healthy (more or
less) and sane (more or less).
Finally, I thank my cat Trash, and my other 4-footed
friends, Harold, Rhett, Yoda, and Barclay, for their
unquestioned and unending affection.
This research was supported by National Science
Foundation grants OCE-8858012 and DPP-8717854 and by the
Donors of the Petroleum Research Fund administered by the
American Chemical Society.
iv


PREFACE
The late Paleogene (middle Eocene through Oligocene)
was an important time in the Earth's biogeographic,
paleoclimatic, paleogeographic, and paleoceanographic
history. It marked the transition between the warm
Cretaceous and early Paleogene, and the cooler Neogene
[Miller et al., 1991a,b]. The continents of the Earth
were approaching the present positions, as Australia and
India moved northward, the Atlantic Ocean continued to
widen, and the Tethys progressively constricted [Dercourt
et al., 1986]. It was during this time that major seaways
opened around Antarctica [Kennett et al., 1972], allowing
the Antarctic to cool to the point where glacial ice could
eventually build up.
A band of siliceous deposition widened as a result,
and the planetary temperature gradient increased, causing
more restricted biogeographic zonation. Deep thermohaline
circulation in the oceans also developed at this time.
In the Oligocene, a series of deposits in the
Atlantic and Indian Oceans (Braarudosphaera oozes) suggest
that unusual conditions repeatedly occurred, for reasons
that are poorly understood. Biostratigraphic markers are
relatively rare in the Oligocene, leading to poor
resolution in stratigraphy.
Isotopic studies of geochemical tracers have been
important parts of research into this time period in the


oceans. Progressive oxygen isotopic enrichment of
planktic and benthic microfossils has documented oceanic
cooling, mostly of high latitudes, with a major shift in
the earliest Oligocene suggesting probable glaciation
(confirmed by the finding of ice rafted debris in Indian
Ocean drill holes [Zachos et al., 1992b]) and cooling,
especially of bottom waters. Additional glaciations have
been postulated during the Oligocene as a result of other
in two sites just off the Antarctic coastline has been
interpreted to result from the production of warm saline
bottom water in the Eocene and warm saline deep water with
an underlying psychrosphere [Kennett and Stott, 1990].
The oceanic record of carbon isotopes has shown that
changes have occurred more uniformly, with less geographic
variability. The changes have been inferred to be overall
reservoir changes rather than basin-to-basin
fractionation. The low paleogeographic variation has been
suggested to be due to low productivity, especially during
the Oligocene [Miller, 1992; Wright and Miller, 1993].
The balance between continental weathering of 87Sr-
rich rocks and hydrothermal weathering of 86Sr-rich
oceanic crustal rocks controls the marine 87Sr/86Sr ratio
preserved in foraminiferal shells. The middle to late
Eocene 87Sr/86Sr ratio was held low at a near constant
level, indicating a stable continental weathering to
hydrothermal weathering ratio, with domination by the


hydrothermal component. A transition to steadily
increasing conditions in the late Eocene has been used to
infer increased continental weathering rates, either
through increased mountain building [Richter et al., 1992]
or increasing glaciation [Oslick et al., 1994], Not
incidentally, the steadily increasing 87Sr/86Sr values
also allow improvements in Oligocene stratigraphic
resolution, as biostratigraphic events can be dated more
precisely even in the absence of magnetostratigraphy, and
as the ratio itself can provide a unique date once the
curve is correlated to the Geomagnetic Polarity Time Scale
(GPTS).
This dissertation is composed of four chapters, which
serve to examine aspects of late Paleogene
paleoceanography. It is concentrated in the South
Atlantic, but many of its conclusions may be extrapolated
world-wide.
Since the research in this dissertation refers back
to several older data sets, most dated magnetostratigraph-
ically, the first chapter presents a compilation and
correlation of the 12 comprehensive GPTSs developed since
1968. The correlations enable conversion of dates
generated under older time scales to newer ones. The
correlation table has been placed on-line at the National
Geophysical Data Center to enable other researchers to
have easy access to it.


Strontium isotope stratigraphy from the middle Eocene
through the Oligocene is the topic of Chapter 2. A well
dated, nearly continuously deposited and recovered section
from Ocean Drilling Program (ODP) Hole 689B, a key high
latitude site, was used to generate a fairly high
resolution (166 ky/sample) 87Sr/86Sr record. This record
was used in two ways. First, the 87Sr/86Sr record was
used to form a refined strontium isotope stratigraphy from
the mid Eocene to the end of the Oligocene. This will
allow further researchers to correlate biostratigraphic
markers to the GPTS. Second, the record forms the basis
for modelling experiments to determine the tectonic
changes controlling the marine 87Sr/86Sr ratio. Since
both hydrothermal activity [Edmond et al., 1979; Owen and
Rea, 1985; Lyle et al., 1987] and mountain building [Raymo
and Ruddiman, 1992] can have implications for affecting
the Earth's climatic history, it is important to
understand what these changes are.
The third chapter uses 6180 and <513C to examine deep
circulation changes in the Antarctic and South Atlantic
Oceans, especially with respect to Warm Saline Deep Water.
It uses pre-existing data from ODP Sites 689 and 690
[Kennett and Stott, 1990] with new data from ODP Sites 699
and 703 to examine the vertical in this region during the late Paleogene, and examines the
changes in circulation which may have taken place during


those times, the loci of deep- and bottom-water formation
and the influence of glaciation.
The final chapter examines the genesis of
Braarudosphaera oozes in the South Atlantic by examining
613C and Drilling Project sections. These data are analyzed in
terms of their possible relationships to vital effects,
productivity, shallow- and intermediate-depth circulation
and how the circulation may have been affected by
Oligocene glaciation.
Collectively, these studies have two purposes.
First, they act to improve stratigraphic correlation and
resolution in marine sections of the late Paleogene.
Second, these studies are intended to help delineate
circulation in shallow, intermediate, and deep parts of
the South Atlantic, and to help us understand how this
circulation affected and was affected by changes in the
Earth's paleoclimatic, paleogeographic, paleoceanographic
and paleotectonic regimes.


TABLE OF CONTENTS
Page
ACKNOWLEDGMENTS iii
PREFACE V
ABSTRACT X
CHAPTERS
1 THE CORRELATION OF GEOMAGNETIC
POLARITY TIME SCALES 1
Introduction 1
Calibration of the Time Scales 5
Correlation Between Time Scales 13
Discussion 21
Conclusion 26
2 CONTROLS ON THE 87Sr/86Sr COMPOSITION OF
SEAWATER FROM THE MIDDLE EOCENE TO OLIGOCENE:
HOLE 689B, MAUD RISE, ANTARCTICA 28
Introduction 28
Methods And Materials 33
Results 43
Discussion 50
Conclusion 92
3 LATE EOCENE TO OLIGOCENE VERTICAL OXYGEN
ISOTOPIC GRADIENTS IN THE SOUTH ATLANTIC:
IMPLICATIONS FOR WARM SALINE DEEP WATER 95
Introduction 95
Methods 99
Results 113
Discussion 121
Conclusion 151
4 OLIGOCENE BRAARUDOSPHAERA OOZES AND CHALKS:
GENESIS AND PALEOCEANOGRAPHIC IMPLICATIONS
FROM ISOTOPIC ANALYSES OF FINE FRACTION AND
FORAMINIFERA 154
Introduction 154
Methods 166
Results 171
Discussion 204
Conclusion 220
APPENDICES
A LARGE DATA TABLES 223
1 Correlated Geomagnetic Polarity
Time Scales from 1968 to 1994 224


2 Sr, O, C Isotope and Sr/Ca Data,
ODP Hole 689B 239
3 Oxygen and carbon isotope analyses of
Cibicidoides spp. for ODP Hole 703A 253
4 Oxygen and carbon isotope analyses of
Cibicidoides spp. for ODP Hole 699A 262
5 Oxygen and carbon isotope data from
DSDP Site 363, cores 5, 6, and 7 269
B DETAILED METHODS AND DERIVATIONS 276
1 Instructions for obtaining and using
on-line data tables 277
2 Strontium extraction method 280
3 Operating instructions for thermal
ionization mass spectrometer 282
4 Preparation of foraminifera for
oxygen and carbon isotopic analysis 287
5 Derivation of equations 291
REFERENCES 294
BIOGRAPHICAL SKETCH 315


Abstract of Dissertation Presented to the Graduate School of
the University of Florida in Partial Fulfillment of the
Requirements for the Degree of Doctor of Philosophy
PALEOCEANOGRAPHY OF THE SOUTH ATLANTIC OCEAN
FROM THE MIDDLE EOCENE THROUGH THE OLIGOCENE
By
GREGORY A. MEAD
May, 1995
Chairman: Dr. David A. Hodell
Major Department: Geology
In middle Eocene through Oligocene time, the Earth
experienced major changes, including major plate tectonic
events and the transition from relatively warm conditions to
polar ice and high planetary thermal gradients. These
paleoclimatic and tectonic events led to changes in the
isotope geochemistry of the oceans, the basis for this
dissertation.
Marine microfossil warmest temperatures of the Cenozoic were during the early
Eocene. Planktic and deep-sea benthic foraminiferal 5180
data suggest that temperatures began to decrease from that
time in a series of rapid changes separated by periods of
relative stasis. Despite these Eocene climatic changes, the
marine 87Sr/86Sr ratio stayed relatively constant, implying
that the balance between continental erosion and
hydrothermal input to the oceans was unchanging.


During the Eocene, Program (ODP) Sites 689 and 690 show a reversal in the
normal temperature gradient, interpreted as mid-water warm,
salty deep waters underlying cooler, fresher, shallow
waters. Additional data from ODP Sites 699 and 703 show
that at least by the late middle Eocene, this warm, saline
deep water mass extended to abyssal depths.
Strontium isotopic data from ODP Site 689 show that in
the late Eocene (about 35.5 Ma), a profound, sudden change
in the crust's geochemical balance occurred as marine
87Sr/86Sr ratios began to rise rapidly, probably as
hydrothermal alteration decreased relative to continental
weathering. The timing of this event suggests that it was
not due to increased continental weathering due to either
glaciation or uplift. In the earliest Oligocene (about 33.5
Ma), marine 6180 values increased by about l/oo at all
paleodepths in the deep South Atlantic, indicating an
increase in continental ice volume and/or lower temperature.
This change, estimated to be due 35% to increased Antarctic
ice volume and 65% to temperature decrease, left the ocean's
vertical temperature structure unchanged. Oxygen isotopic
evidence suggests that at this time, glacial ice became
semi-permanent on Antarctica, waxing and waning through the
mid-Oligocene and beyond.


During the Oligocene, repeated Braarudosphaera oozes
were deposited. These unusual deposits may have resulted
from upwelling of cold, low-salinity, nutrient-rich water
originating from melting of an Antarctic ice shelf during
glacial/interglacial transitions.


CHAPTER 1
THE CORRELATION OF GEOMAGNETIC POLARITY TIME SCALES
Introduction
Knowledge in geology advances with time, just as it
does with all sciences. However, geology differs from other
sciences in that the baseline by which all other data are
calibrated, the geologic time scale, changes as the
knowledge changes. In most cases, the ordering of events in
the geologic time scale does not change. However, spacing
between events does change. This causes our understanding
of rate-dependent geologic processes to change (e.g.,
sedimentation rates, freguency of glaciations, tectonic
plate motions, or evolutionary rates) [Kominz, 1984; Delaney
and Boyle, 1988],
The geologic time scale, especially among marine
geologists, is calibrated to the marine magnetic anomaly
pattern (the Geomagnetic Polarity Time Scale, GPTS) for the
past 100 Ma or so. Since the advent of plate tectonic
theory [Vine and Matthews, 1963] when it was realized that
marine magnetic anomalies (MMA) are created during the
seafloor spreading process, several different GPTSs based on
the MMA pattern have been developed, beginning with the
1


2
pioneering time scale of Heirtzler et al. [1968], A GPTS is
based on the concept that seafloor anomalies record the
history of geomagnetic polarity reversals. Time scales
based on MMAs differ in a number of ways, including the
particular anomaly pattern used, the method of dating tie-
points within the pattern, the method used in dating
anomalies between tie-points, the radiometric decay constant
used for radiometric dates, and the particular anomalies
included or excluded from the time scale.
Since 1968, at least 12 different, comprehensive,
geomagnetic polarity time scales covering the last 80 or 100
Ma (back to Anomalies 33 or 34) have been constructed, along
with a host of partial time scales which cover smaller parts
of the polarity record. As a result, comparing studies
which use different time scales can be difficult, because
events may be dated at different times which may be
simultaneous, or events which appear to be simultaneous may
in fact be diachronous.
Unfortunately, the problem is worsened by the fact that
in very few of the revisions are direct comparisons made
between the revised and older time scales. Usually, the
ages of periods of normal polarity are given, often with
little or no identification. Since revisions often distort
some of the polarity intervals, and insert or delete others,
it becomes a non-trivial matter to determine which polarity
interval of one time scale is correlative to a polarity
interval of another time scale.


3
This paper attempts to give a comprehensive correlation
of 12 geomagnetic polarity time scales over the past 26
years. It does not in general discuss the relative merits
of each time scale. Comparisons of preceding time scales
generally can be found in each successive revision, but
especially in Ness et al. [1981] and Harland et al. [1990].
This paper instead is meant to be a resource by which
investigators can calibrate studies which use different time
scales, or to bring older studies up to date. It presents a
brief description of the method used for calibration of each
time scale to geochronology, and a graphic and tabular
correlation of the time scales to one another. The means by
which time scales are correlated are discussed, and
ambiguous regions are noted. In addition, during the
preparation of this paper, a total of 6 errors were
identified in 5 published time scales. They are listed in
Table 1-1 along with the way that the errors were
identified. In most cases, it is unlikely that researchers
re-entering the time scales for their own uses would not
notice them, but they are presented here for the sake of
completeness. Finally, to help others avoid the onerous
task of typing in each of the time scales for their own use,
a copy of the table which correlates the time scales will be
placed on-line (see Appendix Al) so that it may be
downloaded for the use of other researchers (see Appendix B1
for instructions) .


4
TABLE 1-1
List of typographical errors in published
Geomagnetic Polarity Time Scales
Time
Scale
Datum
Incorrect
Date
Correct
Date
How Error Was
Recoanized
HDHPL68
Within lr
1.93
0.93
Out of order
TM7 6
Top 5111
12.49
16.49
Out of order
MD79
Top 12
33.32
33.22
Not correctly
converted from
LKC77.
NLC81
Within 17
38.57
39.57
Out of order
NLC81
Within 17
38.64
39.64
Out of order
GTS 8 2
Top of
0.92
0.90
Supposed to be
Jaramillo
identical with
MD79.


5
Calibration of the Time Scales
The Heirtzler et al. [1968] (HDHPL68) time scale was
based on a single magnetic anomaly record from the South
Atlantic, the V-20SA magnetic profile from about 30S, on
the western side of the Mid Atlantic Ridge. Heirtzler et
al. [1968] examined several magnetic profiles from the
Atlantic, Pacific, Indian, and Antarctic Oceans and
concluded that the South Atlantic was most likely to have
undergone spreading at a constant rate. This anomaly
profile was to become the standard, with modifications, for
most subsequent time scales. Only two dates for the profile
were used: 0 Ma for the center of the Mid Atlantic Ridge,
and 3.35 Ma for the Gauss/Gilbert boundary [Doell et al.,
1966]. This resulted in a calculation of 1.9 cm/yr as a
spreading rate, which they extrapolated to the entire
anomaly pattern. The result was an age of 76.33 Ma for the
top of Anomaly 32, and a slightly adjusted age of 3.37 Ma
for the Gauss/Gilbert boundary. The ages of normal polarity
intervals are presented in their Table 1 without
identification, although the major anomalies are identified
graphically in their Figure 3 (which in some cases differs
from the table; see the last two normal polarity intervals
in their table and figure, for instance). Despite the
extrapolation of constant spreading rates to more than 22
times the length of the dated portion of the anomaly
pattern, the HDHPL68 time scale has remained within a few
percent of succeeding time scales.


6
Tarling and Mitchell [1976] (TM76) revised the HDHPL68
time scale in two major ways. First, they introduced
additional radiometric calibration points from sedimentary
sequences for the anomaly pattern. They dated the top of
Anomaly 6 at 19.5 Ma, with Anomalies 13, 24, and 30 centered
at 35, 48, and 64-65 Ma, respectively. They then linearly
corrected all reversal boundaries to these tie points.
Secondly, they inserted additional polarity intervals into
Anomaly 2 based on reversals found in deep sea sections, and
into Anomaly 5 based on reversals found in deep sea sections
and in MMA patterns by Blakely [1974]. They designated
additional polarity intervals with superscript roman
numerals such as Anomalies 21, 5IV, etc. As with the
HDHPL68 time scale, ages of normal polarity intervals are
not identified in the tabular listing of ages, only
graphically.
A widely used major revision of the GPTS was developed
by LaBrecque et al. [1977] (LKC77). They started with the
original HDHPL68 time scale, and added an additional
calibration point at the base of Anomaly 29 of 64.9 Ma.
This was based on an age of 65 Ma for the
Cretaceous/Tertiary (K/T) boundary, which was found just
below Anomaly 29 in a section at Gubbio, Italy [Lowrie et
al., 1976]. They also used a more recent [Dalrymple, 1972]
age for the base of the Gauss (Anomaly 2A) of 3.32 Ma,
adjusted the length of Anomaly 4A, included the short
reversed intervals in Anomaly 5 of Blakely [1974], inserted


7
an extra reversal in Anomalies 23 and 24, spliced in a
revised sequence from Anomalies 29 to 34, and most
importantly, removed Anomaly 14 from the GPTS, which was not
found in most anomaly patterns. In addition, LKC77 used
radiometric ages directly for anomalies younger than 2A
[Klitgord et al., 1975], and extrapolated the GPTS through
Anomaly 34, to 108.17 Ma. The new time scale was presented
with brief identification of periods of normal polarity in
both tabular and graphic form. This time scale re-appeared
in the Larson et al. [1982] time scale, with extensions to
the Mesozoic by Larson and Hilde [1975] and Van Hinte
[1976a,b].
Later in 1977, new decay constants were adopted for
K/Ar dating [Steiger and Jger, 1977]. All subsequent time
scales used the new decay constants in calculating K/Ar
dates. To take into account the new decay constants,
Mankinen and Dalrymple [1979] (MD79) revised the LKC77 time
scale by recalculating each of the LKC77 time scale datums
to new values. They presented their results in a side-by-
side comparison to the LKC77 time scale, along with anomaly
identifications, simplifying conversion from one to the
other time scale. MD79 further revised the late Neogene (5-
0 Ma) based on additional radiometric dates later in their
paper (their Figure 3); those dates are presented in the 5-0
Ma portion of this compilation.
A comprehensive review of previous time scales, along
with the development of a new one, was presented by Ness et


8
al. [1980] (NLC80). The new NLC80 time scale was based in
general on the LKC77 time scale corrected to the new K/Ar
decay constants, but used segments from several other time
scales to adjust the length of anomalies, and to insert a
few new reversals. Yet another terminology was used in the
NLC80 time scale, with decimal numbers, and then primes, to
designate successively smaller intervals after the main
anomaly numbers. A total of 4 calibration points, plus the
present, were taken from a wide variety of sources, and
dates were interpolated between them. Anomaly 2.3'(o) (for
"older end") was fixed at 3.40 Ma, with dates extrapolated
to Anomaly 3.4; Anomaly 4.1'(y) was fixed at 7.81 Ma;
Anomaly 24(o) was fixed at the Paleocene/Eocene boundary
(54.9 Ma); and the K/T boundary, just below Anomaly 29(o),
was fixed at 66.7 Ma. The study compared in graphical form
the dated anomaly sequences and time scales used in all the
previous studies discussed in the present paper, as well as
several more fragmentary sequences. A synthetic anomaly
profile was generated which facilitates comparison to other
anomaly sequences. Major anomalies and correlations between
the various time scales were made (and in general are
followed in this paper) although a few ambiguities remained.
Dates attached to the anomaly patterns were unfortunately
printed in a font so small as to be nearly unreadable.
Another time scale appeared just a year later by Lowrie
and Alvarez [1981] (LA81). This time scale used a total of
11 biostratigraphically defined calibration points (geologic


9
period boundaries), from Italian pelagic sections (as well
as the 0 Ma point), to modify the LKC77 time scale,
corrected for new K/Ar decay constants. The
biostratigraphic correlations allowed them to adjust the
positions of the anomaly boundaries to agree with additional
radiometric dates. As in most other time scales, the ages
between calibration points were determined by linear
interpolation. Most of the anomaly boundaries were
identified in the table of ages provided.
A geologic time scale covering the entire Phanerozoic
was published in 1982 by Harland et al. (GTS82). It dated
MMA from the present back into the Mesozoic, although for
the sake of consistency it is reported here only through
Anomaly 34. From 5-0 Ma, radiometric dates were used to fix
reversal boundaries. From 83-5 Ma, GTS82 is based on MMA,
modified from LA81. Two modifications were made. First,
the Paleocene and Eocene boundary calibration points were
removed because they implied improbably rapid spreading rate
changes. Points at the beginning and ends of the Paleocene
and Eocene were kept, however. The second modification was
to place the K/T boundary at 65 Ma; some other adjustments
were made to reduce apparent spreading rate changes as well.
Overall, NLC80 was stretched between the new calibration
points from 10.3 to 83 Ma, NLC80 was used directly for 3.4
to 10.3 Ma, and MD79 was used directly from 0-3.4 Ma.
An extremely widely used time scale was next published
by Berggren et al. [1985a] (BKFV85). The same time scale


10
has appeared as Berggren et al. [1985b,c] separately for the
Paleogene and Neogene, and, with extensions, in Kent and
Gradstein [1986] (back to the Jurassic). The major advance
of this time scale was that it integrated biochronologic,
magnetochronologic, and radiochronologic data into a single
time scale. The BKFV85 time scale used fewer calibration
points than some of the recently preceding time scales, only
6 (plus the present). The younger ends of Anomalies 2A, 5,
12, 13, 21, and 34 were fixed at 3.40, 8.87, 32.4, 34.6,
49.5, and 84.0 Ma, respectively, by reference to a host of
different studies. Three best-fit line segments were
constructed from these data, from 0 to 10.42 Ma (the base of
Anomaly 5), from 10.42 to 56.14 Ma (the base of Anomaly 24),
and 56.14 to 84 Ma (the top of Anomaly 34). Ages for the
other paleomagnetic boundaries were interpolated from LKC77
and reported in tabular and graphical form, with most
intervals of normal polarity identified. Perhaps most
importantly, BKFV85 also correlated marine calcareous
planktonic biostratigraphic zonations to the GPTS for the
entire Cenozoic. BKFV85 used the chron nomenclature
introduced by LaBrecque et al. [1983], with a "C" prefix
designating the time during which a given normal and
reversed anomaly couplet was created, and a "N" or "R"
designating the polarity of the Earth's field at the time
within the chron.
Haq et al. [1988] (HHV88) recognized that using a
single MMA pattern or a patchwork of MMA patterns strung


11
together suffers from the problem that spreading rate
changes may not be taken into account. In addition, errors
in radiometric dates of calibration points may introduce
errors in calculated spreading rates. To overcome these
uncertainties, HHV88 plotted numerous radiometric dates
against MMA patterns from 3 ocean basins the North and
South Pacific and the South Atlantic. They calibrated the
MMA to best-fit age-to-length lines, and then stacked (i.e.,
averaged) the ages of the tops of the major anomalies from
each basin to arrive at their time scale. The time scale
has a coarse resolution, with only the tops of Anomalies 1
to 32, 5A to 5E, and 6A to 6C listed.
Harland et al. [1990] (GTS89) modified the original
HDHPL68 anomaly spacing by splicing in parts from the
several different sources, and adjusting spacing using MMA
pattern segments from fast spreading ridges. To date this
modified HDHPL68 record, GTS89 used numerous radiometric
dates to define 3 best-fit line segments, from 0 to 3.4 Ma,
9 to about 50 Ma, and about 50 to about 85 Ma. No data was
used for 3.4 to 9 Ma. These 3 line segments were used to
constrain and convert the modified MMA pattern of HDHPL68,
resulting in the new time scale. No correlation to previous
time scales is shown. However, through an intermediate step
(2 tables), GTS89 does provide an exact match to HDHPL68,
except for Anomalies 4A through 5, which are "very
different."


12
Cande and Kent [1992] (CK92) used two new approaches in
constructing their new time scale. One of the techniques
was the creation of a synthetic MMA pattern in order to
remove the problems associated with using a single MMA
pattern (V-20SA) as a standard. Cande and Kent [1992]
stacked different segments of MMA patterns from the South
Atlantic to reduce the noise and error associated with a
single MMA pattern, using from 5 to 9 different MMA patterns
for every part of the time scale from Anomaly 1 to 34. They
next assumed that spreading rates were not necessarily
constant but instead smoothly varying, and fitted 9 points
(finite rotation poles from Cande et al. [1988]) plus the
ridge axis with a cubic spline function. Their aim was to
reduce the apparent spreading rate changes in the synthetic
MMA record. They also identified numerous "tiny wiggles"
(their Table 7) which they did not include in the main time
scale (their Table 6). Some of the tiny wiggles they
identified (C2r.2r-1, C5n.2n-1 to -3, C5Dr-l, and C13nl) are
included in Appendix Al of this work, because they are
commonly observed in previous time scales. The second
technique use by CK92 was to use two astronomically tuned
dates (0.780 Ma for the Bruhnes/Matuyama boundary and 2.600
Ma for the Matuyama/Gauss boundary) of Shackleton et al.
[1990] and Hilgen [1991]. These dates were obtained by
assuming that Milankovitch periods seen in geologic records
are known more precisely and accurately than radiometric
dates over the same intervals, and counting back to magnetic


13
reversals in these records. Interestingly, tuned ages used
in this time scale and the Cande and Kent [1995, in press]
time scale are all older than radiometric ages by 5 to 10%.
It has been suggested that this is due to an error in the
decay constant of Steiger and Jger [1977] [Shackleton et
al., 1990] or to diffusional loss of radiogenic Ar [Hilgen
1991].
Cande and Kent [1995, in press] (CK95) is modified in
two ways from the CK92 timescale. The later part uses a
total of 19 additional tie-points added according to the
astronomical time scales of Hilgen [1991] from the base of
C3n.4n to the base of C2An.ln and Shackleton et al. [1990]
from the top of C2An.ln to the present. The early part was
adjusted so that the Cretaceous/Tertiary (K/T) boundary
falls at 65 Ma, and another cubic spline used to fit the
data from the top of C34n to the base of C3n.4n as with
CK92. Because of the use of different cubic splines in CK92
and CK95, the difference between the two time scales varies
sinusoidally between 0.2 my except near the K/T boundary,
when CK95 is over 1 my younger than CK92. This time scale
will be incorporated into a comprehensive revision [Berggren
et al., in preparation] of Cenozoic microfossil datums
[Kent, personal communication].
Correlation Between Time Scales
Ideally, every time a new time scale was published,
each period of normal and/or reversed polarity would be


14
completely identified so as to facilitate comparison to
older and newer time scales. Failing that, it would be
helpful if the polarity intervals in each time scale were of
equal number and of roughly equal relative size. Both of
these possibilities are rarely seen.
Another option would be to have each time scale
correlated by referring back to the original anomaly pattern
each is based upon. Unfortunately, the anomaly patterns are
rarely published, and are presented in only 3 of the papers
discussed here (HDHPL68, NLC80, and CK92). It is clear that
these papers, each published 12 years apart, are easily
correlated by comparison of the anomaly patterns (Figure
1-1), despite the different methods used to construct them.
In lieu of the above choices, pattern matching must be
done, similar to the way magnetostratigraphic records in
sedimentary sequences are matched to the GPTS. The
comparison is made easier because, at least in this case,
the entire record is in fact present. This has been done
graphically in two of the papers discussed, NLC80 and CK92.
NLC80 presented their new time scale correlated to other
partial and comprehensive time scales (their Figure 2), with
a few lines of correlation, and much of their technique is
followed here in Figure 1-2. Few lines of correlation were
presented, leaving a great deal of ambiguity in the
correlations. CK92 used a different method, displaying each
polarity record as a series of lines shifting back and forth


AXIS 2 2A 3 3A 4 4A S 9A 9AC 98 5C 50 5E 6A 08 AC 7 0 10 II 12 13 IS 16 17 10 19 20 21 22 23 24 25 26 2 7 28 29 30 31
Figure 1-1: Marine Magnetic Anomaly patterns from HDHPL68, NLC80, and CK92. The MMA
patterns from HDHPL68 and NLC80 are synthetic. The MMA pattern from HDHPL68 was
constructed using the identified anomalies found in profile V-20SA, while the MMA pattern
from NLC80 was constructed with the modified magnetic anomaly sequence used to define the
timescale itself. The MMA from CK92 is the sum of between 5 and 9 anomaly patterns from
the South Atlantic. Note the excellent correspondence, despite the different methods used
to construct each MMA pattern.
ui


Figure 1-2: Correlated paleomagnetic intervals from 12
time scales. Black indicates an interval of normal
polarity; white indicates an interval of reversed
polarity. Lines indicate tops and bottoms of equivalent
polarity intervals. In the region from 34 to 64 Ma, some
lines are dotted or dashed. The differences are merely to
help guide the eye. The abbreviated headings are
explained in the text.


I I I I I I I I I I I I I I I I I I I I I I 1 I I I I I I I 1 I I 1 1 I I I I I I I I
HDHPL68
TM76
LKC77
MD79
NLC80
LA8 I
GTS8 2
BKFV85
HHV8 8
GTS89
CK9 2
CK95
Q
o
a


I I .-11-
r~T
i
HDHPL68
TM7 6
LKC77
MD79
NLC80
LA8 I
GTS8 2
BKFV85
HHV8 8
GTS89
CK9 2
CK95
00
O
o
CD
U1
O)
O
Ol
(J1
o
o
Q
Q
Q
CM
Figure 1-2continued


LJ I I I I LJ I I I I I I I I I|L
O
HDHPL68
TM76
LKC7 7
MD79
NLC80
LA8 I
GTS8 2
BKFV85
HHV88
GTS89
CK9 2
CK95
VO
Figure 1-2continued


20
from the right (Normal polarity) to the left (Reversed
polarity) for their time scale and 5 preceding time scales.
In this case, no lines of correlation were used, and
the figure was published at such a small scale that fine-
scale correlations cannot be made.
A hierarchy of methods is used in this paper to
correlate the time scales. First and foremost, the authors'
designation of a polarity interval, if present, was used to
make a first-order correlation. Although not all time
scales were published with identifying labels in the table
of ages, generally accompanying figures could be used for
the same purpose. Second, in the absence of sufficient
labeling (and it should be noted that only GTS82, CK92, and
CK95 time scales had sufficient unique identifiers for all
periods of normal and/or reversed polarity), the relative
lengths of the polarity intervals (i.e., pattern matching)
were used to make correlations. This method succeeded in
making almost all of the detailed correlations left after
the first-order correlations. The ambiguities that remain
generally result from multiple short reversed or normal
events which are not present in all time scales, and whose
relative position cannot be used to determine which of the
multiple events may or may not be present in a particular
time scale.
The result is Appendix A1 and Figure 1-2. Within
Figure 1-2, lines of correlation are drawn for the tops of
each major polarity chron, with additional lines drawn to


21
clear up any residual ambiguities. Asterisks (*) indicate
the ambiguities which remain.
Discussion
A GPTS is usually used to date stratigraphic seguences.
An age model for a seguence is developed using identifiable
magnetostratigraphic events within the GPTS, or
biostratigraphic datums that are calibrated to the GPTS.
Sample ages are then calculated by linear interpolation from
the position of samples within the stratigraphic sequence.
Linear interpolation will also generally be used to make the
conversion between one time scale and another. Two factors
must be taken into account when converting between time
scales, however.
First, in constructing a conversion table of two GPTSs
from Appendix Al, only dates (i.e., paleomagnetic chrons)
common to both time scales can be used. For example, if
conversion is made between the BKFV85 and CK92 time scales,
Chron C2r.2r-1 is present only in CK92, not in BKFV85, and
must be left out of a conversion table. Similarly, Chron
C23R-1, found in NLC80, GTS82, and GTS89, is not present in
either BKFV85 or CK92. Obviously, Chron C23R-1 cannot be
used in the conversion table, either. Therefore, Appendix
Al must be edited to remove the unusable chron intervals.
Second, even an edited Appendix Al cannot in general be
used to convert ages from one time scale to another directly
from the nearest two dates common to both time scales. The


22
reason is that the age model used to date the samples,
usually generated from paleomagnetic boundaries, usually
only uses selected polarity transitions in a time scale.
Since the relationship between time scales is usually not
linear, the age model must be converted, and that converted
age model used to determine ages directly from stratigraphic
sequences. For example, suppose columns 2 and 3 of Table 1-
2 show the depths and ages of several samples in a given
stratigraphic sequence, using the BKFV85 time scale. If the
datums listed in column 1 are known, then the ages of the
samples can be easily converted by using the relationship of
BKFV85 to CK92 for all points, and are listed in column 4.
But if only 2 datums, the top of C12 and the top of C20, are
known (e.g., because of coring disturbance or lithologic
problems in the section between those datums), then the age
model used to calculate ages from the depths would consist
of only 2 points, 32.46 Ma at 12.46 m, and 44.66 Ma at
24.66 m, and ages calculated by linear interpolation between
those points. Converting those dates to 30.452 Ma and
42.629 Ma of CK92, respectively, and using linear regression
to calculate the new dates would result in the ages given in
column 5, which differ from the column 4 ages by up to
0.844 Ma (Column 6), a potentially significant error. The
reason that the differences exist is that the relationship
between BKFV85 and CK92 is not linear in this interval
(Figure 1-3).


23
BKFV85 Time Scale (Ma)
Figure 1-3: The CK92 time scale plotted against the BKFV85
time scale from 30-45 Ma. Note that the relationship is
not linear.


24
For a similar reason, if biostratigraphic datums are
being used that have originally been dated
paleomagnetically, then to re-calculate the ages of those
datums, the original citation must be re-examined, and the
ages of the biostratigraphic datums re-calculated based on
the age model used in the original citation. Then the new
ages of the biostratigraphic datums can be used in a given
age model to calculate ages at a particular site.
Therefore, the following steps must be taken in order
to convert a series of dated samples from one GPTS to
another. First, Appendix A1 of this paper must be edited to
remove polarity boundaries not common to both time scales in
question. Second, using linear interpolation, the age model
originally used to calculate ages for the samples must be
converted from the original time scale to the new one.
Finally, the newly converted age model may be used to
calculate ages of the samples.
The set of steps listed above raises two additional
points. In many papers, ages are given without reference to
a particular time scale and/or the age model used to
calculate those ages. Without this information, it is
impossible to evaluate the synchroneity or diachroneity of
events within one study with those of another. Therefore,
in any geologic study, both the time scale and the age model
used should be explicitly stated.
It is of course highly unlikely that the CK92 time
scale will be the final word in GPTSs. The utility of


25
TABLE 1-2
Example of a conversion from one time scale to another
1
2
3
4
5
6
Sample
Direct
Age-Model
Depth
Age
Age
Age
Datum
fm)
BKFV85)
(CK92}
(CK92)
Difference
top
C12
12.46
32.46
30.452
30.452
0.000
top
C13
15.29
35.29
33.050
33.277
0.227
top
C15
17.24
37.24
34.669
35.223
0.554
top
C16
18.10
38.10
35.368
36.081
0.713
top
C17
19.53
39.53
36.665
37.509
0.844
top
C18
21.29
41.29
38.500
39.265
0.765
top
C19
23.60
43.60
41.353
41.571
0.218
top
C20
24.66
44.66
42.629
42.629
0.000


26
future time scale revisions would be greatly enhanced if
they included a detailed point-by-point comparison to at
least one previous version of the GPTS, probably the most
recent preceding one.
Conclusion
Since the development of the first comprehensive GPTS
in 1968, numerous time scales have been developed based on
the seguence of reversals of the Earth's magnetic field.
Twelve GPTSs are examined in this paper and correlated to
each other as well as possible. The correlations are
presented graphically and in table form. Some ambiguities
remain, but the correlations should enable conversion from
one time scale to another.
Direct conversion from one time scale to another cannot
in general be used to re-calculate ages within a
stratigraphic section. Instead, the original age model used
to calculate ages must be converted, and the new age model
must then be used to calculate ages within the stratigraphic
section. The reason is that the relationship between GPTSs
is rarely linear, and it would be very unusual for every
reversal boundary to be identified in any given
stratigraphic section.
The need for this paper results from the fact that most
revisions of the GPTS do not include detailed, useful
comparisons to previous versions of the GPTS. Users of
future revisions would be helped if the authors of those


27
revisions would include a detailed comparison to at least
one pre-existing version.


CHAPTER 2
CONTROLS ON THE 87Sr/86Sr COMPOSITION OF SEAWATER FROM THE
MIDDLE EOCENE TO OLIGOCENE: HOLE 689B, MAUD RISE, ANTARCTICA
Introduction
The time from the middle Eocene through Oligocene saw
major changes in the Earth's climatic and biotic realms.
During this time, the transition from "Greenhouse" to
"Icehouse" conditions occurred [Miller et al., 1991a,b].
For example, ice rafted debris (IRD) from Antarctic glaciers
first appeared on the Kerguelen Plateau in the earliest
Oligocene at about 35.8 Ma (Berggren et al., 1985 time
scale), coeval with the world wide increase in S180 values
of marine carbonates [Zachos et al., 1992b; Breza and Wise,
1992; Barrera and Huber, 1993]. By the late Eocene, the
planetary thermal gradient increased [Keigwin and Corliss,
1986; Zachos et al., 1992a]. Major plate tectonic events in
the middle Eocene to Oligocene include the opening of the
Drake Passage [Barker and Burrell, 1977], the progressive
closure of the Tethys and uplift of the Alps [Dercourt et
al., 1986], collision of the Indian subcontinent with
Eurasia and subsequent uplift of the Himalayan-Tibetan
Plateau [Harrison et al., 1992], and a middle Eocene change
28


29
in the direction of motion of the Pacific plate [Van Andel
et al., 1975; Clague et al., 1975; Kennett et al., 1985].
The ratio of 87Sr to 86Sr in the oceans is a function
of the balance between continental weathering via riverine
input, dissolution of marine carbonates, and hydrothermal
exchange with mid-ocean ridge basalts [Palmer and
Elderfield, 1985]. When biogenic marine carbonates
precipitate, the 87Sr/86Sr ratio of ocean water is preserved
with no fractionation as Sr substitutes diadochically for
Ca. The Sr isotopic composition of the oceans is constant
at any given time due to the long residence time of Sr in
the oceans (about 2.5 Ma [Hodell et al., 1990]). The
87Sr/86Sr ratio thus has great potential as a stratigraphic
tool. Although the controls on the Sr budget of the ocean
are thought to be well understood [Palmer and Elderfield,
1985], the actual factors that influence changes in the
87Sr/86Sr seawater curve are often difficult to determine
[Hodell, 1994].
The ratio of 87Sr to 86Sr in marine carbonates has
changed irregularly through the Phanerozoic [Burke et al.,
1982, Hess et al., 1986], and increased, with some intervals
of stasis, throughout the Cenozoic [Palmer and Elderfield,
1985, Koepnick et al., 1985, Hess et al., 1986]. During the
Eocene, oceanic 87Sr/86Sr ratios were approximately
constant, although somewhat variable, and averaged from
about 0.7077 to 0.7078 [DePaolo and Ingram, 1985, Hess et
al., 1986]. In the late Eocene, oceanic 87Sr/86Sr ratios


30
began to increase, rising fairly linearly to about 0.7082 at
the end of the Oligocene with a slope reported at from
34xl0-6/m.y. [Miller et al., 1988] to 29xl06/m.y. [Hess et
al.f 1989], The transition to more rapidly increasing
Oligocene values is not well documented, however. Lower
resolution studies [Burke et al., 1982; Koepnick et al.,
1985; Palmer and Elderfield, 1985; Hess et al., 1986] place
the change between 40 and 35 Ma, according to the Berggren
et al. [1985] time scale. The most complete published
87Sr/86Sr records available for the Oligocene do not reach
the transition and/or are not paleomagnetically dated
[Miller et al., 1988; Hess et al., 1989]. In addition, Site
522 (analyzed by Miller et al. [1988]) is at least partially
recrystallized [Hess et al., 1989]. Finally, the resolution
of the two studies was not particularly high (about 0.4-0.5
m.y./sample), especially in light of recent studies that
suggest short term 87Sr/86Sr variation [Capo and DePaolo,
1990; Dia et al., 1992; Clemens et al., 1993].
The cause of the inflections in the 87Sr/86Sr curve in
the late Eocene is not known. There are a number of
tectonic and environmental factors that may have influenced
the input of Sr into the oceans near this time. Glaciation
occurred on Antarctica in the earliest Oligocene, which
could have increased the weathering rates on that continent
[Zachos et al., 1992b; in preparation]. Uplift of the
Himalayan-Tibetan Plateau following the collision of India
with Eurasia may have increased continental weathering rates


31
[Richter et al., 1992]. In addition, major sea level drops
in the middle Eocene, earliest, "mid", and late Oligocene
occurred [Haq et al., 1988], and the Carbonate Compensation
Depth (CCD) deepened near the Eocene/Oligocene boundary [Van
Andel et al., 1977], possibly reflecting a change in oceanic
circulation, shallow/deep calcium carbonate fractionation,
or continental weathering rates. Plate reorganization in
the Pacific may also have affected the hydrothermal
component and/or direction of plate convergence [Lyle et
al., 1987].
This paper presents the 87Sr/86Sr record from the
middle Eocene to Oligocene of ODP Hole 689B (Maud Rise, near
the coast of Antarctica; Figure 2-1). The purpose of this
study is (1) to refine the Sr isotope curve during the
middle Eocene-01igocene, (2) to help calibrate Antarctic
biostratigraphy to the Geomagnetic Polarity Time Scale
(GPTS), and (3) to improve our understanding of geochemical
cycling of Sr during that time. In addition, additional
oxygen and carbon isotopes of Cibicidoides were analyzed in
order to cross calibrate the results from the work of
Kennett and Stott [1990] and Chapter 3, and to improve the
resolution of the 5180 and This site contains a nearly continuous and well preserved
calcareous and siliceous biostratigraphic record of the
Eocene to Oligocene, from very high latitudes. As such, it
constitutes a vital tie-point for biostratigraphic


32
Figure 2-1: Location map, showing ODP Site 689 and nearby
Site 690.


33
correlation. It also is an important site for biogeographic
and isotopic paleoceanographic studies [Barker, Kennett et
al., 1988, 1990; Chapter 3 of this dissertation].
Methods And Materials
Stratigraphy
Cenozoic stratigraphy is currently in a state of
change, with the standard time scale of Berggren et al.
[1985] (BKFV85) being replaced by Cande and Kent [1992]
(CK92). Ages in this paper will be given using the newer
time scale, although the BKFV85 ages are reported in the
data tables as well as the newer CK92 ages. Where
appropriate, BKFV85 ages are given in parentheses.
The age model used is based on the magnetostratigraphy
of Spie/3 [1990] and the age model of Kennett and Stott
[1990], modified to the CK92 time scale (Table 2-1, Figure
2-2). Modifications have been made to these models in a few
instances discussed below.
First, an unconformity exists at 161.54 mbsf that was
not recognized by Kennett and Stott [1990] but was
identified by Spie/9 [1990]. Without this hiatus, it is
difficult to reconcile the pattern of normal and reversed
polarity events between about 135 and 190 mbsf. This
unconformity has been added, with the ages of the upper and
lower surfaces (defined by extrapolation of sedimentation


34
TABLE 2-1
Age Model of ODP Hole 689B
Preferred
BKFV85
CK92
Depth
Model
Age
Age
(mbsf)
Datums
(Ma)
(Ma)
60.40,
Top C5Dr-2
18.14
17.893
61.76,
Base C5Dr-2
18.56
18.317
65.50,
Hiatus, C5En
19.09
18.817
65.50,
Hiatus, by extrap.*!
20.05
19.770
66.86,
Hiatus, C6n
20.45
20.162
66.86,
Hiatus, >Top C7nlr
25.60
24.772
67.11,
Top C7n-2
25.67
24.826
75.97,
Base C8n
27.74
26.533
SK
79.46,
Top C9n
28.15
27.004
SK
91.93,
Base ClOn
30.33
28.716
SK
103.37,
Base Clin
32.06
30.071
SK
106.87,
Base C12n
32.90
30.915
SK
116.71,
Top C13n
35.29
33.050
S
119.69,
Base C13n
35.87
33.543
SK*3
124.09,
Top C15n
37.24
34.669
125.07,
Base C15n
37.68
34.959
128.32,
Top C16n
38.10
35.368
K
134.02,
Base C16n-1
38.79
35.977


35
TABLE 2-1continued
Depth
(mbsf)
Preferred
Model
Datums
BKFV85
Age
(Ma)
CK92
Age
(Ma)
135.75,
Top C17n
39.53
36.665
K*4
152.73,
Base C17n3
41.11
38.183
SK
153.70,
Top C18n
41.29
38.500
SK*5
161.17,
Top C18r-2
42.73
40.221
S
161.54,
Hiatus, by extrap.*2
42.80
40.312
S
161.54,
Hiatus, by extrap.*1
43.66
41.387
S
163.16,
Base C19n
44.06
41.617
SK
165.55,
Top C20n
44.66
42.629
SK
171.11,
Base C20n
46.17
43.868
S
183.23,
Top C21n
48.75
46.284
S
S: Datum from Spie/J [1990].
K: Datum from Kennett and Stott [1990].
*1: Age calculated by assuming constant sedi
mentation rate from underlying section.
*2: Age calculated by assuming constant sedi
mentation rate from overlying section.
*3: Datum listed by KS90 as Top C13n.
*4: Age listed by KS90 as 39.52 Ma.
*5: Age listed by KS90 as 41.21 Ma.


Figure 2-2: Magnetostratigraphy and interpretations, 87Sr/86Sr and 689B, plotted against depth.


0.7085H
0.7084
0.7083
0.7082-
0.7081-
0.7080
0.7079
0.7078
0.7077-
81ao (PDB)
87 n ,86 n n .
Sr/ Sr Ratio
I upper Oligo. lower Oligocene I upper Eo.
++
60
1
1
III
!sC7l
?6i res
C9
C10
C11|
[C12
C13
o
O)
o
CO
_o
i i r v
)0 110 120
C16
middle Eocene
mi hi
o
1
2
C17
i
C18$ IC20
C9|
r^~it 1 r
C21
Depth (mbsf)
OJ


38
rates to the position of the unconformity) converted from
the BKFV85 to CK92 time scale.
Second, there is a disagreement between the age models
from 128.33 to 137.77 mbsf (lower upper Eocene). The series
of polarity intervals identified as Chrons C15n-l/2 to C15r-
2 by Spie/3 [1990] are identified as Chrons C16n-1 to C16r-3
by Kennett and Stott [1990]. The CP15b/CP15a calcareous
nannofossil boundary (recognized at 131.7-132.7 mbsf by Wei
and Wise [1990]) is closely associated with the C15n-l/2 to
C15r-2 boundary [Berggren et al., 1985], and this controls
Spiel's [1990] position of that paleomagnetic boundary at
134.02 mbsf. However, Wei [1991] and Wei and Wise [1992]
showed that the first occurrence of Isthmolithus recurvus.
which defines the CP15b/CP15a boundary, occurs somewhat
earlier in the Southern Ocean, within C16. Using Spie/3's
[1990] identification also introduces large distortions into
the sedimentation history, with sudden jumps in the
sedimentation rates within C15n and C16n. As a result, I
accept the chron assignments of Kennett and Stott [1990] in
this interval. This chron assignment also re-defines the
normally magnetized interval between 124.09 and 125.07 mbsf
from a normal within C13r-2 [Spie0, 1990] to C15n,
considerably improving the pattern match between the
magnetic stratigraphy and the seafloor magnetic anomaly
record.
Finally, two unconformities occur in the upper part of
the section: one at approximately 66.86 mbsf and the other


39
at about 65.5 mbsf. Only the lower one is recognized
paleomagnetically [Spie/3, 1990]. The second falls within a
normally magnetized interval and was identified only through
the use of biostratigraphy [Gersonde and Burckle, 1990].
Neither hiatus is recognized by Kennett and Stott [1990],
However, both are clearly identified by the use of Sr
isotope stratigraphy (see RESULTS). The 87Sr/86Sr ratio of
sediments within the short normally-magnetized interval
between the unconformities (about 0.70843) indicates that
the age is approximately 20 Ma [Hodell and Woodruff, 1994],
assignable to Chron C6n [Cande and Kent, 1992]. The
sediments immediately above the upper unconformity have a
87Sr/86Sr ratio of about 0.70851, and are approximately 18.7
Ma [Hodell and Woodruff, 1994], assignable to Chron C5En
[Cande and Kent, 1992].
Materials
Benthic and planktic foraminifera from the >250/m,
>212/xm, or >150nn fraction (depending on abundance) were
picked from washed (>63/im) samples from Hole 689B.
Preservation of foraminifera is generally good [Stott and
Kennett, 1990, Kennett and Stott, 1990]. In addition,
analysis of interstitial waters from Hole 689B [Egeberg et
al., 1990] indicates unusually low recrystallization rates
of carbonate. To further check the degree of preservation,
Sr/Ca ratios were measured as well (see Results).


The sampling interval is somewhat variable, ranging
between 0.5 and 3 m, or from 50 to 660 ky. The average
40
sampling interval was 0.96 m or 166 ky. Below 142 mbsf, the
sampling interval was coarser, with one sample per 1.44 m or
211 ky, while above that depth, the sampling interval
averaged one per 0.81 m or 134 ky. Because the amplitude of
the 87Sr/86Sr signal is much smaller below 135 mbsf, the
lower sampling rate below 142 mbsf is adequate to delineate
the 87Sr/86Sr curve.
The samples were ultrasonically cleaned in methanol to
remove adhering fine debris and dried in a 60C oven. They
were then dissolved in 1 to 2 ml of 0.25 N HC1 for about 1
hour at 200-250C, then dried under a laminar flow hood.
Sample sizes range from about 1.3 to 10.0 mg. Strontium
concentrations were about 2000 ppm according to isotope
dilution techniques, and about 1350 ppm according to atomic
absorption spectrophotometry, giving sample Sr masses of 1.8
to 20 nq.
Ion exchange columns, made of Teflon shrink-tubing and
packed with ElChroM Industries, Inc. Sr-selective crown
ether resin, were used to extract strontium from the
samples. All extractions were made under a laminar flow
hood. A modification of the method of Pin and Bassin [1992]
was used with slightly differing volumes of HN03 and H20
(Appendix B2). Yields were approximately 75%, with
excellent separation of Sr from Ca [Pin and Bassin, 1992].


41
Mass Spectrometry
Strontium isotopes
All Sr samples were analyzed in a VG Micromass 354
thermal ionization mass spectrometer in dynamic mode
(instructions for operation are listed in Appendix B3).
Samples were loaded onto oxidized Ta single filaments by
dissolving the extracted Sr in 2-3il of 0.5 N H3P04 acid.
An analysis of NBS-987 was run with each turret of 14
samples. Analyses of 17 NBS-987 standards averaged
0.710240, with an uncertainty (2 standard deviations or 2a)
of 0.000022. All samples are normalized to NBS-987 =
0.710235 and 86Sr/88Sr = 0.1194. Samples for which the
86Sr/88Sr ratio was determined to be less than 0.1194 are
assumed to have been fractionated and those values were not
reported.
To examine reproducibility of the results, two
different types of replicates were analyzed. Separate
analyses of planktonic and benthic foraminifera were made in
17 samples. The 87Sr/86Sr ratios of planktonic foraminifera
averaged 8xl06 (ct=20x106) higher than those of benthic
foraminifera, showing no systematic difference between Sr
isotopic ratios of the different types. Because in some
cases less than perfectly preserved specimens of
foraminifera had to be picked due to lack of sample, I also
analyzed the difference between well preserved and poorly
preserved (visibly damaged, probably dissolved) specimens


42
when possible (5 cases). The 87Sr/86Sr ratios of well
preserved specimens averaged only 2xl0-6 (a=13xl06) higher
than those of poorly preserved specimens, which, despite the
small number of analyses, suggests that no systematic
difference exists between Sr isotopic ratios of
differentially preserved specimens. The poor preservation
of some foraminiferal tests is attributed to dissolution
rather than recrystallization. The average difference of
all 22 replicates is 16xl0-6 (a=12xl0-6), not much larger
than the standard deviation of the NBS-987 standard.
Oxygen and carbon isotopes
Well-preserved specimens of the benthic foraminiferal
genus Cibicidoides were picked from the washed >150^m
fraction of 35 samples from Hole 689B. Standard methods
were used to prepare and analyze the carbonate (see Chapter
3). The samples were analyzed in a VG Prism triple
collector mass spectrometer with an automated preparation
system. Of the 35 samples run, 14 were replicate analyses
of samples run by Kennett and Stott [1990] in an attempt to
calibrate their data with that of Chapter 3. The analytical
precision of standards (1 for S180.
Atomic Absorption Spectrophotometry
A Perkin-Elmer 3100 Atomic Absorption Spectrometer was
used to measure Sr/Ca ratios. About 50 ng of carbonate


43
(typically 3-6 foraminifera) was dissolved in 5 ml of 0.5 N
Optima (trace metal grade) nitric acid. The Ca
concentrations (on the order of 3 ppm by weight) were
measured in flame mode, while the Sr measurements (on the
order of 10 ppb by weight) were analyzed using the graphite
furnace, on the same aliquot of sample. Both sets of
measurements were compared to standard absorbance curves to
calculate concentrations. Results are reported in both mass
and molar ratios (Appendix A2).
Results
Strontium Isotopic Measurements
From the base of the section analyzed in this study
(183 mbsf) to 133.2 mbsf, the value of the 87Sr/86Sr ratio
shows little variability, averaging 0.707735 (a=0.000018)
(Figure 2-2, Appendix A2). When plotted against age, the
87Sr/86Sr ratios remained steady or possibly decreased
slightly from an average of 0.707754 (a=0.000014) at 46.2 Ma
to the middle Eocene hiatus beginning at 41.4 Ma (162.54
mbsf) (Figure 2-3). By the time sedimentation resumed at
40.3 Ma, the 87Sr/86Sr ratio had dropped to near 0.70771.
The 87Sr/86Sr ratio averages 0.707728 (a=0.000015) between
40.3-35.9 Ma (161.4-133.25 mbsf), but rose slightly during
this interval, at a rate of about SxlO-6/111^* This rise is
poorly constrained, with a regression coefficient (R2) of
only 0.459. None of the variability within this section is


Figure 2-3: ODP Hole 689B strontium isotope data compared to the model curves, plotted
against age. The time scale used is that of Cande and Kent [1992]. Curving dotted
lines on either side of the regressions are the 95% confidence interval of the
regresssion. Upward curving lines are the 95% confidence intervals for ages
calculated from the 87Sr/86Sr ratio of one sample.


87Sr/86Sr Ratio
Age (Ma)
4*
U1
Age Uncertainty (m.y.)


likely to be of stratigraphic utility, because the total
range is so small.
46
During the late Eocene, a brief increase occurred in
four samples from 35.8-35.6 Ma (132.52-130.25 mbsf), with
the 87Sr/86Sr ratio rising to a constant 0.70778
(a=0.000005) before falling again to about 0.70775 (Figure
2-3). Although the increase is very consistent in this
site, it must be confirmed before any interpretations can be
drawn from it.
A steady rise began in the late Eocene at about 35.5 Ma
(128.8 mbsf, 38.2 Ma BKFV85), with 87Sr/86Sr ratios rising
to near 0.70812 by 25.2 Ma (68.8 mbsf), at a rate of
40xl0-6/m.y. This increase is well constrained, with R2 =
0.972. Two samples above 68.8 mbsf (25.0-24.8 Ma; 68.12 and
67.37 mbsf) show somewhat higher 87Sr/86Sr ratios of about
0.70818, suggesting the possibility of another small hiatus.
Across the hiatus in sedimentation from 24.8 to 20.2 Ma
(66.86 mbsf), 87Sr/86Sr ratios increased sharply to 0.70843
at 20 Ma. Another sudden jump in the ratio occurred across
the unconformity from 19.8 to 18.8 Ma (65.50 mbsf), to an
average of 0.70851 between 18.8 Ma and the top of the
section at 18.5 Ma (62.87 mbsf), confirming the unconformity
identified by Spie/3 [1990] and Gersonde and Burckle [1990].
Oxygen and Carbon Isotopic Measurements
In general, the purpose of the additional 35 isotope
measurements was to improve the resolution of the curve of


47
Kennett and Stott [1990], in particular near the earliest
Oligocene oxygen isotope shift. This was successfully done,
constraining the shift to between 118.44 and 121.30 mbsf
(33.33 to 33.96 Ma CK92, 35.63 to 36.37 Ma BKFV85) (Appendix
A2, Figure 2-2). The shift in this section has a duration
somewhat longer than reported elsewhere [e.g., Kennett and
Shackleton, 1976; Oberhnsli et al., 1984; Zachos et al.,
1992a; Barrera and Huber, 1993; Chapter 3], lasting about
0.6 m.y. During this time, the approximately 1.19/oo to 2.51/oo PDB, with the increase
delineated by 4 separate samples (two of which are
replicated from Kennett and Stott [1990]).
The difference between samples measured in 14
replicates by this study and those measured by Kennett and
Stott [1990] is statistically insignificant. Values for
<5180 averaged 0.07/oo (a=0.29) lower in this study than
Kennett and Stott [1990], while values for 0.03/oo (a=0.18) higher in this study than in Kennett and
Stott [1990]. These differences are insignificant compared
to an analytical error of 0.1/oo for stable isotopic
measurements.
Strontium/Calcium Ratios and Diaoenetic Alteration
Strontium to calcium ratios were measured to assess the
effects of diagenesis in this section. Recrystallization
of calcite results in lower Sr/Ca ratios than in biogenic
calcite because of the lower distribution coefficient of


48
authigenically precipitated calcite present in slowly
altering limestone as compared to rapidly precipitating
biogenic calcite [Lorens, 1981]. Measurement of Sr/Ca
ratios can therefore be used to evaluate the degree of
recrystallization [Hess et al., 1986]. Sr/Ca ratios of 42
samples were measured from ODP Hole 689B (Figure 2-4,
Appendix A2) to evaluate recrystallization in this section.
The Sr/Ca ratios of the samples averaged 3.4 (a=0.6) /g/mg
(mass ratio) or 1.5 (0.3) mmol/mol (molar ratio). The
Sr/Ca ratios rise very slightly up section, from 2.9 (0.3)
jug/mg (mass ratio) or 1.3 (0.1) mmol/mol (molar ratio) in
the bottom 20 m to 3.7 (0.5) /xg/mg (mass ratio) or 1.7
(0.2) mmol/mol (molar ratio) in the top 20 m of section.
These values are comparable or higher than those reported by
other studies [Graham et al., 1982; Hess et al., 1986,1989;
Barrera et al., 1991; Egeberg et al., 1990] Figure 2-4 and I
therefore conclude that the 87Sr/86Sr ratios of the samples
used in this study have not been altered by diagenesis.
The 87Sr/86Sr and Sr/Ca ratios of bulk carbonate
(Egeberg et al. [1990]) correspond well with values measured
on cleaned foraminiferal calcite (Figure 2-4) as well. Bulk
carbonate, with a large component of fine fraction
carbonate, tends to be more susceptible to diagenetic
alteration than coarse fraction carbonate [Garrison, 1981].
The fact that 87Sr/86Sr ratios of bulk carbonate (at least
for the 4 measurements of Egeberg et al. [1990]) and


49
Depth (mbsf)
Figure 2-4: Strontium geochemical measurements,
a: X's and empty boxes are Sr/Ca ratios in mmol/mol and
/xg/mg measured on cleaned foraminiferal samples (this
study), respectively, while solid triangles and boxes are
Sr/Ca ratios in mmol/mol and /g/mg measured on bulk
carbonate from Egeberg et al. [1990].
b: Crosses are 87Sr/86Sr ratios measured on cleaned
foraminiferal samples (this study), while W's and B's are
interstitial water and bulk carbonate 87Sr/86Sr ratios,
respectively, from Egeberg et al. [1990].


50
87Sr/86Sr ratios of foraminiferal calcite are so similar
suggests that little or no diagenetic alteration of these
carbonates has occurred. Lastly, the high 87Sr/86Sr ratios
of interstitial water (Figure 2-4) and relatively low
concentrations of interstitial Sr also argue for low rates
of recrystallization in ODP Hole 689B [Egeberg et al.,
1990]. If recrystallization had occurred, porewater
87Sr/86Sr values would be much lower and bulk carbonate
ratios would be higher. Egeberg et al. [1990] concluded
that pore waters at Site 689 displayed an almost complete
lack of recrystallization.
Discussion
Strontium Isotope Stratigraphy
Strontium isotope stratigraphy is used today in two
primary ways: for determining ages of marine samples based
on their measured 87Sr/86Sr ratio, and for investigating
geochemical cycling of strontium on the Earth's surface
[Hodell, 1994]. The variation in 87Sr/86Sr has been
described using regression analyses, either linear [McKenzie
et al., 1988; Miller et al., 1988, 1991a; Hess et al., 1989;
Hodell et al., 1989, 1990; Oslick et al., 1994] or
curvilinear [Miller et al., 1991a; Hodell and Woodruff,
1994; Oslick et al., 1994]. Oslick et al. [1994] suggested
that in the absence of a theory mandating a polynomial
(curvilinear) fit, the simpler linear regression should be


51
used. The present study uses line segments and linear
regression to model the data because of the relative
simplicity of the method and because linear regression seems
to fit these data well.
I use two linear regressions to model the data. First,
a regression for the Site 689 data from 40.3-35.8 Ma (42.8-
38.6 Ma BKFV85) was calculated. This regression, with 35
data points, is
87Sr/86Sr = 0.708024 7.8588xl06 x Age (Ma) (1)
but gives a rather poor fit (R2 = 0.459) as previously
noted, and a standard error of llxlO-6. The equation can be
inverted for the purposes of age calculation to
Age (Ma) = 90093.124 127245.848 x 87Sr/86Sr (2)
and is valid for values of 87Sr/86Sr from 0.707707 to
0.707743. Second, I calculated a regression for the data
from 35.5-24.8 Ma (38.3-25.4 Ma BKFV85). This regression
equation, with 86 data points, is
87Sr/86Sr = 0.709154 39.6584X10-6 x Age (Ma) (3)
and gives an excellent fit (R2 = 0.972), and a standard
error of 21xl06. The equation can be inverted to


52
Age (Ma) = 17881.554 25215.329 x 87Sr/86Sr. (4)
and is valid for values of 87Sr/86Sr from 0.707746 to
0.708171. The two regressions cross at a calculated age of
35.535 Ma and a 87Sr/86Sr ratio of 0.707745. I do not
evaluate the 4 data points from 35.8-35.6 Ma, which fall off
of both these trends. Figure 2-3 shows the fit of this
model to the data.
In order to evaluate the reliability of this curve in
examining controls on the marine 87Sr/86Sr ratio and in
predicting ages of samples, two statistical analyses of the
regressions were performed. The first analysis calculates
the 95% confidence interval of the regression itself (i.e.,
the uncertainty in the estimate that the regression
correctly models the actual increase in marine 87Sr/86Sr
ratios). The second analysis calculates the uncertainty in
age calculations made from one or more 87Sr/86Sr analyses of
an unknown sample. In the first analysis, I follow the
procedures outlined in Draper and Smith [1981], as suggested
by Miller et al. [1991a] in calculating uncertainties in the
regression. I use the exact solution in calculating
uncertainty in age estimation due to regression uncertainty
rather than the approximate solution used by Miller et al.
[1991a] because the calculations show that the approximate
solution slightly underestimates the uncertainty. In
addition, (2F)^ was used instead of the t statistic. This
allows the 95% confidence curve for the entire regression to


53
be calculated rather than the 95% confidence interval for a
given point in the regression [Draper and Smith, 1981, p.31]
(Figure 2-3). In the second case, the procedure outlined in
Miller et al. [1991a, equation 6] is used to calculate the
uncertainty in age determinations.
A caveat in these calculations is that they are based
on the assumption that 87Sr/86Sr measurements are
distributed normally about each point on the regression. If
there is small-scale structure along the line, as suggested
by several recent studies [Capo and DePaolo, 1990; Dia et
al., 1992; Clemens et al., 1993; Oslick et al., 1994], then
deviations from the regression are systematic. In this
case, calculations of ages and uncertainty using a
regression will be invalid.
The linear regression for the interval from 40.3 to
35.8 Ma shows a rather low regression coefficient
(R2=0.459). This, coupled with the low slope in this
interval, results in a high uncertainty of the regression
and raises the question of whether the slope differs
significantly from zero. A t-test of the regression [Draper
and Smith, 1981] was used to test this hypothesis. The t-
test shows that the slope is significantly different from
zero, at the 99.9% confidence level, despite the low
coefficient. A t-test of the interval from 35.5-24.8 Ma,
with a much greater slope and higher regression coefficient,
not surprisingly also shows that the slope differs
significantly from zero at this confidence level.


54
The uncertainty in the regression from 40.3 to 35.8 Ma
ranges from a minimum of 4.9xl0-6 near the middle of the
interval to 8.6xl06 and 10.9xl06 at the upper and lower
ends of the interval, respectively (Table 2-2). The
corresponding age uncertainties are not quite symmetrically
distributed about the regression line because of the curve
of the lines of uncertainty, but are about 1.37, 0.71, and
1.76 m.y. for the upper, middle, and lower parts of the
interval, respectively (Figure 2-3). These calculations do
not take into account the uncertainty in 87Sr/86Sr
measurements, however. Given the 2a uncertainty in
87Sr/86Sr measurement of 22xl0-6 (comparable to a 95%
confidence interval), and a slope of 7.86xl0-6/m.y., we can
divide the uncertainty by the slope to obtain an estimate of
the measurement uncertainty. We arrive at an uncertainty of
about 2.80 m.y. due to the measurement uncertainty.
Applying the formula given by Miller et al. [1991, eq. 6],
which more rigorously calculates the uncertainty in age
calculations based on a number of 87Sr/86Sr measurements,
the uncertainty for one measurement of an unknown sample at
the 95% confidence interval is approximately 3 m.y., and
decreases by about a third for 3 measurements. Clearly, the
uncertainty in this interval (with a total duration of only
4.5 m.y.) makes this regression useless for precise age
calculations.
For the upper (35.5-24.8 Ma) interval, the
corresponding uncertainties due to the regression in


TABLE 2-2
Uncertainties in regression, due to measurement error,
and due to age calculation from regression
Age Range: 35.5-24.8 Ma 42.8-35.8 Ma
Slope: 39.66xl0-6/my 7.86xl0-6/my
Part of Curve:
Upper
Middle
Lower
Upper
Middle
Lower
Uncertainties:
Regression:
87Sr/86Sr (xlO-6)
10.6
5.7
12.0
8.6
4.9
10.9
Age (m.y., total)
0.53
0.29
0.61
2.74
1.43
3.53
From Measurment
Error (m.v.):
0.55
0.55
0.55
2.80
2.80
2.80
95% Confidence
Interval (m.v.):
1 measurement:
1.08
1.06
1.08
3.05
2.97
3.13
3 measurements:
0.64
0.62
0.65
1.90
1.76
2.01
ui
U1


56
87Sr/86Sr are somewhat larger, at 10.6xl0~6, 5.7xl0-6, and
12.0xl0~6. Due to the higher slope, however, the age
uncertainties due to the regression are smaller, and are
about 0.27, 0.14, and 0.30 Ma for the upper, middle, and
lower parts of the curve, respectively (Table 2-2, Figure 2-
3). The age uncertainty obtained by dividing 2a uncertainty
by the slope is 0.55 m.y. Uncertainty based on 1
measurement is about 1.1 m.y. at the 95% confidence level,
decreasing to about 0.6 m.y. for 3 measurements, suggesting
that regression error is less important than errors in
87Sr/86Sr ratios in calculating ages for this part of the
curve.
Comparisons with Other Records
Three studies have used 87Sr/86Sr records for the
latest Eocene to Oligocene to calculate linear regression of
87Sr/86Sr ratio vs. age [Miller et al., 1988; Hess et al.,
1989; Oslick et al., 1994]. In addition, a fourth study
[Zachos et al., in prep.] has also generated a record from
the middle Eocene up to the earliest Miocene. In comparing
these records (from DSDP and ODP sites) to the Site 689
data, two modifications were made. For records originally
calibrated to the BKFV85 time scale, sample ages were
converted from the BKFV85 time scale to the CK92 time scale
in one of two ways. If magnetostratigraphy of a site was
available, a new age model tied to the CK92 time scale was
used to calculate ages. If magnetostratigraphy of a site


57
was not available, sample ages were calculated by comparison
of the magnetic anomaly tie points used by the BKFV85 and
CK92 age models. The new time scale can significantly
change the pattern observed in the marine 87Sr/86Sr ratios.
For example, the inflection identified by Hess et al. [1989]
at 26 Ma (BKFV85) is even more notable when the Hess et al.
[1989] record is calibrated to the CK92 time scale, where
the inflection is dated at near 24.8 Ma. Coincidentally,
this is near where the continuous portion of the record
ends, and so the slope change is not recognized in the data
from Hole 689B. Second, all 87Sr/86Sr results from these
studies were normalized to the standard value of 0.710235
for NBS-987. For example, Miller et al. [1988] reported a
87Sr/86Sr ratio of 0.710250 for NBS-987; therefore, 0.000015
was subtracted from their data. Hess et al. [1989] used the
Eimer and Amend standard, which converts to a value of
0.710220 for the NBS-987 standard; therefore 0.000015 was
added to all of their data. Oslick et al. [1994] reported a
87Sr/86Sr ratio of 0.710255 for NBS-987; therefore, 0.000020
was subtracted from their data.
In general, the records are quite similar, with the greatest
differences occurring during the early Oligocene and late
Eocene (Figure 2-5). With one exception, the data suggest
that the increase in 87Sr/86Sr ratios began earlier, but
initially at a lower rate, than previous studies. Data from
ODP Hole 748B [Zachos et al., in prep.] also suggests that a
lower rate of increase in 87Sr/86Sr began in the middle or


Figure 2-5: Sr isotope data from ODP Hole 689B (crosses), compared to the Sr curves of
Miller et al. [1988] (Holes 522 and 549A), Hess et al. [1989] (Holes 516, 516F, and
563), Hodell and Woodruff, [1994] (Hole 289, labelled Hodell & W. 1994), Oslick et al.
[1994] (Hole 747B), and Zachos et al. [in prep.] (Hole 748A, labeled Zachos .. in
prep.).


Marine 87Sr/86Sr Ratio
0.7088
0.7087
0.7086-
0.7085-
0.7084
0.7083
0.7082
0.7081
0.7080
0.7079H
0.7078
0.7077H
x
*
*

Miller et al. 1988
Hess et al. 1989
Oslick et al. 1994
Hodell & W. 1994
x
Zachos .. in prep.
This paper
EARLY
LATE
EARLY
MIOCENE
OLIGOCENE
EOCENE
i i i i i i i i i i i i i i i i i i i iiiiiritir
LATE
MIDDLE
16 18 20 22 24 26 28 30 32 34 36 38 40 42 44 46
Age (CK 92, Ma)
VO


60
late Eocene, but the resolution of this record is low
compared to that of Hole 689B. I recalculated the linear
regressions of Miller et al. [1988] and Hess et al. [1989],
and include the regressions of Oslick et al. [1994] for
comparison to this model (Figure 2-6). The regressions of
the different records give quite similar results, with the
exception that the regression of Hess et al. [1989] does
not match the other regressions in the late Eocene to early
Oligocene. This suggests that there may be a problem with
the age model used by Hess et al. [1989]. Alternatively
(see Figure 2-5), there may not be enough points in the data
set to adequately constrain the early part of the records
from sites 516 and 563. It is interesting to note also that
there is a very consistent offset between the data of Miller
et al. [1988] and the data, even after normalization. The
data of Oslick et al. [1994] are offset on average by about
the same amount. The results in the present study are
consistently about 15xl0~6 below those of Miller et al.
[1988] and Oslick et al. [1994]. This offset is in the same
direction, but not quite as large as that observed by Hodell
and Woodruff [1994]. Oslick et al. [1994] attributed the
offset between data from the University of Florida (Site
588; Hodell and Woodruff [1994]) and Rutgers (Hole 747A) to
errors in stratigraphic correlation. However, Holes 689B
and 522 [Miller et al., 1988] are well constrained by
magnetostratigraphy. This does not guarantee that the


Figure 2-6: Regressions from Miller et al. [1988], Hess et al. [1989] (both converted
to CK92), and Oslick et al. [1994] compared with the regression curve from ODP Hole
689B.


Model Marine 87Sr/86Sr Ratio
CT\
K)


63
correlation between these sites is correct, as small
unconformities and mis-identifications can be present which
would not be visible in the magnetostratigraphy, but it
makes it less likely that mis-correlations are the cause of
the offset between the holes. As yet, the reasons for the
offsets between laboratories remain unknown.
Some small scale variability can be seen in the record
from Hole 689B, with local peaks in 87Sr/86Sr values at
35.8-35.6 (as previously mentioned), 31.0, 29.0, and 26.6-
27.5 Ma, and a period of rapid rise from 33.0 to 32.1 Ma
(Figure 2-3). All of these values are within the expected
measurement error, and thus their significance is suspect.
However, they are all represented by more than one
measurement. Very small scale, high frequency variability
in 87Sr/86Sr ratios has been reported in the Plio-
Pleistocene and related to glacial/interglacial changes
[Capo and DePaolo, 1990; Dia et al., 1992; Clemens et al.,
1993], but these variations remain controversial. In
addition, Oslick et al. [1994] identified several episodes
of small scale 87Sr/86Sr variability in the Oligocene and
Miocene, and suggested that they lagged deglaciation events
by approximately 0.9-1.4 m.y. Whether the specific 32 and
28 Ma events identified by Oslick et al. [1994] correlate
with the 31 Ma (or 33 to 32 Ma) and 29 Ma events identified
here is unknown, but the high quality magnetostratigraphy in
the two sections involved (DSDP Hole 522 and ODP Hole 689B)
makes it unlikely. It is possible that finer scale


64
variation does exist in this core, but more detailed studies
will have to be undertaken to confirm this.
Controls on the 87Sr/86Sr Curve
Several possible controls on the strontium isotopic
composition of the ocean have been identified. Most of them
depend on changing inputs or 87Sr/86Sr ratios from
continental erosion (i.e., riverine input) of one type or
another. Some sources of strontium may be dismissed as a
control almost immediately. Dissolution of marine
carbonates only buffers seawater 87Sr/86Sr ratios from
changing, and cannot be invoked as a control [Hodell et al.,
1989]. This also eliminates changes in the CCD as a forcing
mechanism. Input into the ocean from continental erosion
has a 87Sr/86Sr ratio considerably above that of modern
seawater (0.7119 vs 0.7092, Palmer and Edmond, 1992], and it
has been argued that as a result it is much easier to change
the seawater 87Sr/86Sr ratio by altering that source than by
altering hydrothermal input or carbonate dissolution.
Different proposed ways of changing the Sr input from the
continents are: (1) by opening a previously isolated Arctic
Ocean, with inputs from the Canadian and Siberian shields,
(2) by increasing erosion through the effects of the changes
in glaciation [Armstrong, 1971; Hodell et al., 1990; Miller
et al., 1991a; Oslick et al., 1994; Zachos et al., in
prep.], or (3) by increasing erosion via uplift of mountain
ranges such as the Himalayan-Tibetan Plateau [Raymo et al.,


65
1988; Hodell et al., 1989, 1990; Edmond, 1992; Richter et
al., 1992]. Hydrothermal input is usually considered to
change too slowly to be effective [Palmer and Elderfield,
1985; Elderfield, 1986]. However, the difference between
the 87Sr/86Sr ratio of hydrothermal input (0.703) and the
Eocene seawater 87Sr/86Sr ratio (0.7077) is even greater
than between riverine input and seawater 87Sr/86Sr ratio
(0.0047 vs. 0.0042), so hydrothermal input is also
considered as a way of significantly changing the marine
87Sr/86Sr ratio.
To investigate these controls, I used the method of
Hodell et al. [1989], which solves the box model of Palmer
and Elderfield [1985] for non-equilibrium conditions, to
investigate possible changes in fluxes. In calculating
initial conditions, I assumed present day fluxes for
carbonate dissolution and hydrothermal exchange, use the
present day 87Sr/86Sr ratio for hydrothermal exchange and
riverine input. The 87Sr/86Sr ratio for riverine input
(0.7119) should probably be lower in the absence of input
from the Himalayan-Tibetan Plateau [Palmer and Edmond,
1992]. However, by using the higher value, river flux is
allowed to vary more conservatively. I assume that prior to
the increase in marine 87Sr/86Sr in the middle Eocene, the
oceans were in equilibrium at 0.707707. This allows us to
use 0.707707 as the 87Sr/86Sr ratio of carbonate dissolution
(Table 2-3). I can then calculate the equilibrium riverine
flux as a starting point. At equilibrium, the riverine flux


TABLE 2-3
Model Parameters, Stable Latest Eocene Values
Modern Arctic Opening Model Glacial Uplift Model Hydrothermal
Values Arctic non-Arctic Model Model 1 Model 2 Model
Fluxes (109 mol Sr/yr):
Riverine
33.3a
1.00a
14.3289
Hydrothermal
13.202b
0.4388b
12.7 6b
Carbonate
3.4a
0.1130a
3.287a
Arctic Connection (post-
conn.)
20.0e
87Sr/86Sr:
Riverine
0.7119a
0.7128a
0.7119a
Hydrothermal
0.703c
0.703c
0.703c
Carbonate
0.7084a
0.7 098llf
0.707707
Glacial
Seawater
0.709172d
0.709811*?
0.707707
NBS-987
0.710235d
14.8189
14.8189
15.0589
33.3a
13.2b
13.2b
13.2b
29.6599
3.4a
3.4a
3.4a
3.4a
0.7119a
0.7119a
0.7115399
0.7119a
0.703c
0.703c
0.703c
0.703c
0.7077 07b
o^is1
0.707707h
0.7 07707h
0.707707
0.707707h
0.707707h
0.707707h
0.707707
c\


TABLE 2-3continued
a: From Palmer and Edmond (1989), with Arctic and non-Arctic values partitioned if
necessary according to their respective areas (from Knauss, 1978).
b: Calculated from non-equilibrium equations, with Arctic and non-Arctic values
partitioned if necessary according to their respective areas (from Knauss, 1978).
c: From Richter et al. (1992).
d: From Hodell et al. (1990).
e: Approximately 1/100 of Gulf Stream.
f: Defined from Arctic Seawater equilibrium value.
g: Calculated from equlibrium equations.
h: Defined from Hole 689B average for late Eocene.
i: Approximate Gondwana value from Palmer and Edmond (1989).
CTi


68
in the middle Eocene is calculated to be less than 15xl09
moles Sr/yr, about half of today's riverine flux of 33.3xl09
moles Sr/yr.
I model the marine 87Sr/86Sr curve as a series of
straight line segments. This method of modeling can create
artificially abrupt changes in fluxes at changes in slope of
the marine 87Sr/86Sr curve, but is simpler to model and
easier to understand than using a continuously varying
model. Each step in the model is approximated by a portion
of an exponential curve. As a result, it is important to
use a time step which is small relative to the residence
time in order to accurately approximate the model. I use a
step of 0.1 m.y., less than 5% of the residence time of 2.5
Ma for the modern ocean [Hodell et al., 1990] to 3.4-4.0
m.y. for the Eocene and Oligocene (this study, as a result
of smaller riverine fluxes).
Arctic opening model
Plate tectonic reconstructions of the opening of the
Norwegian-Greenland Sea suggest that opening occurred along
a fracture zone, leading to a gateway effect in which the
Arctic was suddenly put into communication with the North
Atlantic Ocean over a short period of time. Prior to 72 Ma
(Cretaceous), the North Atlantic Ocean had not yet begun to
form and the region between Scandinavia and Greenland was
closed with no intervening ocean floor. The first extension
in the Arctic region is believed to have begun at Chron 30


69
time [Rowley and Lottes, 1988] and the oldest seafloor
anomaly to be identified in the North Atlantic and Eurasian
basins is Anomaly 24 [Vogt et al., 1981], at about 53 my
[Cande and Kent, 1992], of early Eocene age. A close
connection is maintained between Greenland and Europe
(Svalbard) during the Eocene as Svalbard was transformed
eastward relative to Greenland along the Greenland fracture
zone. At Anomaly 13 time (33.4 Ma), the contact between
Svalbard and Greenland was tenuous and the final separation
between the European and American plates may have occurred
at this time. This disagrees with interpretations of Lawver
et al. [1990] which suggested a later opening (15-10 Ma).
However, as we shall see below, only a small opening would
be necessary to have a large effect on marine 87Sr/86Sr
ratios. Because the seafloor in the North Atlantic and the
Arctic Ocean would have been at depths of over 4000 m, the
opening may have occurred rapidly, over hundreds or
thousands of years. The motion of the Arctic basin may have
occurred similar to a sliding door, allowing water from the
surface to 4 km depth to be released suddenly into the
Atlantic Ocean.
This sudden influx of cold deep water should produce
severe erosion of abyssal sediments in the North Atlantic
Ocean. Severe erosion is, in fact, observed beginning in
earliest Oligocene time (C13) (e.g., the reflector R4 of
Miller and Tucholke [1983]). If the Arctic had been
isolated up to that point, it would then have mixed with the


70
world ocean, and added its riverine input, hydrothermal
exchange, and carbonate dissolution to those of the World
Ocean.
By using a few assumptions about those sources (Table
2-3), we can calculate what the equilibrium 87Sr/86Sr ratio
of a closed Arctic Ocean would have been (0.709811) and what
riverine input into the world ocean (in the absence of
Arctic input) would be at the equilibrium marine isotopic
value of 0.707707 (14.328xl09 moles Sr/yr, as opposed to
14.818xl09 moles Sr/yr with Arctic input). At the point
when the marine 87Sr/86Sr ratio begins to increase, I model
the connection between the Arctic Ocean and the world ocean
as a small flux between the two oceans, on the order of
1/100 of the flux of Sr carried by the Gulf Stream. I then
use the non-equilibrium model of Hodell et al. [1989] to
model the response of the world ocean to these new inputs
(Figure 2-7a).
Rather than recreating the marine Sr isotope curve, it
is clear that there are fundamental differences between the
modelled and the real curve. First, there is an immediate,
sudden jump in the world ocean 87Sr/86Sr ratio, as the
isotopically more radiogenic Arctic Ocean mixes with the
world oceans, even with the relatively small flux between
the two. Secondly, the modelled marine 87Sr/86Sr ratio
asymptotically approaches equilibrium with the new riverine
input from the Arctic Ocean, rising steeply at first, then
less steeply as time goes on. The modelled Arctic Ocean


Figure 2-7. Modelled response of marine 87Sr/86Sr ratio to different forcing
functions.
a: Opening of the Arctic Ocean, compared to the actual marine 87Sr/86Sr ratio
b: Change in Sr flux from Antarctic glaciation.
c: Change in riverine Sr flux and in both riverine Sr flux and isotopic ratio
d: Change in hydrothermal Sr flux.


87Sr/86Sr Ratio
Age (Ma)
M
87Sr/86Sr Ratio (Arctic)


87Sr/86Sr Ratio
Figure 2-7continued
Glacial Sr Flux (109 mol/yr)


87Sr/86Sr Ratio
Figure 2-7continued
Riverine Sr Flux (109 mol/yr)


87Sr/86Sr Ratio
Figure 2-7continued
Hydrothermal Sr Flux (109 mol/yr)


76
87Sr/86Sr ratio decreases suddenly and then asymptotically
approaches equilibrium with the world ocean. It is clear
that this model cannot adequately explain the nearly linear
increase in marine strontium isotope composition from 35 to
26 Ma. It is also clear that strontium isotopic
measurements of Arctic carbonate sediments may help to
constrain the time of connection of this basin to the world
ocean.
Glaciation model
Recent evidence suggests that glaciation on Antarctica
began near the time that the marine 87Sr/86Sr ratio began to
increase, and that glaciation may in fact be a cause for the
rising 87Sr/86Sr ratios of seawater [Armstrong, 1971; Hodell
et al., 1990; Miller et al., 1991a; Hodell, 1994; Oslick et
al., 1994; Zachos et al., in prep.]. It has been known for
some time that a major increase occurred in 6180 at about
33.4-33.6 Ma (35.7-35.9 of BKFV85). The cause was
originally suggested to be the onset of cold bottom water
production [Shackleton and Kennett, 1975] and more recently,
due to the onset of Antarctic glaciation [Matthews and
Poore, 1980]. A pulse of IRD has been found on the
Kerguelen Plateau [Zachos et al., 1992b; Breza and Wise,
1992; Barrera et al., 1994] and in lesser amounts on the
Maud Rise [Ehrmann and Mackensen, 1992], which coincided
with the earliest Oligocene increase in suggests that glaciation on Antarctica reached the coastline


77
at least briefly at about 33.4-33.6 Ma. At this time,
benthic foraminiferal about 1 /oo, then decreased by about 0.3-0.4 /oo [Miller
et al., 1987; Zachos et al., 1992b; Chapter 3]. This
suggests a sudden advance, followed by a retreat of
continental ice sheets.
Other evidence may point to a somewhat earlier onset of
glaciation. Sedimentological studies of ODP Legs 113 and
119 [Ehrmann and Mackensen, 1992] suggest that glaciers may
have affected Antarctica as early the late middle Eocene,
43.3 Ma (45.5 Ma BKFV85), with a further expansion near the
middle/late Eocene boundary at about 37.1 Ma (40 Ma).
Somewhat controversial evidence of terrigenous grains first
occurred on the Kerguelen Plateau and possibly the Maud Rise
in the middle Eocene from 43.3-39.1 Ma [Wise et al., 1992],
suggesting ice-rafting events during the Eocene. In
addition, massive diamictites drilled in Prydz Bay,
Antarctica, and interpreted as till deposits [Ehrmann et
al., 1992] have been tentatively dated between 37.9 and 34.8
Ma (40.8-37.5 Ma), suggesting a late middle to late Eocene
age for Antarctic glaciation. An increase of chlorite and
kaolinite at the expense of smectite on the Maud Rise in the
late Eocene is suggested to have resulted from increased
physical weathering on Antarctica attributed to increased
Antarctic glaciation [Ehrmann and Mackensen, 1992; Robert
and Chamley, 1992], although a moderate, humid climate is
still postulated for that time. The major sedimentologic


78
evidence cited by Ehrmann and Mackensen [1992] still
suggests the major onset of glaciation in the earliest
Oligocene at about 33.5 Ma. This is the time when the
largest change from smectite to illite occurs in marine clay
mineral assemblages, indicating a change from chemical to
physical weathering conditions on Antarctica [Ehrmann et
al., 1992].
In modeling the effects of glaciation on Antarctica,
equilibrium values for riverine input, hydrothermal
exchange, and carbonate dissolution (Table 2-3) are used.
As a rough average 87Sr/86Sr of typical "Gondwana" shields,
0.715 is used by averaging the values for rivers from
Australia and southern Africa [Palmer and Edmond, 1989]. I
use this value for a new, glacial erosion input (holding
riverine input, hydrothermal exchange, and carbonate
dissolution constant at their equilibrium values; Table 2-3)
and calculate the glacial erosive flux of Sr necessary to
produce the observed record, using the non-equilibrium
method of Hodell et al. [1989] (Figure 2-7b).
In this case we can compare the model's predicted
glacial activity with the observed indications of
glaciation. The comparison is not encouraging. The onset
of modelled glaciation occurred at 40.4 Ma, defined by the
time when the marine 87Sr/86Sr began to increase, and rose
from a low of just 0.24X109 to 0.52xl09 moles Sr/yr by 35.5
Ma (Figure 2-7b). The modelled glacial Sr flux then jumped
to 1.50xl09 moles Sr, about 1/10 of the riverine flux of


79
14.8xl09 moles Sr/yr, reflecting the increased slope of the
87Sr/86Sr curve, and finally rose to a high of 5.19xl09
moles Sr/yr by 24.8 Ma. In contrast, most sedimentologic
and oxygen isotopic evidence points to a somewhat later
onset of glaciation, at about 33.5 Ma, 2 m.y. later than the
first increase in marine 87Sr/86Sr [Rea, 1992 and this
study]. This glaciation, rather than starting at low
intensity and gradually increasing over the next 8 m.y. or
so, as suggested by the strontium isotopic modelling, seems
to have been episodic based on oxygen isotopic evidence,
with maxima at about 33, 29, and 24 Ma (35, 31, 25 Ma;
Miller et al., 1987, 1991a; Wise et al., 1991].
Oslick et al. [1994] also examine the possibility that
glaciation and the 87Sr/86Sr curve are related. They
identified 6 events in the Oligocene and Miocene in which
very small scale 87Sr/86Sr (from ODP Hole 747A) increases
lag deglacial events by 0.9 to 1.4 m.y. The deglaciation
events were defined by decreases in followed previously identified glacioeustatic episodes
[Miller et al., 1991b, Wright and Miller, 1992]. They
attribute the overall rise in marine 87Sr/86Sr ratios in the
Oligocene and the shorter scale 87Sr/86Sr increases in the
Miocene to glaciation, with the overall rise in the Miocene
due to uplift of the Himalyan-Tibetan Plateau. The 0.9 to
1.4 m.y. lag in response is attributed to the long residence
time of the oceanic Sr reservoir.


80
I disagree with this interpretation on three points.
First, the long residence time of Sr will create a lag
between the initial 87Sr/86Sr response and its final
equilibrium value, not between the driving mechanism
(weathering of the Antarctic craton during deglaciation
[Oslick et al., 1994]) and initial response of marine
87Sr/86Sr values. If in fact the lagged association between
deglaciations and 87Sr/86Sr increases is real, another
mechanism must be invoked to delay the release of radiogenic
Sr to the ocean until 0.9 to 1.4 m.y. after the
deglaciation. Second, as stated above, the continued rise
in marine 87Sr/86Sr ratios implies a gradually increasing
level of glacial weathering throughout the Oligocene (Figure
2-7b). Evidence for glaciation in the Oligocene instead
suggests a more episodic behavior. Unless the amplitude of
these episodic glaciations slowly increased (which is not
implied by Antarctic craton gradually increased during each deglacial
episode (perhaps through dissolution of gradually increasing
deposits of finely ground, easily weathered cratonic till
[Hodell et al., 1990, B. Opdyke, pers. comm., 1994]), it
would not be possible to produce the observed 87Sr/86Sr
curve from Oligocene glaciation. Third, the inflection to a
high slope of the 87Sr/86Sr curve clearly begins before the
first glaciation event cited by Oslick et al. [1994] (about
35.5 Ma vs. about 33.5 Ma). The time of onset of the slower
rise in 87Sr/86Sr began even earlier, around 40 Ma.


81
Even if interpretations of sedimentologic evidence
suggesting earlier onset of glaciation are correct, it may
be hard to ascribe the increase in marine 87Sr/86Sr to
glaciation. Two factors support this conclusion. First,
the 87Sr/86Sr ratio in river water draining continental
shields under most circumstances is likely to be lower than
the bulk ratio of the shield rocks, because the minerals
containing highly radiogenic 87Sr/86Sr ratios are more
resistant to erosion [Palmer and Edmond, 1992; Edmond,
1992]. Higher (more radiogenic) 87Sr/86Sr ratios are
present in minerals with high Rb contents (i.e., containing
incompatible monovalent cations such as potassium feldspar,
sodic plagioclase, and muscovite) which are also fairly
resistant to weathering. Biotite has been suggested as one
mineral containing highly radiogenic Sr which is easily
weathered in glacial deposits, but the amount of Sr
contained in biotite is low [Blum and Erel, 1995]. Less
radiogenic 87Sr/86Sr ratios are generally present in more
easily weathered minerals containing divalent cations such
as calcic plagioclase and olivine. Strontium isotopic
ratios derived from shield areas are therefore usually
either not high enough, or with strontium fluxes too low, to
control the marine 87Sr/86Sr ratio [Edmond, 1992], For
example, the marine 87Sr/86Sr ratio of rivers draining the
Canadian Shield averages only 0.7111, with more radiogenic
ratios present in rivers with lower Sr concentrations
[Wadleigh et al., 1985], Special circumstances must be


82
invoked, such as deep-seated metamorphic homogenization, for
rivers with high 87Sr/86Sr to also have large Sr fluxes
[Edmond, 1992]. Second, although the increase in illite in
the late Eocene suggests climatic deterioration possibly
associated with glaciation, it is suggestive of glaciation
precisely because it indicates a change in the mode of
weathering, from chemical to physical. This, however, may
not necessarily reduce the amount of Sr weathered from the
craton, because mechanical grinding by glaciers may lead to
enhanced chemical dissolution [Armstrong, 1971; Hodell et
al., 1990]. There is evidence that chemical denudation
rates of cations in glacial environments are high [Reynolds
and Johnson, 1972], and it is likely that the temporary
glacial events suggested for the Oligocene were temperate
(wet based) in nature.
Uplift model
The Himalayan-Tibetan Plateau system is unique in a
number of ways. First, it is the highest mountain range on
Earth, with an average elevation of 5000 m [Raymo and
Ruddiman, 1992]. Second, rivers draining it have high
87Sr/86Sr ratios for the flux of Sr that they carry [Palmer
and Edmond, 1989, 1992]. It has been suggested that the
anomalous co-occurrence of high 87Sr/86Sr ratios with high
Sr fluxes in Himalayan rivers is due to unique tectonic
aspects of the region [Palmer and Edmond, 1992; Edmond,
1992]. Uplift of the Himalayan orogen followed unusually


83
deep seated metamorphism due to the collision of India with
Eurasia and subsequent crustal thickening, allowing
radiogenic Sr (formerly trapped within hard-to-weather Na-
and K-bearing silicates) to be remobilized into easier-to-
weather Ca-silicates [Edmond, 1992]. Uplift exposes more,
and deeper, continental crust to weathering, and because of
mass wasting in areas of high relief, fresh rock is
repeatedly exposed to continued weathering. As a result,
several studies [Raymo et al, 1988; Hodell et al., 1989,
1990; Raymo and Ruddiman, 1992; Richter et al., 1992; Hodell
and Woodruff, 1994] have suggested that uplift of the
Himalayas controls the marine 87Sr/86Sr ratio during the
Cenozoic by increasing the amount of continental weathering
and delivery of radiogenic Sr to the world's oceans.
Raymo et al. [1988], Hodell et al. [1989, 1990], and
Raymo and Ruddiman [1992] suggested Himalayan control of
marine Sr isotopic ratios during the Neogene. Richter et
al. [1992] suggested that this control extended back to the
Paleogene, with the first effects of the uplift on marine
87Sr/86Sr ratios occurring at about 37 Ma (40 Ma BKFV85,
their date for the first increase in the Sr isotopic ratio).
They modelled the increase using a non-equilibrium model in
which riverine input of Sr controls the marine 87Sr/86Sr
curve. I re-calculated the curve for this time period for a
number of reasons. First, the model is based on the data
from Hole 689B, which is more highly refined than previous
records. Second, the time-step used by Richter et al.


84
[1992], 1 m.y., was too small to accurately portray the
large changes in flux implied by the data. An example of
this is the change in 87Sr/86Sr ratios that occurred in the
late Eocene. The sudden change in slope of the marine
87Sr/86Sr curve implies that large and rapid changes in
marine 87Sr/86Sr ratios occurred at this time. By using a 1
m.y. time step, they underestimated these changes. Third,
the model used by Richter et al. [1992] ignores the effect
of carbonate dissolution in calculating fluxes. Although
small, this flux is significant, and leaving out its
buffering effect [Raymo and Ruddiman, 1992] on the Sr
isotope budget gives underestimates of the fluxes necessary
to change the marine 87Sr/86Sr curve. Fourth, I re
calculate the curve based on the new Cande and Kent [1992]
time scale. These calculations are based on rates-of-
change, which are sensitive to changes in time scale
[Delaney and Boyle, 1988].
I model the change in two ways. Richter et al., [1992]
noted that in increasing erosion from the Himalayan-Tibetan
Plateau, both riverine Sr flux and 87Sr/86Sr may change. I
first assume that world riverine 87Sr/86Sr stays constant at
the current value, 0.7119 [Palmer and Edmond, 1989], The
result is that a small jump in flux occurred, from the
equilibrium value of 14.82xl09 to 15.06xl09 moles Sr/yr at
the beginning of the increase in 87Sr/86Sr at 40.4 Ma. The
riverine flux gradually rose to 15.34xl09 moles Sr/yr by
35.5 Ma. At that time, a sudden increase in riverine


85
strontium flux occurred, from 15.34xl09 to 16.32xl09 moles
Sr/yr, an increase of about 6.4%. The sudden jump is
probably an artifact resulting from the modelling of the
data as having a sudden change in slope, from nearly stable
87Sr/86Sr values to steadily increasing values. If instead
of suddenly changing, the marine 87Sr/86Sr ratio ramps
gradually up to increasing values, the increase in riverine
flux would be more gradual. The model flux grows steadily
from 35.5 Ma to a maximum of 20xl09 moles Sr/yr at 24.8 Ma
(Figure 2-7c, Table 2-3). For the second model, I assume
that the riverine Sr flux began to increase slowly, as the
Himalayas were uplifted, but that the 87Sr/86Sr ratio of the
rivers draining the Himalayas was higher, such that the Sr
flux increased from an equilibrium value of 15.06xl09 moles
Sr/yr. In this case, it is a simple equilibrium calculation
to determine what the riverine 87Sr/86Sr ratio would have
been before the increase. Using equilibrium calculations,
if the Eocene riverine Sr flux was 15.06xl09 moles Sr/yr,
the riverine 87Sr/86Sr ratio must have been 0.711833. These
results are similar to the results of Richter et al. [1992],
except that (1) the calculated fluxes are higher, and (2)
the timing of course is different, in part because of the
different time scales, and in part because of the different
87Sr/86Sr curves employed.
Superficially, this seems to be a reasonable simulation
of the events leading to the change in marine 87Sr/86Sr from
the Eocene through the Oligocene. The next step is to


Full Text
I certify that I have read this study and that in my
opinion it conforms to acceptable standards of scholarly
presentation and is fully adequate^ in scope and quality, as
a dissertation for the degree o^/Dpctor of Philosophy.
David A. Hodell, Chair
Associate Professor of Geology
I certify that I have read this study and that in my
opinion it conforms to acceptable standards of scholarly
presentation and is fully adequate, in scope and quality, as
a dissertation for the degree of Doctor of Philosophy.
11*-!'
/>
/
J- 'f- {
Paul F. Ciesielski
Associate Professor of Geology
I certify that I have read this study and that in my
opinion it conforms to acceptable standards of scholarly
presentation and is fully adequate, in scope and quality, as
a dissertation for the degree of Doctor of Philosophy.
Dougl
Professor
I certify that I have read this study and that in my
opinion it conforms to acceptable standards of scholarly
presentation and is fully adequate, in scope and quality, as
a dissertation for the degree of Doctor, of Philosop
S
Opdyke /
Neil D. f
Professor of Geology
I certify that I have read this study and that in my
opinion it conforms to acceptable standards of scholarly
presentation and is fully adequate, in scope and quality, as
a dissertation for the degree of Doctor of Philosophy.
4
; r
Claire L. Schelske
Carl S. Swisher Eminent
Scholar in Water Resources
This dissertation was submitted to the Graduate Faculty
of the Department of Geology in the College of Liberal Arts
and Sciences and to the Graduate School and was accepted as
partial fulfillment of the requirements for the degree of
Doctor of Philosophy.
May, 1995
Dean, Graduate School


1780
199^
.mvi
UNIVERSITY OF FLORIDA



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