Citation
Magnetic Stratigraphy and Environmental Magnetism of Oceanic Sediments

Material Information

Title:
Magnetic Stratigraphy and Environmental Magnetism of Oceanic Sediments
Creator:
EVANS, HELEN F. ( Author, Primary )
Copyright Date:
2008

Subjects

Subjects / Keywords:
Axial tilt ( jstor )
Demagnetization ( jstor )
Grain size ( jstor )
Magnetic recording ( jstor )
Magnetism ( jstor )
Oceans ( jstor )
Oxygen isotopes ( jstor )
Reflectance ( jstor )
Sediments ( jstor )
Stratigraphy ( jstor )

Record Information

Source Institution:
University of Florida
Holding Location:
University of Florida
Rights Management:
Copyright Helen F. Evans. Permission granted to University of Florida to digitize and display this item for non-profit research and educational purposes. Any reuse of this item in excess of fair use or other copyright exemptions requires permission of the copyright holder.
Embargo Date:
3/1/2007
Resource Identifier:
659561078 ( OCLC )

Downloads

This item has the following downloads:

evans_h ( .pdf )

evans_h_Page_065.txt

evans_h_Page_011.txt

evans_h_Page_019.txt

evans_h_Page_051.txt

evans_h_Page_055.txt

evans_h_Page_012.txt

evans_h_Page_037.txt

evans_h_Page_158.txt

evans_h_Page_164.txt

evans_h_Page_110.txt

evans_h_Page_093.txt

evans_h_Page_078.txt

evans_h_Page_034.txt

evans_h_Page_028.txt

evans_h_Page_189.txt

evans_h_Page_168.txt

evans_h_Page_081.txt

evans_h_Page_174.txt

evans_h_Page_003.txt

evans_h_Page_173.txt

evans_h_Page_030.txt

evans_h_Page_015.txt

evans_h_Page_139.txt

evans_h_Page_202.txt

evans_h_Page_182.txt

evans_h_Page_129.txt

evans_h_Page_199.txt

evans_h_Page_080.txt

evans_h_Page_009.txt

evans_h_Page_136.txt

evans_h_Page_138.txt

evans_h_Page_140.txt

evans_h_Page_188.txt

evans_h_Page_113.txt

evans_h_Page_170.txt

evans_h_Page_153.txt

evans_h_Page_085.txt

evans_h_Page_008.txt

evans_h_Page_029.txt

evans_h_Page_014.txt

evans_h_Page_104.txt

evans_h_Page_097.txt

evans_h_Page_123.txt

evans_h_Page_177.txt

evans_h_Page_044.txt

evans_h_Page_149.txt

evans_h_Page_031.txt

evans_h_Page_133.txt

evans_h_Page_038.txt

evans_h_Page_111.txt

evans_h_Page_099.txt

evans_h_Page_183.txt

evans_h_Page_120.txt

evans_h_Page_047.txt

evans_h_Page_198.txt

evans_h_Page_042.txt

evans_h_Page_194.txt

evans_h_Page_007.txt

evans_h_Page_190.txt

evans_h_Page_083.txt

evans_h_Page_137.txt

evans_h_Page_163.txt

evans_h_Page_197.txt

evans_h_Page_101.txt

evans_h_Page_125.txt

evans_h_Page_061.txt

evans_h_Page_058.txt

evans_h_Page_002.txt

evans_h_Page_020.txt

evans_h_Page_195.txt

evans_h_Page_150.txt

evans_h_Page_062.txt

evans_h_Page_079.txt

evans_h_Page_075.txt

evans_h_Page_155.txt

evans_h_Page_171.txt

evans_h_Page_132.txt

evans_h_Page_092.txt

evans_h_Page_090.txt

evans_h_Page_180.txt

evans_h_Page_201.txt

evans_h_Page_006.txt

evans_h_Page_082.txt

evans_h_Page_050.txt

evans_h_Page_115.txt

evans_h_Page_105.txt

evans_h_Page_041.txt

evans_h_Page_130.txt

evans_h_Page_039.txt

evans_h_Page_073.txt

evans_h_pdf.txt

evans_h_Page_159.txt

evans_h_Page_069.txt

evans_h_Page_074.txt

evans_h_Page_176.txt

evans_h_Page_186.txt

evans_h_Page_154.txt

evans_h_Page_128.txt

evans_h_Page_017.txt

evans_h_Page_134.txt

evans_h_Page_167.txt

evans_h_Page_160.txt

evans_h_Page_116.txt

evans_h_Page_059.txt

evans_h_Page_004.txt

evans_h_Page_021.txt

evans_h_Page_200.txt

evans_h_Page_148.txt

evans_h_Page_193.txt

evans_h_Page_001.txt

evans_h_Page_067.txt

evans_h_Page_151.txt

evans_h_Page_109.txt

evans_h_Page_091.txt

evans_h_Page_066.txt

evans_h_Page_094.txt

evans_h_Page_095.txt

evans_h_Page_071.txt

evans_h_Page_143.txt

evans_h_Page_114.txt

evans_h_Page_121.txt

evans_h_Page_022.txt

evans_h_Page_107.txt

evans_h_Page_064.txt

evans_h_Page_057.txt

evans_h_Page_077.txt

evans_h_Page_010.txt

evans_h_Page_016.txt

evans_h_Page_027.txt

evans_h_Page_161.txt

evans_h_Page_048.txt

evans_h_Page_122.txt

evans_h_Page_156.txt

evans_h_Page_084.txt

evans_h_Page_053.txt

evans_h_Page_063.txt

evans_h_Page_126.txt

evans_h_Page_144.txt

evans_h_Page_025.txt

evans_h_Page_096.txt

evans_h_Page_045.txt

evans_h_Page_086.txt

evans_h_Page_165.txt

evans_h_Page_169.txt

evans_h_Page_103.txt

evans_h_Page_175.txt

evans_h_Page_076.txt

evans_h_Page_098.txt

evans_h_Page_135.txt

evans_h_Page_102.txt

evans_h_Page_157.txt

evans_h_Page_118.txt

evans_h_Page_124.txt

evans_h_Page_106.txt

evans_h_Page_023.txt

evans_h_Page_203.txt

evans_h_Page_013.txt

evans_h_Page_185.txt

evans_h_Page_036.txt

evans_h_Page_026.txt

evans_h_Page_196.txt

evans_h_Page_172.txt

evans_h_Page_060.txt

evans_h_Page_131.txt

evans_h_Page_178.txt

evans_h_Page_108.txt

evans_h_Page_145.txt

evans_h_Page_088.txt

evans_h_Page_184.txt

evans_h_Page_052.txt

evans_h_Page_191.txt

evans_h_Page_068.txt

evans_h_Page_035.txt

evans_h_Page_204.txt

evans_h_Page_179.txt

evans_h_Page_162.txt

evans_h_Page_181.txt

evans_h_Page_089.txt

evans_h_Page_152.txt

evans_h_Page_070.txt

evans_h_Page_100.txt

evans_h_Page_032.txt

evans_h_Page_024.txt

evans_h_Page_117.txt

evans_h_Page_112.txt

evans_h_Page_142.txt

evans_h_Page_166.txt

evans_h_Page_043.txt

evans_h_Page_141.txt

evans_h_Page_040.txt

evans_h_Page_187.txt

evans_h_Page_087.txt

evans_h_Page_018.txt

evans_h_Page_046.txt

evans_h_Page_033.txt

evans_h_Page_005.txt

evans_h_Page_192.txt

evans_h_Page_049.txt

evans_h_Page_054.txt

evans_h_Page_119.txt

evans_h_Page_127.txt

evans_h_Page_147.txt

evans_h_Page_056.txt

evans_h_Page_146.txt

evans_h_Page_072.txt


Full Text





MAGNETIC STRATIGRAPHY AND ENVIRONMENTAL MAGNETISM OF OCEANIC
SEDIMENTS



















By

HELEN F. EVANS


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

2006

































Copyright 2006

by

Helen F. Evans


































For Jane and Eryl









ACKNOWLEDGMENTS

I would like to thank my advisor, Jim Channell, for giving me the opportunity to study in

Florida and for all his help and support over the last 7 years. I also thank my committee members

Ellen Martin, Neil Opdyke, John Jaeger and Bo Gustafson for agreeing to supervise my research

over the last five years and for all their help, academic and otherwise. I also had the pleasure of

collaborating and interacting with a number of talented individuals without whose help I would

not have been able to accomplish this work: Gary Acton, Paul Bown, Yohan Guyodo, Sean

Higgins, Claude Hillaire-Marcel, Dave Hodell, Kainian Huang, Mark Leckie, Ulla Rohl, Joseph

Stoner, Ray Thomas, Thomas Westerhold and many others. My research was made possible by

the technical and financial support of several organizations including the National Science

Foundation, the JOIDES U.S. Sciences Support Advisory Committee (USSAC), and the

Graduate Student Council. Support was also provided by the Institute for Rock Magnetism, the

College of Liberal Arts and Sciences, the Graduate School, the McLaughlin family, and the

department of Geological Sciences at the University of Florida.

I thank all those who gave me their moral support during my years of study, which often

consisted of many hours of festivities in the numerous bars and restaurants in Gainesville, San

Francisco and further afield. I thank in particular Gillian Rosen, Joe Graves, Joann and Jason

Hochstein, Howie Scher, George and Katherin Kamenov, Cara Gentry, Jen Mays, Steve Volpe,

Phil D'Amo, Bricky Way, Kendall Fountain, Adi Gilli, Simon Nielsen, Dave Hodell, Mike

Rosenmeier, William Kenney, Yohan Guyodo and Victoria Meija. Finally I would like to

express my gratitude to my family without whose support I would not have been able to

complete this work. I thank my late mother Eryl, my father Terry and my brother Michael. I also

thank Diane and John Thomas, Judy and Robin Ganz, and Molly and Reg Beynon.










TABLE OF CONTENTS

page

ACKNOWLEDGMENT S............................................................ ....... .............. 4

LIST OF TABLES ..................................................... 7

LIST OF FIGURES .............. ........................................................

AB STRACT ................................................. ....... ....... ........ .............. 12

CHAPTER

1 IN TR O D U C TIO N ..................................................... 14

2 LATE MIOCENE-HOLOCENE MAGNETIC POLARITY STRATIGRAPHY AND
ASTROCHRONOLOGY FROM ODP LEG 198-SHATSKY RISE ..... ..................... 18

Intro du action ..................................................... 18
M eth o d s ....................................................................... 19
M agnetostratigraphic Interpretation ............................................................................. ......... 20
A stro ch ro n o lo g y .................................................................. ...................................... 2 3
D iscu ssio n ..................................................... 2 4
C on clu sion s ..................................................... 2 7

3 INTEGRATED NEOGENE MAGNETIC, CYCLE AND BIO- STRATIGRAPHY
FROM ODP SITE 1208 (SHATSKY RISE, PACIFIC OCEAN) ........ ..............48

Introduction ......... ........ ........ ................. ...... ....... ...... ......... 48
Site Location and Lithology .............. ............ .. .................................. 51
M magnetic Stratigraphy ............................................................................................... 51
C y cle S tratig rap h y ................................................................ .................................... 5 3
C alcareous N annofossils ................................................................................................ 54
Planktonic Foraminifera .................................................................. .... ........ 56
C o n c lu sio n s .......................................................................................... 5 8

4 PALEOINTENSITY-ASSISTED CHRONOSTRATIGRAPHY OF DETRITAL
LAYERS ON THE EIRIK DRIFT (NORTH ATLANTIC) SINCE MARINE ISOTOPE
S T A G E 1 1 .............. .... ............. ................. ........................................................7 5

In tro d u ctio n .............. .... ............. ................. ......................................... 7 5
M methods ..................................... ..... .......... ........... .............. 76
NRM and Normalized Remanence Record ................................................................78
P olarity E excursions ................................................................................................ 79
Relative Paleointensity ................................................................. ... ........ 79
C h ro n o lo g y ........................................................................................................... 8 0
Detrital Layer Stratigraphy.. .............................................................. ............... 81









D iscu ssio n ........................ ................. ................4
C o n clu sio n s ........................ ................. ................7

5 RELATIVE PALEOINTENSITY STACK FOR THE LAST 85 KYR ON A REVISED
GISP CHRONOLOGY, AND ENVIRONMENTAL MAGNETISM OF THE
G A R D A R D R IF T .............. .... ............. ................. ........................................... 10 5

Introdu action ......... ..... ............. ................................. ........................... 10 5
S ite L o catio n s .............. .... ............. ................. ........................ ................. 10 7
M methods .............. ............................................ ................. ......... 108
Directional Magnetic Data ................. ........ ........... ........ 110
Normalized Remanence ................ ............... ...... ............. 111
Stable Isotope Data and Age Models.......................................... 112
Bulk Magnetic and Physical Parameters .................. ............. .......... 112
R elative Paleointensity Stack ..................................................................... 113
Environm mental M agnetism ............................................................................................ 115
C o n c lu sio n s ......................................................................................... 1 1 8

6 RELATIVE GEOMAGNETIC PALEOINTENSITY IN THE GAUSS AND GILBERT
CHRONS FROM IODP SITE U1313 (NORTH ATLANTIC) .............. ............ 136

In tro d u ctio n ......................................................................................... 13 6
M eth o d s ................... ................... ............................3 8
R e su lts ................... ................... ............................3 9
D isc u ssio n .......................................................................................... 14 2

7 ODP SITE 1092 REVISED COMPOSITE DEPTH SECTION HAS IMPLICATIONS
FOR UPPER MIOCENE "CRYPTOCHRONS".................................. 161

Introduction .......................................................... ..................... ........ 161
Revised Composite Depths (rmcd) .............................................................. ............. 162
Implications for Magnetic Stratigraphy ...... ........ ............. .......... ......... 163

8 ASTRONOMICAL AGES FOR MIOCENE POLARITY CHRONS C4AR-C5R (9.3-
11.2 MA), AND FOR THREE EXCURSION CHRONS WITHIN C5N.2N..................... 171

Introduction ............ ......... ......... .................................... .............................. 17 1
Methods and Results .......... ......... ......... ... ............. 173
Comparison with Other Timescales...................................... 175
Excursion Chrons............................ .................. 178

9 CONCLUSIONS AND FUTURE WORK ............................ ................... 189

L IST O F R E F E R E N C E S .................... ......... .................................................... ............. 19 1

B IO G R A PH IC A L SK E T C H ....................................................... ....................................... 204









LIST OF TABLES
Table page

2-1 Latitude, longitude, w ater depth ...................... .... .......... .................... .............. 28

2-2 M agnetostratigraphic age m odel ........................................................ .............. 29

2-3 Comparison of astrochronological age models............... ................................... 30

2-4 A strochronological ages for Leg 198 .................................... .......................... ......... 31

3-1 Depths of reversal boundaries from ODP Site 1208...................................................... 59

3-2 Astronomically calibrated ages for reversal boundaries from ODP Site 1208............... 60

3-3 Nannofossil datum s for ODP Site 1208 ................................................. ....... ...... 61

3-4 Plio-Pleistocene foraminfer datum s ................................................................... 62

3-5 M iocene foram inifer datum s.......................................................... .......................... 63

4-1 Core, latitude, longitude, water depth and base age of the core. ................................ 89

4-2 DC and LDC layer properties in Core MD99-2227 ............. ..... ...................90

4-3 Detrital Layers from other studies considered to be correlative to detrital layers
identified on Eirik drift.............. ... ................ ......... ............... .... .......... 91

5-1 Summary of the cores used in this study and the eleven cores used in the relative
paleointensity stack ................................................................. ... ......... 120

6-1 Depth of polarity chrons from IODP Site U1313 ............................................ 146

6-2 Polarity reversal ages determined at Site U1313 ................ ............................ 147

7-1 Adjusted depths of core tops from ODP site 1092 ..................................................... 166

7-2 Position of the polarity zone boundaries at site 1092........................................... 167

8-1 Astronomical ages from recent timescales compared with those inferred at ODP Site
1 0 9 2 ................... ................... ................... ................................. .. 1 8 1









LIST OF FIGURES


Figure page

2-1 B athym etric m ap of Shatsky R ise........................................................ .. .................. 32

2-2 Representative orthogonal projections of AF demagnetization data. ........................... 33

2-3 Site 1207 component inclination and declination from discrete samples for 0-80
m eters ...............................................................................................3 4

2-4 Site 1207 component inclination and declination from discrete samples for 80-160
m ete rs ...................................... ........................................ ............... 3 5

2-5 Interval sedimentation rates and age versus depth................................................. 36

2-6 Site 1209 component inclination and declination from discrete samples ....................... 37

2-7 Site 1210 component inclination and declination from discrete samples ........................38

2-8 Site 1211 component inclination and declination from discrete samples ........................39

2-9 Site 1212 component inclination and declination from discrete samples ........................40

2-10 Pow er spectra ..................................................................... ......... ........ 41

2-11 The astronomical solution for obliquity compared with tuned L* reflectance data
from Site 1207 ......... ......... .......................................... ........................... 42

2-12 The astronomical solution for obliquity compared with tuned L* reflectance data
from Site 1208 ......... ......... .......................................... ........................... 43

2-13 The astronomical solution for obliquity compared with tuned L* reflectance data
from Site 1209 ......... ......... .......................................... ........................... 44

2-14 The astronomical solution for obliquity compared with tuned L* reflectance data
from Site 1210 ................. ............................................. ......... 45

2-15 The astronomical solution for obliquity compared with tuned L* reflectance data
from Site 12 11 ......... .... ..... ......... ................................ .......................... 46

2-16 C ross-spectral analy sis ...................................... ............ ............... ........................... 47

3-1 Bathymetric map showing the location of Shatsky Rise in the Pacific Ocean ............... 64

3-2 Inclination, declination and MAD values plotted against meters below sea floor .......... 65

3-3 Inclination, declination and M AD values............................................. .. .................. 66









3-4 Inclination, declination and MAD values ..................... ............... ................ 67

3-5 Orthogonal projections showing AF demagnetization data .......................... ......... 68

3-6 Interval sedimentation rates ............................ ...... ......... 69

3-7 R eflectance (L *) data .......... ................ ....................................... ............................ 70

3-8 Plio-Pleistocene planktonic foraminifer and calcareous nannofossil datums................... 71

3-9 Miocene planktonic foraminifer and calcareous nannofossil datums.............................72

3-10 Calcareous nannofossil biostratigraphy...................................................................... 73

3-11 A proposed biostratigraphy for the mid-latitude North Pacific .............. ................... 74

4-1 Location m ap showing the Labrador Sea ................................... ................................. 92

4-2 Component inclination, corrected component declination and maximum angular
dev nation ............................................................................................ 93

4-3 Component inclination, declination and maximum angular deviation (MAD) values
recording Laschamp and Iceland Basin polarity excursions .........................................94

4-4 Anhysteretic susceptibility (kam) plotted against volume susceptibility (k) ..................95

4-5 NRM, ARM, IRM and volume susceptibility..................... ......................... 96

4-6 JPC 19: Relative paleointensity record correlated to that from ODP Site 983................ 97

4-7 JPC 18: Relative paleointensity data correlated to ODP Site 983 ............................98

4-8 JPC 15: Relative paleointensity data correlated to ODP Site 983 ............. ................99

4-9 MD99-2227: Relative paleointensity data correlated to ODP Site 983 ......................... 100

4-10 kam/k and m agnetic susceptibility versus age..................................... ................... 101

4-11 Core MD99-2227: kam/k, magnetic susceptibility, bulk (GRAPE) density ................ 102

4-12 Photographs and X-radiographs of three detrital......................................................... 103

4-13 Hysteresis ratios Mr/Ms plotted versus Her/Hc .......... ................. ................... 104

5-1 Location map for cores analyzed in this study............................ 121

5-2 Correlation of the magnetic susceptibility records .............................................. 122

5-3 Orthogonal projections of alternating field demagnetization data ........................... 123









5-4 Component inclination, declination and maximum angular deviation (MAD) values ... 124

5-5 Plot of anhysteretic susceptibility (karm) versus volume susceptibility (k)..................... 125

5-6 Paleointensity proxies.................... ................................................ ........... ........ ..... 126

5-7 Core JPC13 benthic oxygen isotope record ........................................... ............. 127

5-8 Relative paleointensity records from Cores JPC2, JPC5 correlated to Core JPC13....... 128

5-9 Interval sedimentation rates for Cores JPC2, JPC5 and JPC13 .................................. 129

5-10 Core JPC13: GRA bulk density ............................................... 130

5-11 Anhysteretic susceptibility divided by volume magnetic susceptibility ........................ 131

5-12 Eleven relative paleointensity records from the North Atlantic Ocean........................ 132

5-13 The new relative paleointensity stack ................................... ............................ ........ 133

5-14 Comparison of the EHC06 paleointensity stack to 36C1 flux............... ............. 134

5-15 Comparison of the EHC06 paleointensity stack .......................................................... 135

6-1 Location map for IODP Site U 1313............................................................... 148

6-2 Magnetic polarity stratigraphy from IODP Site U1313 in the 120-200 mcd interval..... 149

6-3 Magnetic polarity stratigraphy from IODP Site U1313 in the 200-280 mcd interval .... 150

6-4 Vector end-point projections of AF demagnetization data........................................... 151

6-5 Interval sedim entation rates ....................................................................................... 152

6-6 G auss C hronozone at Site U 1313 ............. ............. ................................. ............. 153

6-7 The magnetic grain size proxy, anhysteretic susceptibility divided by susceptibility .... 154

6-8 Later part of the Gilbert Chronozone at Site U 1313 .................................................... 155

6-9 Relative paleointensity records from IODP Site U1313 .............................................. 156

6-10 Volume magnetic susceptibility from u-channel samples................................. 157

6-11 Volume magnetic susceptibility from u-channel samples and L* reflectance data
m measured shipboard ................................................................. .... ........ 158

6-12 M ean volum e m agnetic susceptibility ................................... .................................... 159

6-13 Output of a gaussian filter centered on a period of 41 kyr ............................................ 160









7-1 Fe intensity (XRF) data plotted as a five-point moving average.................................. 168

7-2 Inclination of the characteristic magnetization component .................. ................... 169

7-3 Site 1092 ............ .... ............................................................................... .. 170

8-1 Magnetic component inclination for the C4Ar. n-C5r. n interval............. ............. 182

8-2 Oxygen isotope records from the C4An-C5r.ln interval at ODP Site 1092 ................ 183

8-3 Power spectrum generated from the oxygen isotope stack in the depth domain............ 184

8-4 Upper plot shows the correlation of the filtered (filter centered at 0.0244 + 0.0073
kyr1) oxygen isotope stack to the astronomical solution for obliquity......................... 185

8-5 Interval sedimentation rates for the C4Ar. ln-C5r. In interval ............. ................ 186

8-6 Comparison of the age estimates of polarity chrons at ODP Site 1092 .................... 187

8-7 The Site 1092 relative paleointensity record for C5n.2n......................................... 188









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

MAGNETIC STRATIGRAPHY AND ENVIRONMENTAL MAGNETISM OF OCEANIC
SEDIMENTS

By

Helen F. Evans

December 2006

Chair: James E. T. Channell
Major: Geology

This dissertation presents the results of chronostratigraphic studies on marine sediment

cores from three Oceans. Using a combination of magnetic stratigraphy, biostratigraphy and

cycle stratigraphy it is possible to produce chronostratigraphies that exceed the resolution of

any individual technique.

In the North Atlantic, environmental magnetic records from Eirik Drift, south of

Greenland, record detrital signals related to the melting of the Greenland and Laurentide Ice

Sheets. The detrital layer stratigraphy has been placed in a paleointensity-assisted

chronostratigraphic template, based on paleointensity and stable isotope data, to enhance

correlation of detrital layers across the North Atlantic region. In the central Atlantic, on Gardar

Drift, correlation of a benthic oxygen isotope record to the Greenland and Vostok Ice cores has

placed cores from the drift on a revised GISP chronology. A stack of relative paleointensity

records was developed and placed on the revised GISP chronology. In marine isotope stage 3, a

benthic isotope record appears to record changes in bottom water temperature that are coeval

with magnetic grain size changes.

IODP Site U1313 from the North Atlantic produced a high-resolution polarity

stratigraphy and relative paleointensity record between 2.5 and 6.0 Ma. This is one of a handful









of paleointensity records for this interval. Cycles in magnetic susceptibility allowed age-

calibration by correlation to a benthic oxygen isotope stack.

Sediment cores from the Pacific Ocean produced excellent magnetic stratigraphies, and

cycles in the sediment allowed astronomic calibration of reversal boundaries. Based on the

correlation of planktonic foraminifer datums to the magnetic stratigraphy at ODP Site 1208, a

new planktonic foraminifer zonation for the northwest Pacific Ocean can be precisely correlated

to polarity chrons and astronomically calibrated ages. Numerous paleomagnetic excursions are

tentatively identified for the first time in Pacific sediments.

Oxygen isotope records from the Late Miocene (9.3-11.2 Ma) at ODP Site 1092 (South

Atlantic) allowed astronomic calibration of ages of reversal boundaries and three polarity

excursions within Chron 5. This is the first time astronomically calibrated ages have been

assigned to these polarity excursion chrons and indicate a duration for the excursions of 3-4 kyrs.









CHAPTER 1
INTRODUCTION

Stratigraphy is a fundamental part of Geology. Earth processes unfold over a great range of

time scales from millions of years to minutes and seconds. One of the challenges in stratigraphy

is to be able to assign dates to events in the geologic record. The geologic timescale is the means

by which we can understand the history of the Earth and magnetic reversal stratigraphy provides

the central framework for the geologic timescale to which other dating techniques

(biostratigraphic, radiometric, orbital) can be tied. This is because magnetic reversals are, on

geologic timescales, globally synchronous, environmentally independent events.

The geomagnetic timescale of Heirtzler et al. (1968) was one of the foundations of the

plate tectonic revolution. They proposed a geomagnetic polarity timescale for the Late

Cretaceous to Recent based on a few long magnetic anomaly profiles. The evolution of the

polarity timescale since 1968 has involved two types of revisions: adjustments of the relative

spacing of some anomalies and calibration of the polarity sequence in time (Cande and Kent,

1992). Over the past forty years the pattern of normal and reversed polarities has been

extensively studied and most of its large-scale features for the past 200 million years are now

well understood (Gradstein et al., 2005).

Classic magnetic polarity reversal stratigraphy lacks the resolution necessary for the

high-resolution (millennial-scale) climate studies being conducted today. This led to the

development of high-resolution cryogenic magnetometers capable of measuring whole-core

samples or u-channel samples. This in turn led to the development of "composite sections" for

marine sediment cores whereby multiple cores were taken at a single site and spliced together to

provide a complete stratigraphic section (Hagelberg et al., 1995).









Changes in the intensity of the Earth's magnetic field occur over much shorter timescales

than polarity reversals. These changes can be measured in sedimentary cores to produce records

of relative geomagnetic paleointensity. This is done by normalizing the natural remanent

magnetization by an artificial remanence to remove intensity changes due to changes in

concentration of magnetic material in the core. Records of relative paleointensity have been

shown to be globally correlative on millennial timescales for the last glacial cycle (Laj et al.

2004).

In attempting to understand the time-depth relationship in marine sediment cores and

therefore understand more about the Earth's climate and evolution my work covers three Oceans,

the South Atlantic, the North Atlantic and the Pacific. Below is a summary of the work presented

in this dissertation. The nature of this work is collaborative and, as such, data provided by my

colleagues is included in this dissertation. Their contribution is acknowledged and clearly

detailed in the following summary.

In Chapter 2 magnetostratigraphic and cyclostratigraphic results are presented for the 0-

12 Ma interval from sites drilled during ODP Leg 198 to Shatsky Rise. Cyclic alternations in the

percentage of calcium carbonate, as shown by color reflectance data and gamma ray attenuation

bulk density measured on the sediments, allowed astronomic calibration of the magnetic

stratigraphy from six ODP Sites. This chapter was published in the Scientific Results Volume for

Ocean Drilling Program (ODP) Leg 198 (Evans et al., 2005). Chapter 3 is a continuation of this

work and has produced an integrated magneto- bio- and cyclostratigraphy from ODP Site 1208

for the 1-12 Ma interval. Biostratigraphic data included in this chapter were provided by

Nicholas Venti, Mark Leckie (U. Massachusetts, foraminifer) and Paul Bown (University

College London, nannofossils).









In Chapters 4 and 5, I used sedimentary relative paleointensity records to correlate

between cores collected on drift deposits in the North Atlantic. In Chapter 4, I present a study of

sediments from the Eirik drift for the 0-400 ka interval. Detrital layers identified within four

cores are placed in a paleointensity assisted chronostratigraphic framework. Environmental

magnetic records from climatically sensitive regions such as the North Atlantic can provide

information about changes in the strength of bottom currents and ice sheet dynamics both of

which are climatically sensitive. Oxygen isotope data used in this chapter were provided by Jim

Wright and Lauren Nietzke (Rutgers University) and Claude Hillaire-Marcel (GEOTOP) (Core

MD99-2227). This chapter is under review in the journal Geophysics, Geochemistry and

Geosystems. In Chapter 5 cores from the Gardar Drift provide records of changes in magnetic

grain size over glacial/interglacial and stadial/interstadial cycles for the last 130 ka. These

changes are interpreted as changes in the speed of bottom currents forming the drift deposits over

glacial/interglacial cycles and stadial/interstadial cycles. David Hodell (UF) provided oxygen

isotope data in Chapter 5.

In Chapter 6 a paleomagnetic study of IODP Site U1313 from the North Atlantic is

presented. The magnetic stratigraphy spans the interval from 2.5-6.3 Ma including the Gauss and

Gilbert chronozones. A relative paleointensity record for the Gauss and Gilbert chrons, is one of

only a handful of such records for this time interval. Cycles in magnetic susceptibility have

allowed astronomic calibration of the ages of reversal boundaries.

In 2001, my MS thesis consisted of a paleomagnetic study of ODP Site 1092 from the

South Atlantic. Chapter 7 presents a revision of the composite depth scale from ODP Site 1092.

X-Ray fluorescence (XRF) scanning data were provided by Thomas Westerhold (University

Bremen). This chapter was published in Geophysical Journal International (Evans et al. 2004).









In Chapter 8 we use cyclic alternations in a stack of three oxygen isotope records (Paulsen et al.,

in press) from ODP Site 1092 in the South Atlantic to astronomically tune the magnetic

stratigraphy from 9.3-11.2 Ma. This includes the long normal polarity chron C5n.2n and three

short reverse polarity intervals within it identified by Evans and Channell (2003). It also includes

a critical age tie-point from the Cande and Kent (1995) Geomagnetic Polarity Timescale (GPTS)

at the base of C5n.2n. This chapter is under review at Earth and Planetary Science Letters.









CHAPTER 2
LATE MIOCENE-HOLOCENE MAGNETIC POLARITY STRATIGRAPHY AND
ASTROCHRONOLOGY FROM ODP LEG 198-SHATSKY RISE

Introduction

Shatsky Rise is a medium-sized large igneous province in the west-Central Pacific Ocean

(Figure 2-1) and is possibly the oldest existing oceanic plateau. The rise consists of three

prominent topographic highs. Sites 1209, 1210, 1211 and 1212 were cored on the Southern High

(Bralower, Premoli Silva, Malone et al., 2002). Eight sites on the Southern High of the rise were

drilled during Deep Sea Drilling Project (DSDP) and earlier Ocean Drilling Program (ODP) legs

(Sites 47, 48, 49, 50, 305, 306, 577, and 810). Of these, ODP Sites 577 and 810 provided

interpretable Neogene magnetic stratigraphies.

Sites 1207 and 1208, from the Northern and Central Highs, provided unexpectedly

expanded late Miocene (12.5 Ma) to Holocene sequences. These locations had not been cored

during previous DSDP/ODP expeditions. The initial age model for all of the sites was based on

correlation of the sequence of polarity zones to the geomagnetic polarity timescale (GPTS)

(Cande and Kent, 1992, 1995). Mean sedimentation rates at the five sites vary from 1- to 4

cm/k.y. Latitude and longitude of the sites and basal ages of the Neogene sediments are given in

Table 2-1. Neogene sediments at the sites consisted mostly of light gray to pale orange

nannofossil oozes with varying amounts of clay, radiolarians, and diatoms. Magnetic

susceptibility is low (< 2 x10-5 SI) at all the sites and shows a decreasing trend from the

Quaternary to the late Miocene. Composite sections were constructed shipboard for four of the

sites (1209, 1210, 1211, and 1212) using multi-sensor track (MST) data including magnetic

susceptibility, gamma ray attenuation (GRA) bulk density, and reflectance data. Sites 1207 and

1208 were not double-cored, and depths at these sites are in meters below sea floor (mbsf). The









magnetic stratigraphy from the six sites (1207, 1208, 1209, 1210, 1211, and 1212), was based on

shipboard pass-through magnetometer measurements and discrete samples measured post-cruise.

Sediments from five of the sites (1207, 1208, 1209, 1210, and 1211) showed a prominent

cyclicity in reflectance data for parts of the sections, and this is the basis for the construction of

an astronomically tuned age model for the 0- to 8-Ma interval. The astronomically calibrated

polarity timescale has been well established for the 0- to 6-Ma interval (Shackleton et al., 1990,

1995; Hilgen 1991a, 1991b). Hilgen (1991a, 1991b) produced his astronomically calibrated

polarity timescale for the 2- to 5.23-Ma interval using sapropel occurrences and carbonate

content in Mediterranean sections. These polarity chron ages were incorporated into the GPTS of

Cande and Kent (1995).

In this study we produced an astronomically calibrated magnetic reversal stratigraphy for

the 0- to 8-Ma interval. This is in good agreement with Hilgen (1991a, 1991b) and Shackleton et

al., (1995) in the 0- to 6-Ma interval. In the 6-to 8-Ma interval, polarity chron ages are in better

agreement with the Shackleton et al. (1995) timescale, differing by up to -200 k.y. from that of

Hilgen et al. (1995) and the ATNTS 2004 of Lourens et al. (2004). This chapter was published in

the Scientific Results Volume for ODP Leg 198 (Evans et al., 2005).

Methods

Two types of paleomagnetic measurements were made on sediments collected during ODP

Leg 198; pass-through measurements on half-cores and discrete sample measurements. Discrete

sample cubes (2cm x 2cm) were collected during Leg 198 to augment measurements using the

shipboard pass-through magnetometer. Shipboard measurements on half-cores were made at 5-

cm intervals. A total of 747 discrete samples were taken at 50-cm intervals. Discrete samples

were collected from the center of the half-cores to avoid deformation at the outer edges of the

core. Magnetic measurements on the cubes were performed in the magnetically shielded room at









the University of Florida using a 2G-Enterprises cryogenic magnetometer. The samples were

step-wise alternating-field (AF) demagnetized using a D-Tech D2000 AF demagnetizer.

Magnetization component directions were determined using the method of Kirschvink (1980),

applied to the 20- to 60 mT peak field demagnetization interval.

The astrochronology developed for Sites 1207, 1208, 1209, 1210, and 1211 was based on

cycles seen in reflectance data (L*) measured shipboard on a purpose-built track. Reflectance of

visible light from soft sediment cores was measured using a spectrophotometer at 2.5-cm

intervals and provided a high-resolution record of color variations for visible wavelengths (400-

700 nm). L* reflectance represents "lightness" of the sediment which is usually controlled by

changes in percent carbonate.

The initial age model for each site was based on correlation of the polarity zone sequence

to the timescale of Cande and Kent (1995). Power spectra using the Blackman-Tukey method

with a Bartlett window from the Analyseries software of Paillard et al. (1996) indicate the

presence of obliquity and eccentricity peaks. The reflectance data were then tuned to the

astronomic solutions for obliquity from Laskar et al. (1993). This allowed astronomically

calibrated ages to be assigned to the polarity reversal boundaries at Sites 1207, 1208, 1209, 1210

and 1211. Site 1212 was not included in the astrochronology, as it contains a hiatus at 4- to 5-

Ma.

Magnetostratigraphic Interpretation

Site 1207 is the only site that has been drilled on the Northern High of Shatsky Rise. The

sequence of sediment recovered was mostly Neogene in age (0-163.8 mbsf) underlain by

Campanian and older oozes and cherts. The sediment consists of nannofossil ooze with diatoms,

radiolarians, and clay in varying amounts (Bralower, Premoli Silva, Malone et al., 2002). The

samples taken for paleomagnetic analysis were AF demagnetized in 5-mT steps up to either 50,









60, or 70 mT, depending on the intensity of the natural remanent magnetization (NRM). Less

than 10% of the NRM remains after demagnetization at these peak fields, indicating a low-

coercivity remanence carrier, most likely magnetite. Orthogonal projections of demagnetization

data (Figure 2-2) show well-defined components for most of the Neogene section after removal

of the steep drilling related overprint at peak AF fields of 20 mT. Maximum angular deviation

(MAD) values are low for most of the section (< 100), indicating well-defined characteristic

magnetization components; however, some intervals, particularly the interval between 50-60

mbsf (Figure 2-3), have slightly higher MAD values and less well-defined components. The

interpretation of the magnetic stratigraphy from shipboard and discrete sample data can be

accomplished by polarity zone pattern fit to the GPTS (Cande and Kent, 1992, 1995) (Table 2-2).

This pattern fit is satisfactory to the base of Subchron C5An. in (Figures 2-3, 2-4). Below the

polarity zone equivalent to Subchron C5An. In, recovery was intermittent and biostratigraphy

indicates a hiatus with Campanian age sediments below (Bralower, Premoli Silva, Malone et al.,

2002). Sedimentation rates average 1-2 cm/k.y. throughout the section with some slightly higher

(3-4 cm/k.y.) rates in the late Pliocene and late Miocene (Figure 2-5A). Component declination

has been corrected for each core using Tensor orientation data measured shipboard. The mean

inclination in normal polarity zones for the Site is 57.80, close to the expected inclination of 560

for a geocentric axial dipole at this site; however, reversed polarity intervals have a mean

inclination of -51.1, shallower than expected. This can be attributed to shallowing of reversed

polarity directions by the steep downward-directed drilling overprint, shown clearly in the

orthogonal projections (Figure 2-2A).

Sites 1209, 1210, 1211, and 1212 are located on the southern high of Shatsky Rise

(Figure 2-1). Multiple holes were drilled at each site and composite sections were constructed









using shipboard MST data. Discrete sample cubes were only collected from Holes 1209A,

1210A, 1211A, and 1212A. The shipboard data from the pass-through magnetometer are

consistent between the different holes at each site and confirms the interpretation of the magnetic

stratigraphy (see Shipboard Scientific Party, 2002).

As for Site 1207, orthogonal projections from discrete sample data show two components:-

a steep downward drilling related overprint and well-defined characteristic components (Figure

2-2 B, C, D, E). In most cases the drilling related overprint was easily removed in peak AF fields

of 10-20 mT. Little of the natural remanent magnetization remained at peak fields of 60 mT.

MAD values are generally <50 throughout the sections. The expected inclination for the

Southern Rise is 51; again, all the sites show slightly steeper than expected inclinations in

normal polarity zones and shallower than expected inclination in reversed polarity zones. The

magnetostratigraphic age models indicate mean sedimentation rates between 1- and 3 cm/k.y. for

most of the Neogene (Figure 2-5 B, C, D, E).

The polarity interpretation at Sites 1209, 1210, and 1211 is unambiguous back to Subchron

C3Bn (Table 2-2) (Figures 2-6, 2-7, 2-8). Below this level, interpretation becomes difficult due

to decreasing sedimentation rates leading to a hiatus recognized at all sites between the upper

Miocene, and Oligocene and older sequences. At Site 1212, a hiatus accounts for the interval

between 4 and 5 Ma (Chron C3), and the polarity interpretation can be accomplished to

Subchron C4n.2n (Figure 2-9). This interpretation of the sequence of polarity zones is confirmed

by the shipboard biostratigraphy. The interpretation of the polarity stratigraphy was carried out

using data measured shipboard augmented with discrete sample cubes. When the

magnetostratigraphic data were placed on the composite depth scale, the reversal boundaries

were found to be consistent between holes, indicating that there is very little error in the depths









of polarity zone boundaries or in composite depth calculations. The magnetic measurements

made shipboard do include a small amount of error due to the response function of the shipboard

magnetometer. The response function of the wide-access magnetometer used to measure half-

cores is -10 cm, resulting in a cm-scale uncertainty in the placement of the reversal boundaries.

Site 1208 is located on the Central High of Shatsky Rise and also provided an expanded

late Miocene to Holocene section. The magnetic stratigraphy from Site 1208 will be presented in

Chapter 3.

Astrochronology

Cycles were visually identifiable in L* reflectance data from all six of the sites in this

study. For Sites 1209, 1210, and 1211, we worked with spliced composite records rather than

data from a single hole. Reflectance data were initially placed on the magnetostratigraphic age

model based on the polarity timescale of Cande and Kent (1995). Power spectra for untuned

sections of reflectance data placed on this age model consistently show a concentration of power

at orbital frequencies, particularly around the 41 k.y. obliquity cycle (Figure 2-10).

The reflectance records were then tuned to the astronomical solution for obliquity from

Laskar et al. (1993), as this was the most visually identifiable cycle in the reflectance data and

the power spectra for different time intervals in all the sites showed a concentration of power at

the obliquity frequency (Figure 2-10). In constructing the astrochronological age model, we

assume that there was no phase lag between the orbital forcing and the response. For

convenience, the reflectance data were broken up into 1 Myr intervals when compared to the

astronomical solution and each site was tuned independently. Cycles were readily apparent in the

reflectance data for all sites, and tuning of the record required a minimum of adjustment of peaks

in the reflectance data to the astronomical solution (Figures 2-11, 2-12, 2-13, 2-14).

Astronomically tuned ages were calculated for polarity reversals in the 1 to 8-Ma interval at Site









1207 (Table 2-3). At Site 1209, tuning was performed in the 1 to 7-Ma interval and at Sites 1210

and 1211 in the 1 to 5-Ma interval. Site 1208 has also provided an astrochronological age model

for the 1 to 6-Ma interval (Figure 2-12) and is included in Table 3. The tuned age models are

compared to each other (Table 2-3) and are compared with other recently published

astrochronologies for this time period (Table 2-4). The output of a band-pass filter centered on

41 k.y. is shown below the astronomical solution for obliquity and the raw reflectance data in

Figures 2-11, 2-12, 2-13, 2-14, and 2-15.

To test the validity of the timescale we used cross-spectral analysis performed using the

Blackman-Tukey method and Analyseries software (Paillard et al., 1996). Coherence between

the reflectance data and the astronomical solution for obliquity was significant at all the sites,

although the coherence values depend on which time interval is being examined. At Site 1207

coherence was 0.8 for the 1.2 to 1.8-Ma and 6.2 to 6.8-Ma intervals (Figure 2-16A). The

coherence values at Site 1208 were > 0.8 for the entire 1 to 6-Ma interval. Sites 1209, 1210, and

1211 also showed coherence values between 0.8 and 1 (Figure 2-16C, 2-16D, 2-16E).

Discussion

Comparison of the tuned ages for polarity reversal boundaries at the five sites in the 1.5

to 2-Ma interval showed that polarity chron ages are in good agreement. For other time intervals

there are some significant differences (more than an obliquity cycle) between sites (Table 2-3).

Intervals with enhanced 41 k.y. power in reflectance data are considered more reliable (italics in

Table 2-3). Site 1208 showed the strongest cyclicity, with Site 1207 also showing a clear signal

in some intervals particularly the 2.1 to 2.7-Ma and 4.5 to 5-Ma intervals.

During ODP Leg 138 to the eastern equatorial Pacific, 11 sites were drilled and most of

them showed a prominent cyclicity in GRA density. Shackleton et al. (1995) used these cycles in

GRA bulk density records to produce an orbitally-tuned age model for the 0 to 12.5-Ma interval.









They worked entirely in the time domain comparing smoothed GRA bulk density records with

the target record of summer insolation at 650N. In their tuning they assumed that no phase lag

existed between insolation and GRA bulk density controlled by proportion of Si02 and CaCO3

(high density), high carbonate content being associated with high Northern Hemisphere

insolation. Age control points were added to the data to align prominent groups of density

maxima. The records were broken into 0.8-m.y. intervals for convenient viewing. Each site was

tuned independently over the chosen time interval. Shackleton et al. (1995) found that some

intervals in these records were more easily tuned than others, similar to results from Leg 198.

Shackleton et al. (1995) noted that it was difficult to tune the 0- to 1-Ma interval, which was also

the case at four of the Leg 198 Sites (1208, 1209, 1210, and 1211). The 1- to 2-Ma interval for

the Leg 138 sites carries a clear 41 k.y. obliquity cycle. For Leg 198 sites, the 1.2- to 1.6-Ma

interval also carries a very clear obliquity cycle (Figures 2-11, 2-12, 2-13, 2-14, 2-15). In the 2.4-

to 2.6-Ma interval, a very strong obliquity cycle was observed at Site 846 (Leg 138), and this

same interval also carries a strong 41 k.y. signal at Sites 1209, 1210, and 1211. Comparison

between the Site 1207 age model and ages from Shackleton et al. (1995) indicate consistency for

the 1-to 8-Ma time interval (Table 2-4).

Hilgen et al. (1995) developed an astronomical timescale for the interval from 3- to 9.7-Ma

using lithologic cyclicity seen in sedimentary sections in the Mediterranean. These sections

comprise open marine sediments that alternate between carbonate-rich and carbonate-poor marls

or homogeneous marls and sapropels. The individual sapropels were related to precession

minima, and the clusters of sapropels to the 400-k.y. eccentricity cycles. In tuning the section,

the target curve used was the 65N summer insolation curve. To obtain an astronomical age for

the youngest polarity reversal in the sequence, Hilgen et al. (1995) took the Shackleton et al.









(1995) age for the onset of Subchron C3An.2n of 6.576 Ma. They then matched the lithologic

cycles in the section to the astronomical solution using the correlation of sapropel clusters to

eccentricity. The age of the calibration point (6.576 Ma) had to be adjusted to 100 k.y. older to

establish a consistent correlation between sapropel clusters and eccentricity maxima. The ages

from Hilgen et al. (1995) differ significantly with those from Leg 198 in the 6-to 8-Ma interval

(Table 4). At the top of Subchron C3Bn the difference is more than 200 k.y. In the interval from

7.2- to 8.1-Ma, the difference is 100 k.y. which is the amount of adjustment of the 6.576-Ma

tie point used by Hilgen et al. (1995) for the age of the youngest polarity reversal in their section.

ODP Site 926 on the Ceara Rise also produced an orbitally tuned timescale from 5- to 14-

Ma (Shackleton and Crowhurst, 1997). This timescale cannot be directly compared with the Leg

198 timescale because of a lack of polarity reversals at Site 926. Backman and Raffi (1997) used

the cyclostratigraphic age model from Site 926 to calibrate ages of the calcareous nannofossil

datums for the late Miocene. These ages were then compared with the biomagnetochronology

from Site 853 (ODP Leg 138). The center of the peak in abundance of transitional morphotypes

of Triquetrorhabdulus rugosus at Site 853 occurred 120-130 k.y. after the corresponding peak at

Site 926. The age estimates of Hilgen et al. (1995) were then applied to the Site 853 data and the

peak center was found to coincide at Sites 853 and 926. Therefore, Backman and Raffi (1997)

considered that the Hilgen et al. (1995) ages are more reliable in this interval than the ages of

Shackleton et al. (1995).

Lourens et al. (2004) have recalibrated the Miocene astronomic timescales of Shackleton

and Crowhurst (1997) and Hilgen et al. (1995) using the astronomic solution of Laskar et al.

(2004). For the last 13 Ma the returning resulted in almost negligible changes in the ages of

reversal boundaries (Lourens et al., 2004). For the 6- to 8- Ma interval the ATNTS2004 is in









close agreement with that of Hilgen et al (1995) and therefore differs significantly with the

results of this study.

Conclusions

Five sites from Shatsky Rise have produced high-quality magnetic stratigraphies from the

late Miocene to Holocene. Cycles identified in reflectance data from Sites 1207, 1208, 1209,

1210, and 1211 have allowed astronomic calibration of the polarity reversal sequence from -8

Ma to present. The assumption that there is no phase lag between sedimentary cyclicity and the

astronomical parameters allowed the cycles to be tuned to the astronomical solution for

obliquity. Cross-spectral analysis on the tuned age model indicated high coherence between the

astronomic solution and the reflectance data and confirms the reliability of the tuning. The age

model has been compared with other published astrochronologies and is found to be in good

agreement with Hilgen (1991a, 1991b) (and, therefore, Cande and Kent [1995]) in the 1-to 6-Ma

interval. In the 6-to 8-Ma interval the age model differs significantly from that of Lourens et al.

(in press) and Hilgen et al. (1995) from the Mediterranean. It is in better agreement with the

ODP Leg 138 timescale of Shackleton et al. (1995) from the Pacific Ocean.











Table 2-1. Latitude, longitude, water depth, the oldest Neogene magnetic polarity chron
identified, and the basal age of the Neogene section.


Latitude

3747.4287' N
32 39.1001'N
32 13.4123'N
32 0.1300'N
32 26.9000'N


Longitude

16245.0530'E
15830.3560'E
15815.5618'E
15750.9999'E
15742.7016'E


Water Basal Chron Basal Age
depth (Ma)
3100m C5An2n 12.184
2387m C3Bn 7.091
2573m C3Bn 7.091
2907m C3Bn 7.091
2682m C4n.2n 8.072


Site

1207
1209
1210
1211
1212













Table 2-2. Magnetostratigraphic age model for Sites 1207, 1209, 1210, 1211 and 1212. Polarity
chron labels are according to Cande and Kent (1992, 1995). Ages of chrons are from

Cande and Kent (1995). Depths are in meters below sea floor (mbsf) for Site 1207
and meters composite depth (mcd) for Site 1209, 1210, 1211 and 1212.


Chron
Cln
base
Clr.1n
base
C2n
base
C2r. ln
base
C2An. in
base
C2An.2n
base
C2An.3n
base
C3n.ln
base
C3n.2n
base
C3n.3n
base
C3n.4n
base
C3An.ln
base
C3An.2n
base
C3Bn
base
C3Br.n n
base
C3Br.2n
base
C4n.ln
base
C4n.2n
base
C4r. ln
base
C4An
base
C4Ar. In
base
C4Ar.2n
base
C5n.ln
base
C5n.2n
base
C5r.ln
base
C5r.2n
base
C5An.ln
base
C5An.2n
base


Ma (CK95)
0.00
0.780
0.990
1.070
1.770
1.950
2.197
2.229
2.581
3.040
3.110
3.220
3.330
3.580
4.180
4.290
4.480
4.620
4.800
4.890
4.980
5.230
5.894
6.137
6.269
6.567
6.935
7.091
7.135
7.170
7.341
7.375
7.432
7.562
7.650
8.072
8.225
8.257
8.699
9.025
9.23
9.308
9.580
9.642
9.740
9.880
9.920
10.949
11.052
11.099
11.476
11.531
11.935
12.078
12.184
12.401


1207 (mbsf)
0.00
12.35
16.26
16.77
24.38
28.40
29.73
30.25
43.13
51.77
53.25
56.79
58.77
66.91
80.23
83.34
87.19
90.46
92.39
94.32
96.84
99.95
105.73
106.77
109.29
114.18
116.56
120.41


1209 (mcd)
0.00
11.28
13.32
14.22
25.28
28.21


37.69
49.43


52.34
58.03
66.22
68.23


73.24
74.08
75.59
78.76
82.94
84.95
86.12
90.13
93.81
96.15


1210 (mcd)
0.00
14.89
18.07
19.71
32.03
34.70
37.68
38.09
46.51
56.88
58.52
60.37
61.81
67.35
75.36
77.62
80.90
82.34
83.78
85.42
86.45
91.17
94.05
95.69
97.02
99.12
100.35
101.46


1211 (mcd)
0.00
8.00
9.550
10.27
16.74
18.38
23.70
30.08
30.90
32.34
33.98
37.37
41.17
42.51
43.94
44.66
46.92
49.28
50.72
51.70
52.69
53.80
54.29
55.15


123.08
123.53
125.16
126.05
126.34
129.01
130.79
132.27
134.50
136.72
137.76
138.35
140.28
140.72
141.32
142.36
142.65
151.40
153.62
154.07
155.11
155.40
157.81
160.77
161.76


1212 (mcd)
0.00
11.95
14.12
14.98
23.62
25.78
26.95
27.78
32.61
39.00
39.81
41.67
43.00
48.68










54.44
56.17
58.31
59.14
59.88
61.03
61.60
63.58
64.57
65.14
65.39
66.95
67.37
70.00











Table 2-3. Comparison of astrochronological age models for sites 1207, 1208, 1209, 1210 and
1211. Italics indicate the most reliable ages in intervals where the cyclicity in
reflectance is best defined. In italics and brackets are the differences between tuned
ages and those of Cande and Kent (1995).


Chron Ka
(CK95)
(Hilgen
1991a,b)
Cln 0
Clr.lr 780
Clr.ln 990
Clr.2r 1070
C2n 1770
C2r. r 1950
C2r.ln 2140
C2r.2r 2150
C2An.ln 2581
C2An.lr 3040
C2An.2n 3110
C2An.2r 3220
C2An.3n 3330
C2Ar 3580
C3n.ln 4180
C3n.lr 4290
C3n.2n 4480
C3n.2r 4620
C3n.3n 4800
C3n.3r 4890
C3n.4n 4980
C3r 5230
C3An.ln 5894
C3An.lr 6137
C3An.2n 6269
C3Ar 6567
C3Bn 6935
C3Br.lr 7091
C3Br.ln 7135
C3Br.2r 7170
C3Br.2n 7341
C3Br.3r 7375
C4n. n 7432
C4n. r 7562
C4n.2n 7650
C4r. r 8072


Site 1207
Ka
(difference)
0
776.7 (-3.3)
992.8 (2.8)
1089.4 (19.4)
1786.4 (16.4)
1954.2 (4.2)
2095.7 (-44.3)
2112.0 (-38)
2620.5 (39.5)
3042.5 (2.5)
3118.0(8)
3236.5 (16.5)
3354.5 (24.5)
3593.3 (13.3)
4154.0 (-26)
4262.5 (-27.5)
4489.8 (9.8)
4637.0(17)
4760.5 (-39.5)
4857.3 (-32.7)
4972.5 (-7.5)
5245.4 (15.4)
5886.0 (8)
6143.0 (6)
6241.5 (-27.5)
6526.2 (-40.8)
6878.0 (-57)
7095.8 (4.8)


75-3 2 (8.2)
7388.3 (13.3)
7453.5 (3.5)
7540.9 (-21.1)
7634.1 (15.9)
8038.0 (-34)


Site 1208
Ka
(difference)




1073.9 (3.9)
1776.2 (6.2)
1948.7 (-1.3)
2133.5 (-6.5)
2170.4 (20.4)
2564.7 (-16.3)
3045.2 (5.2)
3105.8 (-4.2)
3229.8 (9.8)
3340.9 (10.9)
3599.6 (19.6)
4190.9 (10.9)
4351.9 (61.9)
4523.6 (43.6)
4683.8 (63.8)
41'*,. ) (6.9)
41i '9 (-9.1)
4991.8 (11.8)
5201.2 (-28.8)
5952.7 (58.7)


Site 1209
Ka
(difference)




1069.4 (-0.6)
1770.0 (0)
1975.4 (25.4)


2550.3 (-30. 7)
3032.0 (-8)


3361.4 (31.4)
3648.8 (68.8)
4172.5 (-7.5)
4305.9 (43.4)


4809.0 (9)
4880.9 (-9.1)
4981.5 (1.5)
5240.2 (10.2)
5915.8 (21.8)
6073.9 (36.9)
6318.3 (49.3)
6548.3 (-18. 7)
6971.3 (35.3)
7027.7 (-63.3}


Site 1210
Ka
(difference)





1777.8 (7.8)
1972.2 (22.2)


2642.7 (61.7)
3032.9 (-7.1)
3114.9 (4.9)
3248.5 (~\ 5)
3340.5 (10.5)
3597.5 (17.5)
4182.8 (2.8)
4305.9 (15.9)
4501.0 (21)
4665.3 (45.3)
4798.8 (-1.2)
4891.2 (1.2)
4973.3 (-6.7)


Site 1211
Ka
(difference)





1777.8 (7.8)
1972.2 (22.2)


2536.1 (-44.9)
3022.2 (-17.8)
3110.9 (0.9)
3242.3 (22.3)
3352.8 (22.8)
3644.4 (64.4)
4169.4 (-10.6)
4305.6 (15.6)
4457.9 (-22.1)
4589.3 (-30.7)


4950.7 (-29.3)









Table 2-4. Astrochronological ages for Leg 198 compared to ages Hilgen (1991a, 1991b), Hilgen
et al. (1995) and Shackleton et al, (1995). In italics and brackets are the differences
between Leg 198 tuned ages and Hilgen et al. (1995) and Shackleton et al. (1995).


Chron


Cln
base
Clr.ln
base
C2n
base
C2r. ln
base
C2An. in
base
C2An.2n
base
C2An.3n
base
C3n.ln
base
C3n.2n
base
C3n.3n
base
C3n.4n
base
C3An. in
base
C3An.2n
base
C3Bn
base
C3Br.ln
base
C3Br.2n
base
C4n. ln
base
C4n.2n
base


Ka (CK95)
Hilgen
(1991a,b)

0
780
990
1070
1770
1950
2140
2150
2581
3040
3110
3220
3330
3580
4180
4290
4480
4620
4800
4890
4980
5230
5894
6137
6269
6567
6935
7091
7135
7170
7341
7375
7432
7562
7650
8072


Leg 198





776.7
992.8
1089.4
1786.4
1954.2
2133.5
2170.4
2564.7
3042.5
3118.0
3236.5
3354.5
3593.3
4190.9
4351.9
4523.6
4683.8
4806.9
4880.9
4972.5
5201.2
5952.7
6143.0
6241.5
6526.2
6878.0
7095.8


7348.2
7388.3
7453.5
7540.9
7634.1
8038.0


Shackleton
et al. (1995)
(difference
to 198)











3046 (3.5)
3131 (13)
3233 (-3.5)
3331 (-23.5)
3594 (.7)
4199 (8.1)
4316 (-35.9)
4479 (-44.6)
4623 (-60.8)
4781 (-25.9)
4878 (-2.9)
4977 (4.5)
5232 (30.8)
5875 (-77.7)
6122 (-21)
6256 (14.5)
6555
6919 (4)
7072 (-23.8)





7406 (-47.5)
7533 (-7.9)
7618 (-16.1)
8027 (-11)


Hilgen et al.
(1995a)
(difference
to 198)






























6677 (150.8)
7101 (223)
7210 (114.2)
7256
7301
7455 (106.8)
7492 (103.7)
7532 (78.5)
7644 (103.1)
7697 (62.9)
8109 (71)
































Figure 2-1. Bathymetric map of Shatsky Rise showing the location of sites drilled during ODP
Leg 198.













S/U N/Up




W F
198-1207A-2H.2
mb.r= 7 07


L"_lreal =mml
HbI-a 17 H r- t = 50 mT
HL-treaat 0 m
S1,1n -07 30 5
5 n~ Hi~ tret = 50 mT n 5. I1n
Lui rc $/Pn


Lo Ireat = 0 T
Hi treal= 60 n







5/Un
C) N/Up



I -121(1A-H-
mrd = 0.48
Lo treat = OmT
HI treat 60 TI




'U,
5;;


19S1209A 2H
mcd I 55
II tIreat = 0 m
lli trial = t0 j


)A-2H
- 9.,5
=01
601


mcd= 102 55 S/Dn -
Hi-treat= r60 mT

198-l211A-15H-I
mcd = 120 0
Li-treat = 0 inT
Hi-(reat 60 mT


N/UP





'T <





: : ',-- E

"'UP




-3 I -121
iT Hi tea


i I
Q |


1212A 8H 2 1N14
T-rta m aT
H ll Nr \mlrT
s, S I1,


Figure 2-2. Representative orthogonal projections of AF demagnetization data from (A) Site

1207, (B) Site 1209, (C) Site 1210, (D) Site 1211 and (E) Site 1212. Low AF

demagnetization treatment and the high treatment are given, as is mbsf or mcd of the

sample. Open circles represent the vector end-point projection on the vertical plane,

while closed circles represent the vector endpoint projection on the horizontal plane.


N/Up



L-itfea = 0 05







Sin


IA- 3H,
at =0 inT
I 6 in


N/Up





-I--I
ii


.tp
i ,'


V 1212A-KH-I

Hi trc l liniT


s
Hcr;
n


Ifvn












Site 1207


Chrons


10 -t -




I0 r -
20











40
_. -- .



-t4

B _
50

--- __ _^^



60





70



80

80 I


Declination (0)


S tu ou 0 5 10 15 2025
Inclination (0)
MAD (0)


Figure 2-3. Site 1207 component inclination and declination from discrete samples (open
squares) for 0-80 meters. Inclination and rotated declination from the shipboard pass-
through magnetometer after AF demagnetization at peak fields of 20 mT (gray line).
Chrons are labeled according to Cande and Kent (1992). Black indicates normal
polarity, white reversed polarity. Also shown are the MAD values calculated for
discrete sample data (after Kirschvink, 1980).












Site 1207


80 I j -



--


II. _--I










-

110 -- --- ---
110





120






130



7-J


140





150 I I






160 I i


0 50 100 150 200 250 300 350-80
Declination (0)


Chrons


n--

A-I-

-n


i











~i


%i
-





.= -- ; *- -





~ --i-i
-==3'-



c--------


C3n.2n


C3n.3n


C3n.4n


C3r


C3An.In


C3An.2n


C3Ar

C3Bn

C3Br
C4n.,n

C4n.2n

C4r

C4An

C4Ar

C5n.In


C5n.2n
cryptochron
C5n.2n-3 ?

C5r


C5An.ln


0 40 80
Inclination (0)


u
rr;
o

r





',





~

i'

e




~3
rr=
~20


0 5 10 15 2025
MAD (o)


Figure 2-4. Site 1207 component inclination and declination from discrete samples (open
squares) for 80-160 meters. Inclination and rotated declination from the shipboard
pass-through magnetometer after AF demagnetization at peak fields of 20 mT (gray
line). Chrons are labeled according to Cande and Kent (1992). Black indicates normal
polarity, white reversed polarity. Also shown are the MAD values calculated for
discrete sample data (after Kirschvink, 1980).


V-



-c
C
















20-

40




80


c)

140

160
0 2
























e) o
20


40

E60


80










0) 0 _


1207A age model
I I


6 8 10 12
Age (Ma)


1210 age model


3 4 5 6
Age (Ma)


60 b)





40





to d

0 Io-
0


1209 Age model
















S 2 3 4 5 6 7 8
Age (Ma)


1211A age model


50


7 8


2 3


4
Age (Ma)


1212age model


4
Age (Ma)


Figure 2-5. Interval sedimentation rates and age versus depth for the initial age model at a) Site
1207, b) Site 1209, c) Site 1210, d) Site 1211 and e) Site 1212.


30





20
25 |

















20





a
20



1s

S-


iS











Hole 1209A


I I r


20
20 _


40 ,






E -
S601
C,
C)


In-


Ic I I


Ii


-7



0 50 100 150 200 250 300 350-80
Corrected Declination (o)


-40 0 40
Inclination (o)


0 5 10 15 20
MAD (o)


Figure 2-6. Site 1209 component inclination and declination from discrete samples (open
squares). Inclination and rotated declination from the shipboard pass-through
magnetometer after AF demagnetization at peak fields of 20 mT (gray line). Chrons
are labeled according to Cande and Kent (1992). Black indicates normal polarity,
white reverse polarity. Also shown are the MAD values calculated for discrete sample
data (after Kirschvink, 1980).


Chrons


__ jl













Hole 1210A


_
S- I








nrA
i -






-----=^----


I-]-






I-I-

K5-


S--



'-I

i--
^__61~

-r-



i _1




;-l-










-i_ _--
~ ~ I
L1


I- -F-_


-Q


?--F



___ '~


---
I-- ------if--i "


0 50 100 150 200 250300350-80

Corrected Declination (0)


0 40 80

Inclination (0)


0 5 10 15
MAD (0)


Figure 2-7. Site 1210 component inclination and declination from discrete samples (open

squares). Inclination and rotated declination from the shipboard pass-through
magnetometer after AF demagnetization at peak fields of 20 mT (gray line). Chrons
are labeled according to Cande and Kent (1992). Black indicates normal polarity,
white reverse polarity. Also shown are the MAD values calculated for discrete sample
data (after Kirschvink, 1980).


I I


20- _
20

_-J i^


I '


Cd'


100


Chrons




Cln




Clr.lr
Clr.ln



CI r,2r



C2n

C2r.lr
C2r. In

C2r.2r




C2An.ln

C2An.lr
C2An.2n
C2An.2r
C2An.3n



C2Ar


C3n.ln

C3n.2n
C3n.3n

C3n.4n

C3r
C3An.ln

C3An.2n
C3Bn


Hiatus


^^,_



i-..










Hole 1211A


- 30 --O ,

a 2'



40






50-






60 I I 1
0 50 100 150 200250 300350 -80 -40 0 40 80
Corrected Declination (o) Inclination (0)


Chrons


0 5 10 15
MAD (o)


Figure 2-8. Site 1211 component inclination and declination from discrete samples (open
squares). Inclination and rotated declination from the shipboard pass-through
magnetometer after AF demagnetization at peak fields of 20 mT (gray line). Chrons
are labeled according to Cande and Kent (1992). Black indicates normal polarity,
white reverse polarity. Also shown are the MAD values calculated for discrete sample
data (after Kirschvink, 1980).










Hole 1212A Chrons
I II


Cln


10
Clr.lr
SC-:ln Cr.ln


20 I Clr.2r
20- -


C2n .1

------------_ --
30 9 C2r


C2An.ln
-,p


,s i C2An.Ir
40 C--2An.2n
C2An.2r
C2An.3n


50 Hiatus
0 C3r

r C3An.ln
S-o C3An.lr
C3An.2n
60- C3Bn
C3Br
C4n.In
C4n.2n

70 Hiatus
0 50 100 150 200 250 300350-80 -40 0 40 80 0 5 10 15 20
Corrected Declination (o) Inclination (0) MAD (0)

Figure 2-9. Site 1212 component inclination and declination from discrete samples (open
squares). Inclination and rotated declination from the shipboard pass-through
magnetometer after AF demagnetization at peak fields of 20 mT (gray line). Chrons
are labeled according to Cande and Kent (1992). Black indicates normal polarity,
white reverse polarity. Also shown are the MAD values calculated for discrete sample
data (after Kirschvink, 1980).

























0 0.02 0.04 0.06
frequency (k.y. -1)


Site 1207 5.2-5.8 Ma


obliquity


0.08 0.1 0 0.02 0.04 0.06 0.08 0.1
frequency (k.y. I)


Site 1207 7.2-7.7 Ma
obliquity


0 0.02 0.04 0.06
frequency -1)


Siobliquity




II


0.1 0 002


Site 1208 2-3 Ma


0.04 0.06
frequency (k.y. I)


0.08 0.


0 0.02 0.04 0.06
frequency (k.y. 1)


0.08 0.1 0 0.02 0.04 0.06
frequency (k.y. 1)


Site 1210 2.2-2.8 Ma


0 0.02 0.04 0.06
frequency (k.y. -I)


0.08 0.1


0.08 0.1


I I I
obliquity
ot Site 1211 2.2-2.8 Ma










0 0.02 0.04 0.06 0.08 0.1
frequency (k.y. -1)


Figure 2-10. Power spectra from a) Site 1207, b) Site 1208 c) Site 1209, d) Site 1210, and e) Site
1211 for reflectance data placed on a Cande and Kent (1995) age model.















SSite 1207


I ,,. .N -- 0 .

C 30 0.38
4 A-
-4 I
0 200 400 Age (ka) o 0 800o o0







101200 1400 .1600 100 200
00.4


-4-
000 1200 140 Age (ka)
SIJI
2000 2200 2400 2600 280
/ 0.43
740 C42
S60"" 0.4 1cC
.50 q.. Z. /NY








/N- I I I -


S4000 2200 2400 2 ( 4600 2Mo 3000
S, i Ii 03
-4




S000 200 5400 Age (ka) 56 00 5so 4000
M' 0.39 ,
,j 0.38










." I" ,. 41 '
-4


7000 7200 7400 Age (ka) 5600 7500 6onO
85

24-



6DW0 6200 6400 Age (ka) 660o 60{ 700L)

Figure 2-11 The astronomical solution for obliquity (Laskar et al., 1993) compared with tuned
', .',

K 111 I I I
y),
-4





7000 720 7400 Age (ka) 76,o -sm 800


Figure 2-11. The astronomical solution for obliquity (Laskar et al., 1993) compared with tuned

L* reflectance data from Site 1207 for the 0-8 Ma interval. The reflectance data

filtered using a band-pass filter centered on 41kyrs is shown in the lower part of each

frame. Black indicates normal polarity and white reverse polarity. Heavy line on L*

reflectance data indicates intervals where the cyclicity is best developed.













Site 1208


-I




-6---- I -
1000 1200 1400 4.,. 1600 1800 2000

-2


ii0 0J39
I I 0.43
I0
-2

1000 1200 1400 Age 600 1800 1000









2000 2200 2400 Age (ka) 2600 2800 3000
0.42






) 8 0.415$ 0




e 'iu, 22 V Vl' 041 wt





3000 3200 3400 Age (ka) 3600 3800 4t000
= ,'oo.4 03






0 0.394 .
0.38




4000 4200 4400 Age ,, 4600 4800 5000




I ir in r w t
Il.i, ,'
4o ,, '45 R










30 0.38





5000 5200 5400 Age (ka) 5600 5800 6000

Figure 2-12. The astronomical solution for obliquity (Laskar et al., 1993) compared with tuned

L* reflectance data from Site 1208 for the 1-6 Ma interval. The reflectance data
filtered using a band-pass filter centered on 41kyrs is shown in the lower part of each
frame. Black indicates normal polarity and white reverse polarity. Heavy line on L*
reflectance data indicates intervals where the cyclicity is best developed.
Figue 212. he stroomial sluton fr oliquty Laskr e al. 193) cmpaed 00.43ne
Q 70etnedt fo ie10 o he16M nevl Terfetnedt
ct 60 rn adpssfle etee n4kysi hw i h oe pr fec
0 'jilt'0. 4 j 2ml oart adwht rvrs olrty eaylieonL
40- 0.39- idcts nevaswee h yliiyisbs evlpd














Site 1209




1 10 Ag k 60 20 S
I _.0.38 'I .



1200 1400 Age (ka) 1600 1800 2000


9)1 I 1,.43


20M 2200 2400 Age ki 260 2800 3000




-- v4




3000 3200 3400 Age (ka) 3600 3_o 400


3.42 R- -

0.39 ...
0.38






0.43



0.38


., "I. '" l -. .. "

S.3a )
0 38,


4200


4400 Age (ka) 4600


_Ij V V


'V V


5000 5200

960 -





-8I
670 6200'7




600) 6200


I I ~0


1 ..
" '* v- !. -' '. .
\; o.3- ^ ^ -. -


5400 Age (ka) 5600


/t


6400 Age (ka) 66o0


6800


o 43
..
,' I ,

, 4'8
o.s' ,. =.


7000


Figure 2-13. The astronomical solution for obliquity (Laskar et al., 1993) compared with tuned
L* reflectance data from Site 1209 for the 1-7 Ma interval. The reflectance data
filtered using a band-pass filter centered on 41kyrs is shown in the lower part of each
frame. Black indicates normal polarity and white reverse polarity. Heavy line on L*
reflectance data indicates intervals where the cyclicity is best developed.


- TX r 'F '
"I,.-i

'a 'I' () *Q i. il,1'V
-'-F >/1y~xN~~- ~ -


1"- -

'7'

0
-4
_R _


<
(
(
(


--












Site 1210
0.43

0 ,,0.43
40 4 4/ A 0





S1000 1200 1400 Age (ka) 1600 18oo 20o
90 "0.43

0.4 4
4- 039



2000 2200 2400 Age (ka) 2600 2800 low


0 1 t 1 1 0.43
s io- ''' o.0
041
4- 0_n38


3000 3200 3400 Age (ka) 3600 3800 4N0X


90 043

09 0.4l

2 Wi 0.38
:4,

4000 4200 4400 AgeC II 46 480W 50


Figure 2-14. The astronomical solution for obliquity (Laskar et al.,, 1993) compared with tuned
L* reflectance data from Site 1210 for the 1-5 Ma interval. The reflectance data
filtered using a band-pass filter centered on 41kyrs is shown in the lower part of each
frame. Black indicates normal polarity and white reverse polarity. Heavy line on L*
reflectance data indicates intervals where the cyclicity is best developed.












Site 1211




i^6o' A ,'
80 0 43








04 0.38


2000 2200 2400 A Li 2600 2800 3000
00 9
IV 0. '0.38






-4




3000 200 3400 .. 3600 3800 4000






0 I4
-4D







L4 0.38 E


4000 4200 4400 L 4600 4SC 500


Figure 2-15. The astronomical solution for obliquity (Laskar et al., 1993) compared with tuned
L* reflectance data from Site 1211 for the 1-5 Ma interval. The reflectance data
filtered using a band-pass filter centered on 41 kyrs is shown in the lower part of each
frame. Black indicates normal polarity and white reverse polarity. Heavy line on L*
reflectance data indicates intervals where the cyclicity is best developed.



























0 0.02 0.04 0.06 0.08 0.
frequency (k.y. )


Site 1208 2-3 Ma











S /Ii

: -k~j ^V ^v w


0 0 02 0.04 0.06
frequency (k.y. -1)


0.08 0. 1


0.02 0.04 0.06
frequency (k.y. -1)


0 0.02 0.04 0.06
frequency (k.y. 1)


0 0.02 0.04 0.06 0.08 0.1 0 0.02 0.04 0.06 0.08 0.1
frequency (k.y. ') frequency (k.y. 1)


Figure 2-16. Cross-spectral analysis from a) Site 1207, b) Site 1208 c) Site 1209 d) Site 1210 and
e) Site 1211. Power spectra for the tuned reflectance data (black) and for the
astronomical solution for eccentricity and obliquity (red). Coherence values between
the astronomical solutions and reflectance data are shown below.


0.08 0.1


0.08 0.1









CHAPTER 3
INTEGRATED NEOGENE MAGNETIC, CYCLE AND BIO- STRATIGRAPHY FROM ODP
SITE 1208 (SHATSKY RISE, PACIFIC OCEAN)

Introduction

ODP Site 1208 was drilled in 2001 on Shatsky Rise, a large igneous province in the NW

Pacific Ocean. A single hole drilled at the site has produced a magnetic polarity stratigraphy for

the 0-12 Ma interval. Sedimentation rates decrease from 4-5 cm/kyr in the Brunhes and

Matuyama chrons to less than 1 cm/kyr at the base of the studied section. A revised planktonic

foraminifer biostratigraphic zonation has been developed for the NW Pacific Ocean using the

seventeen most isochronous foraminifer datums. This scheme has been integrated with

nannofossil events, and with the magnetic stratigraphy. Cycles in the reflectance (L*) can be

matched to astronomic solution for obliquity allowing astronomic calibration of polarity chron

boundaries, and planktonic foraminifer and calcareous nannofossil datums. Astronomic ages for

polarity chron boundaries are consistent with the ATNTS2004 timescale (Lourens et al., 2004) in

the 1-5.2 Ma interval, however, between 5.2 and 6.2 Ma, astronomic ages from Site 1208 differ

significantly (by 300 kyr) from the ATNTS2004 timescale.

As polarity reversals can be considered globally synchronous, the integration of polarity

chron boundaries and biostratigraphies has become a powerful means of calibrating

biostratigraphic zonations, and determining synchroneity of biostratigraphic events (see

Berggren et al., 1995a,b). Early work on Neogene foraminifer biostratigraphy in the North

Pacific Ocean on Deep Sea Drilling Project (DSDP) Sites 173, 296 and 310 (Keller, 1979a,b,c)

was augmented by foraminifer and nannofossil work on ODP Leg 138 in the eastern equatorial

Pacific (Raffi and Flores, 1995; Shackleton et al., 1995a). ODP Leg 138 biostratigraphies were

integrated into well-defined magnetic stratigraphies (Schneider, 1995) and cyclostratigraphies

based on gamma ray attenuation (GRA) bulk density data. Correlation of GRA bulk density









cycles to astronomical calculations for solar insolation provided robust age models for Late

Miocene to Recent sediments (Shackleton et al., 1995b). The ODP Leg 138 age models were

among the first astrochronologies developed for the Late Miocene to Quaternary, and hence the

biostratigraphies generated from ODP Leg 138 sites were rather precisely calibrated. Berggren

et al. (1995a,b) incorporated ages and bio-magnetostratigraphic data from ODP Leg 138 into

their review of bio-magnetostratigraphic correlations for the Cenozoic and Quaternary.

ODP Site 1208 offers the opportunity to refine biomagnetostratigraphic correlations for the

late Neogene. The attributes of ODP Site 1208 include: good preservation of foraminifers and

calcareous nannofossils, relatively high sedimentation rates compared to ODP Leg 138 sites, and

a robust age model based on magnetic polarity stratigraphy and correlation of reflectance data to

astronomical solutions.

The study of ODP Site 1208 is a continuation of the work presented in Chapter 2, which

deals largely with ODP Leg 198 sites other than Site 1208. An initial astrochronology for ODP

Site 1208 (Chapter 2, Evans et al., 2005) was based on correlation of the shipboard reflectance

(L*) data (Shipboard Scientific Party, 2002a) to the astronomical solution for obliquity of Laskar

et al. (1993). Here, we update the Site 1208 astrochronology using the new astronomic solutions

of Laskar et al. (2004), present the Site 1208 magnetostratigraphy, foraminiferal and nannofossil

biostratigraphy, and link these stratigraphies to the new astrochronology. The recalibration of

the Site 1208 age model makes little difference to the chronology presented in the previous

chapter (and in Evans et al., 2005) because the astronomic solutions in the 0-12 Ma interval do

not change significantly in Laskar's two calculations (Laskar et al., 1993; 2004).

Today, Shatsky Rise (Figure 3-1) lies in a subtropical water mass toward the north end of a

warm-water mass known as the Kuroshio Extension Current (Shipboard Scientific Party, 2002b).









North of the Northern High of Shatsky Rise (Figure 3-1) lies a significant front, a transition

region between subtropical and subarctic water masses. The transition zone waters are derived

from off the coast of northern Japan, where the cold, nutrient-rich Oyashio Current mixes with

the warm, nutrient-poor Kuroshio Extension Current. Middle Miocene calcareous plankton

assemblages are rather uniform and diverse across Shatsky Rise and display warm, subtropical

affinities. Since the Late Miocene, however, a faunal and floral gradient has been established

across Shatsky Rise (Shipboard Scientific Party, 2002b). Calcareous plankton assemblages

progressively loose their warm-water taxa along a traverse from south to north across the

Shatsky Rise. At Sites 1207 and 1208 (Figure 3-1), there is a marked decrease in diversity in

assemblages that assume temperate (occasionally cold-temperate) affinities, relative to sites

further south. The changes in calcareous plankton assemblages are paralleled by a progressive

decrease in calcareous preservation from north to south (Shipboard Scientific Party, 2002b).

One of the most noticeable features of the upper Miocene through Pleistocene sections

recovered at Shatsky Rise is the decimeter- to meter-scale cycles between darker and lighter

lithologies. The darker-colored intervals, in general, contain larger amounts of well-preserved

biosiliceous material, and contain calcareous plankton assemblages that have cold-water

affinities and have undergone relatively enhanced dissolution. Calcareous plankton preservation

is enhanced in the light-colored layers that are poorer in diatoms and represent warmer-water

intervals when Site 1208 was located in a subtropical water mass, similar to the situation at Site

1208 today and for the Southern High through most of the Neogene (Shipboard Scientific Party,

2002b).









Site Location and Lithology

ODP Site 1208 is located in 3346m of water on the Central High of Shatsky Rise (Figure

3-1). The Central High of the Rise had not been drilled prior to ODP Leg 198, and the

sedimentary sequence at the site revealed -260 m of Upper Miocene to Recent sediments with

-60m of more condensed Lower and Middle Miocene below. A total of 314.17 m of Neogene

age sediment was recovered at the site with an average recovery of 95%. The Upper Miocene to

Recent section is composed of nannofossil ooze and nannofossil clay with diatoms and

radiolarians, and an average carbonate content of 53% (Shipboard Scientific Party, 2002b).

Since 3 Ma, the average sedimentation rates were 4.2 cm/kyr. Prior to 3 Ma, sedimentation rates

decrease progressively reaching 1 cm/kyr at -8 Ma. The character of the seismic reflection

record at the site, along with the relatively high sedimentation rate that prevailed during the

Pliocene-Pleistocene, suggests that the stratified lens of sediment at the site constitutes a drift

deposit formed by current redistribution of sediment that settled on the Central High (Shipboard

Scientific Party, 2002b). The sediment drift deposits at Site 1208 are somewhat similar to those

drilled along the Meiji Seamount during ODP Leg 145 in that both sections comprise fine-

grained sediment devoid of sedimentary structures other than bioturbation (Rea et al., 1993).

Magnetic Stratigraphy

Magnetic measurements on half cores from Site 1208, using the shipboard pass-through

magnetometer, revealed an unambiguous magnetic stratigraphy, ranging in age from Recent to

Upper Miocene (Figures 3-2, 3-3 and 3-4). The shipboard data are based on a single

demagnetization step (20 mT). This abbreviated treatment was necessary to preserve the

sediment magnetization for later shore-based study, and to maintain core-flow through the

shipboard core laboratory during the cruise. These shipboard data are supported using discrete

sample cubes (7cm3) collected from the working halves of cores, which were measured at the









University of Florida. The discrete sample cubes were AF demagnetized in 5mT increments up

to peak fields of 80 mT. A steep drilling related overprint was removed by 20 mT peak field

(Figure 3-5), and the primary magnetization was defined using the standard least squares method

(Kirschvink, 1980), giving low maximum angular deviation (MAD) values indicating well

defined component magnetizations (Figures 3-2, 3-3 and 3-4).

An initial age model and initial estimate of interval sedimentation rates were calculated

using the magnetostratigraphy and the geomagnetic polarity timescale (GPTS) of Cande and

Kent (1995) (Figure 3-6a). Sedimentation rates decrease down section averaging 4-5 cm/kyr in

the Pleistocene, 3.5-4 cm/kyr in the Pliocene and 1-2 cm/kyr in the Miocene. The duration of the

Reunion subchron given in the Cande and Kent (1995) GPTS (10 kyr) causes a large increase in

the sedimentation rates in the polarity zone correlative to the Reunion subchron (Figure 3-2).

Using a revised age and duration for the R6union subchron (Channell et al., 2003), sedimentation

rates in the polarity zone correlative to the Reunion subchron are reduced to ~4 cm/kyr in

keeping with surrounding sedimentation rates.

Numerous excursions can be identified in the shipboard magnetic stratigraphy particularly

in the Matuyama Chron, one of which (at 103 mbsf) is confirmed by a single discrete sample

corresponding to an age of 2.283 Ma. Channell et al., (2002) identified seven excursions in the

Matuyama chron at ODP Site 983 in the North Atlantic that have been labeled: Santa Rosa (932

ka), Clr. n. r (1048 ka), Punaruu (1115 ka), Bjorn (1255 ka), Gardar (1472-1480 ka), Gilsa

(1567-1575 ka), and C2r.lr.ln (1977 ka).

At ODP Site 1208, shallow inclinations are seen in shipboard data that appear to be

correlative to Santa Rosa (950 ka), Punaruu (1123 ka), Gardar (1450 ka), Gilsa (1522 ka), and

C2r. Ir. In (1976 ka) (Figure 3-2). An interval of shallow inclination at 896 ka at Site 1208 may









correspond to the Kamikatsura excursion that originates from the work of Maenaka (1983).

Three intervals of shallow inclination are also noted in the Brunhes chron with ages of 134, 193

and 262 ka close to the published ages for the Blake excursion (120 ka), Iceland Basin excursion

(189 ka) and 8a (260 ka) of Lund et al. (2001) (Figure 3-2). One potential excursion is noted in

the Gauss chron (Figure 3-3) and three potential excursions in the Gilbert chron (Figure 3-5).

The ages for the Santa Rosa, Blake, Iceland Basin and 8a are calculated by assuming constant

sedimentation rates within the Brunhes and subchron Cir. Ir. Ages for other excursions are at

Site 1208 are calculated from the astronomic age model described below. All the postulated

excursions in the Site 1208 record should be regarded with some caution as they are based on a

single demagnetization step (20 mT peak field) from shipboard data.

Cycle Stratigraphy

Shipboard gamma ray attenuation bulk density data and L* reflectance data (Shipboard

Scientific Party, 2002), show a prominent cyclicity in the 1-6 Ma interval, that, based on the

initial age model has a period close to 41 kyr (see Chapter 2, Evans et al., 2005). Using the

astronomical solutions of the Laskar et al. (2004), the L* reflectance data was tuned to obliquity

by matching the L* output of a filter centered on 41 kyr to the orbital solution for obliquity

(Figure 3-7). The resulting interval sedimentation rates for the 1-6 Ma interval are given in

Figure 3-6b. The tuned ages for reversal boundaries, based on this match, are given in Table 3-2.

The astronomical calibration of Site 1208 presented here is a recalibration of the astronomical

timescale of Evans et al., (2005) using the updated astronomical solutions of Laskar et al. (2004).

The recalibration to the new astronomic solutions resulted in little change to the astronomic ages

for Site 1208 relative to those given by Evans et al. (2005). Comparison of polarity reversal ages

with other timescales, and with results from IODP Site U1313 (Chapter 6), indicates close









agreement with differences < -60 kyrs between 2.6 Ma and 5 Ma (Table 3-2). Beyond 5.2 Ma,

the differences with respect to other timescales increase to over 100 kyr. In the 5.5-6 Ma interval,

the polarity reversal ages from Site 1208 are closest to those of Shackleton et al. (1995b) from

ODP Leg 138 (equatorial Pacific). The largest discrepancy beyond 5.2 Ma is with ATNTS2004

timescale (Lourens et al., 2004) where the difference in ages is 300 kyrs. The ATNTS2004

timescale uses the work of Hilgen et al. (1995) from the Mediterranean in this interval.

Calcareous Nannofossils

Calcareous nannofossils were semi-quantitatively analyzed using smear slides and

standard light microscope techniques (Bown and Young, 1998). The following abundance and

preservation categories were used: Species abundance: abundant: >10 specimens per field of

view (FOV), common: 1-10 specimens per FOV, few: 1 specimen per 2-10 FOV, rare: 1

specimen per 11-100 FOV. Total nannofossil abundance: abundant: >10%, common: 1%-10%,

few: 0.1%-1%, rare: <0.1%, barren and questionable occurrence. Nannofossil preservation:

good, moderate, poor (See range chart of Bown, 2005). All core catcher samples were examined

and -60 other samples collected through the Late Miocene to Recent section. Biostratigraphy is

described with reference to the zonal scheme of Bukry (1973, 1975; zonal code numbers CN and

CP added and modified by Okada and Bukry, 1980) for Cenozoic calcareous nannofossil

biostratigraphy.

The middle Miocene-Holocene section yielded a beautiful succession of rich and abundant

nannofossil assemblages. Preservation improved up-section but was also dependent upon which

part of the light/dark sedimentary cycle was sampled. The darker, diatom-rich intervals yielded

more poorly preserved nannofossil assemblages (Shipboard Scientific Party, 2002). The Neogene

nannofossil biostratigraphy indicates a relatively complete stratigraphy for the Pliocene-

Pleistocene (Figure 3-8) and Miocene (Figure 3-9), with all nannofossil zones from CN5 through









CN15 identified by their primary zonal fossils (Figure 3-10). Calcareous nannofossil range charts

are shown in Bown (2005). Zones CN1-CN5 could not be easily distinguished because of the

absence of the marker species Sphenolithus belemnos, Helicosphaera ampliaperta, and

Discoaster kugleri. In addition, a number of CN subzones could not be recognized due to the

absence ofD. kugleri (Subzone CN5b), Discoaster loeblichii, Discoaster neorectus (Subzone

CN8b), and Amaurolithus amplificus (subdivisions within Zone CN9) and an anomalously low

last occurrence (LO) of Triquetrorhabdulus rugosus (Subzone CN10b) (Bown, 2005).

The astronomically calibrated ages of Pliocene to Quaternary calcareous nannofossil

datums from Site 1208 (Table 3-3) are generally consistent with ages from Berggren et al.

(1995b), that are based largely on work from the Mediterranean (Rio et al., 1990). The LO of

Discoaster brouweri, however, differs significantly from Berggren et al. (1995b) in both age and

correlative polarity chron (Table 3-3). The age of 1.95 Ma given by Berggren et al. (1995b),

correlative to the onset of the Olduvai subchron, is based on correlation to Deep Sea Drilling

Project (DSDP) Site 606 in the North Atlantic (Backman and Pestiaux, 1987).

At Site 1208, the FO of Discoaster berggrenii in the Late Miocene is 0.5 Myrs younger

than the age reported in Berggren et al. (1995a). This age is based on correlation to polarity

chron C4r.2r from ODP Leg 138. The age is more consistent with that seen at DSDP Site 608

where the datum is correlated to polarity chron C4n (Ruddiman et al., 1987). The FO of

Discoaster hamatus is a controversial datum (Berggren et al., 1995a) that has very inconsistent

correlation to polarity chrons regardless of latitude. In ODP Leg 138 sites, it is correlative to

subchron C5n.2n, as at Site 1208. The FO of Catinaster coalitus is another controversial datum

that, at ODP Site 1208, is correlated to subchron C5n.2n similar to the correlation at ODP Leg

138 sites. Berggren et al. (1995a) give an age of 10.8 Ma for the FO of Coccolithus









miopelagicus, 200 kyrs younger than the age from Site 1208 (Table 3-3), however the correlation

of this datum to polarity subchron C5r. r at Site 1208 is consistent with the correlation at DSDP

Site 608.

Planktonic Foraminifera

158 samples were analyzed for planktonic foraminifers at -1.5 m intervals, together with

core-catcher samples (from the base of each core) collected shipboard from the 320-m-thick

upper Neogene section at ODP Site 1208. The samples were soaked in a slightly basic solution,

shaken, washed over a 63 [tm sieve and dried at 600C. Specimens of planktic and benthic

foraminifers were picked from the >125 [im fraction. Specimens of all recognizable planktic

species were identified following the classic taxonomies of Kennett and Srinivasan (1983), Bolli

and Saunders (1985), Jenkins (1985), and laccarino (1985). Shipboard and shore-based

occurrence tables were combined to determine a planktic foraminifer biostratigraphy.

Occurrence estimates were based on the following percentages: Rare=l%, rare to few=3%,

few=5%, few to common=8%, common=10%, common to abundant=15%, abundant =>20%,

Planktonic foraminiferal abundance varies from abundant to common through the

Pleistocene and upper Pliocene but declines in the Miocene to few to rare relative to siliceous

microfossils and clay. Temperate-water species dominate many of the Neogene planktonic

foraminiferal assemblages at Site 1208 (Shipboard Scientific Party, 2002b).

The magnetostratigraphically-interpolated ages for many foraminiferal datums on Shatsky

Rise differed significantly from those reported from the southwest Pacific, due to regional

migration patterns. Application of zonal schemes proposed for the southwest Pacific (Jenkins,

1985) and the mid-latitudes were complicated by unexpected changes in the sequence of

foraminiferal datums observed at Shatsky Rise. A revised temperate foraminifer biostratigraphy









for the late Neogene uses seventeen of the most isochronous foraminiferal datums at Shatsky

Rise as zonal markers (shown in Figure 3-11).

Discrepancies between published ages for planktonic foraminifer datums (Berggren et al.,

1995a,b; Lourens et al., 2004) and those identified at Site 1208 are large in some cases. (Tables

3-4 and 3-5). The majority of the magneto-biostratigraphic correlations used in Berggren et al.,

(1995a,b) and Lourens et al. (2004) are from the Mediterranean (Hilgen, 1990), South Atlantic

(Hodell and Kennett, 1987) or South Pacific (Srinivasan and Sinha, 1993). The comparison of

the ages of the planktic foraminifer datums is affected by regional differences and varying zonal

schemes (Tables 3-4 and 3-5).

The LO of Gr. tosaensis at ODP Site 1208 is at 0.292 Ma, significantly younger than the

age of 0.65 Ma given by Berggren (1995b). This datum is taken from the work of Berggren et al.

(1985) and Srinivasan and Sinha (1993) from the southern Pacific and Indian Oceans. The LO

datum of Gr. punticulata has an age of 1.882 Ma at Site 1208, however, an age of 2.41 Ma was

obtained at DSDP Site 607 (North Atlantic) where it is correlative to polarity subchron C2An.2n.

The FO of Gr. truncatlinoides occurs at the same stratigraphic level as the FO of Gr.

toseanis at ODP Site 1208. Following Berggren et al. (1995b), the FO of Gr. toseanis has an age

of 3.35 Ma. At Site 1208, the astronomically calibrated age of the event is 2.015 Ma. The Site

1208 age for this datum is more consistent with the astronomically calibrated age for the FO of

Gr. truncatlinoides of 2.39 Ma from ODP Leg 138 in the eastern equatorial Pacific (Shackleton

et al., 1995a). The LO of Gr. margarita was assigned an age of 3.85 Ma in the ATNTS2004

(Lourens et al., 2004) from ODP Sites 925 and 926 from Ceara Rise. The astronomic age for the

datum at Site 1208 is 3.761 Ma.









Conclusions

ODP Site 1208 has produced a clear magnetic stratigraphy for the 0-12 Ma interval with

sedimentation rates in the Brunhes and Matuyama chrons varying in the 4-5 cm/kyr range. These

sedimentation rates are some of the highest sedimentation rates seen in this interval in pelagic

sediments from the mid- and low latitude Pacific Ocean. This anomalously high sedimentation

rate appears to be due to formation of a drift-type deposit on the Central High of Shatsky Rise.

The relatively high sedimentation rates have allowed identification of polarity excursions in the

Matuyama Chron that have not been previously identified in sediments from the Pacific Ocean.

It is important to stress that these excursions are identified in shipboard pass-through magnetic

data, and are not based on identification of magnetization components. For this reason, the

ratification of these excursional directions must await further (u-channel) studies of these

sediments.

Reflectance (L*) cycles identified in the sediments have allowed astronomic calibration of

reversal boundaries and biostratigraphic datums, by correlation ofL* reflectance data to the

astronomic solution for obliquity (Laskar et al., 2004). Calcareous nannofossil biostratigraphy is

largely consistent with the most recent review of bio-magnetostratigraphic correlations for this

time interval (Berggren et al., 1995a, b). Based on the correlation of planktonic foraminifer

datums to the magnetic stratigraphy at Site 1208, a new planktonic foraminifer zonation for the

northwest Pacific Ocean has been developed that can be precisely correlated to polarity chrons

and astronomically calibrated ages.











Table 3-1. Depths of reversal boundaries from ODP Site 1208. Chrons are labeled according to
Cande and Kent (1992). Ages for polarity chrons are from Cande and Kent (1995)
and Channell et al., (2003).


Chron Ma (CK95) mbsf

Cln 0
base 0.78
Clr.ln 0.99
base 1.07
Clr.2r.ln 1.201
base 1.211
C2n 1.77
base 1.95
C2r.ln* 2.115
base* 2.153 1
C2An.ln 2.581 1
base 3.04
C2An.2n 3.11 1
base 3.22 1
C2An.3n 3.33 1
base 3.58 1
C3n.ln 4.18
base 4.29 1
C3n.2n 4.48 1
base 4.62 1
C3n.3n 4.8 1
base 4.89 1
C3n.4n 4.98 1
base 5.23 2
C3An.ln 5.894 2
base 6.137 2
C3An.2n 6.269 2
base 6.567 2
C3Bn 6.935 2
base 7.091 2
* age from Channell et al. ,2'" .1 ')


Chron Ma (CK95) mbsf


0 C3Br.ln
42.92 base
52.57 C3Br.2n
55.85 base
61.18 C4n. ln
61.67 base
85.01 C4n.2n
92.81 base
99.86 C4r.ln
01.01 base
19.45 C4An
137.8 base
40.64 C4Ar.ln
44.58 base
47.76 C4Ar.2n
56.88 base
172.4 C5n. ln
76.47 base
82.51 C5n.2n
85.46 base
89.28 C5r.ln
91.01 base
94.09 C5r.2n
00.04 base
16.47 C5An.ln
21.51 base
22.25 C5An.2n
31.39 base
35.31
40.34


6.946
6.981
7.153
7.187
7.245
7.376
7.464
7.892
8.047
8.079
8.529
8.861
9.069
9.146
9.428
9.491
9.592
9.735
9.777
10.834
10.94
10.989
11.378
11.434
11.852
12
12.108
12.333


240.33
241.08
241.83
242.95
250.78
251.71
252.46
256
260.66
262.53
264.02
265.69
268.12
269.05
271.85

282.1
287.14
287.69
290.49
291.42
292.17
294.03
298.69
299.63










Table 3-2. Astronomically calibrated ages for reversal boundaries from ODP Site 1208 compared
to ATNTS2004 (Lourens et al., 2004), Cande and Kent (1995), IODP Site U1313
(Evans et al., in preparation, Chapter 6), Hilgen et al., (1995) and ODP Leg 138
(Shackleton et al., 1995b). Differences between Site 1208 ages and published ages are
given in parentheses.


Chron 1208 CK95 (Ma) Hilgen et al. ATNTS ODP Leg 138 IODP Site
tuned age (1995) (Ma) 2004 (Ma) U1313
(Ma) Chapter 6
Cln
base 0.780
Clr.ln 0.990
base 1.062 1.070 (-0.008) 1.072 (0.01)
Clr.2r.ln 1.158 1.201 (0.043) 1.173 (0.015)
base 1.167 1.211 (0.044) 1.185 (0.018)
C2n 1.763 1.770 (-0.007) 1.785 (0.022) 1.778 (0.015)
base 1.944 1.950 (-0.006) 1.942 (-0.002) 1.945 (0.001)
C2r.ln 2.204 2.140 (0.064) 2.129 (-0.075) 2.128 (-0.076)
base 2.214 2.150 (0.064) 2.149 (-0.065) 2.148 (-0.066)
C2An.ln 2.616 2.581 (0.035) 2.582 (-0.34) 2.581 (-0.035) 2.600 (-0.016) 2.616 (0)
base 3.048 3.040 (0.008) 3.032 (0.016) 3.032 (-0.016) 3.046 (-0.002) 3.074 (0.026)
C2An.2n 3.091 3.110 (-0.019) 3.116(0.025) 3.116(0.025) 3.131 (0.04) 3.153(0.062)
base 3.207 3.220 (-0.013) 3.207 (0) 3.207(0) 3.233 (0.026) 3.268 (0.061)
C2An.3n 3.350 3.330 (0.020) 3.330 (-0.02) 3.330 (-0.02) 3.331 (-0.019) 3.346 (-0.004)
base 3.584 3.580 (0.004) 3.569 (-0.015) 3.596 (0.012) 3.594 (0.01) 3.549 (-0.035)
C3n. n 4.164 4.180 (-0.016) 4.188 (0.024) 4.187 (0.023) 4.199 (0.035) 4.144 (-0.02)
base 4.307 4.290 (0.017) 4.300 (-0.010) 4.300 (-0.007) 4.316 (0.009) 4.277 (-0.03)
C3n.2n 4.484 4.480 (0.004) 4.493 (+0.009) 4.493 (0.009) 4.479 (-0.005) 4.500 (0.016)
base 4.601 4.620 (-0.019) 4.632 (0.031) 4.631 (0.03) 4.623 (0.022) 4.631 (0.03)
C3n.3n 4.785 4.800 (-0.015) 4.799 (0.014) 4.799 (0.051) 4.781 (-0.004) 4.760 (-0.025)
base 4.897 4.890 (0.007) 4.879 (-0.018) 4.896 (-0.001) 4.878 (-0.019) 4.889 (-0.008)
C3n.4n 4.987 4.980 (0.007) 4.998 (0.011) 4.997 (0.01) 4.977 (-0.01) 5.009 (0.022)
base 5.182 5.230 (-0.048) 5.236 (0.054) 5.235 (0.053) 5.232 (0.05) 5.273 (0.091)
C3An.ln 5.735 5.894 (0.159) 5.952 (0.217) 6.033 (0.298) 5.875 (0.14)
base 5.955 6.137 (0.182) 6.214 (0.259) 6.252 (0.297) 6.122 (0.167)











Table 3-3. Nannofossil datums for ODP Site 1208 (Bown, 2005). Ages for the datums are
interpolated from the magnetic stratigraphy (this work) and the correlative polarity
chron is given. The datums are compared to ages given by Berggren et al. (1995a, b).
Tuned ages for the datums are compared to ATNTS2004 (Lourens et al., 2004) and
ODP Leg 138 ages for nannofossil datums only (Raffi and Flores, 1995; Shackleton
et al., 1995a).




1208 Berggren et ODP 1208 ATNTS
Datum Depth
Datum ( ) Mag. strat. Chron al. 1995a b Leg Tuned age
Age (Ma) Site 1208 (Ma) chron 138 age (Ma) (Ma)


FO Emiliania huxleyi
LO P. lacunosa
FO G. omega
FO G. caribbeanica
LO D. brouweri
LO D. pentaradiatus

LO D. surculus

LO D. tamalis
LO Large
Reticulofenestra
FO D. tamalis
LO Sphenolithus
LO Amaurolithus
FO D. asymmetricus
FO C. cristatus
LO D. quinqueramus
FO Amaurolithus
FO D. quinqueramus
FO D. berggrenii
FO D. hamatus
FO C. calyculus
FO D. hamatus
FO C. coalitus
LO C. miopelagicus
LO C. premacintyrei
LO C. floridanus


14.24
30.30
43.11
87.90
100.16
116.40

119.08

128.70
163.90

166.66
166.66
168.88
168.88
187.90
207.00
235.52
250.80
250.80
265.94
270.10
274.20
279.70
285.06
295.41
295.41


0.258
0.551
0.784
1.837
2.143
2.510


Cln
Cln
Clr.lr
C2n
C2r. In
C2r.2r


2.572 C2r.2r

2.812 C2An.ln
3.851 C2Ar


3.65
3.65
4.56
4.56
4.735
5.551
6.941
8.075
8.075
9.586
9.792
10.125
10.699
11.009
13.19
13.19


C2Ar
C2Ar
C2Ar
C2Ar
C3n.2r
C3r
C3Bn
C4r.lr
C4r. lr
C4Ar.2n
C5n. ln
C5n.2n
C5n.2n
C5r. lr
C5An.lr
C5An.lr


0.26 0.26
0.46 0.46


1.95 Olduvai
2.46-2.56 M/G
boundary
2.55-2.59 M/G
boundary
2.78 top Gauss


3.6 bas


4.2 top


1.841
1.96 2.146
2.52 2.499


2.63 2.556 2.52


2.78 2.802
3.833


3.95
e Gauss 3.66 3.95
4.03
Cochiti 4.13 4.03
4.750
5.6 C3r 5.55 5.472


8.6 C4r.2r
9.4 C4Ar.2r

10.7
10.9
10.8


8.45



10.38



12.65
13.19


10.79
10.55
10.89
11.02
11.21
13.33


0.29
0.44



2.06
2.39


2.80












Table 3-4. Plio-Pleistocene foraminfer datums, with depths, correlative polarity chron, tuned age
and compared to Berggren et al. (1995a, b) and ATNTS 2004 (Lourens et al., 2004)
from ODP Legs 138 and 111.


Event


LO Gr. crassula
LO Gr. tosaensis
LO Gs. bulloideus
LO B. praedigitata
LO Gt. woodi
LO Gs. bollii
LO Gs. obliquus
LO N. acostaensis
LO N humerosa
LO Gt. decoraperta
LO Gr. puncticulata, FO Gs. tenellus,
FO Gs. elongatus
FO Gr. hirsuta
FO Gr. toseansis, LO Gr. cibaoensis, FO Gr.
truncatulinoides
LO Pu. primalis, LO Gr. limbata FO Ga. parkerae
FO Pu. obliquiloculata, FOB. digitata
LO Gq. venezuelana
LO Ss. paenedehiscens, LO Gt. apertura
LO N. "dupac", FO Ge. siphonifera
LO Gr. pseudomiocenica
LO Gr. juanai, FO Gr. bermudezi
LO Gr. plesiotumida, LO Gs. extremus,
LO Gs. triloba
FO Gr. limbata
LO Gq. conglomerate, LO Gr. sphericomiozea
FO Pu. primalis
LO Gr. inflata
FO Gt. rubescens
FO Gr. puncticulata, LO Gr. conoidea
FO Sa. dehiscens
LO Ge. pseudobesa
FO Ge. calida, FO Gr. crassula, LO D. altispira, LO Ss.
seminulina
LOSs. kochi
LO Gr. margaritae


FO Gs. bulloideus
FO Gq. conglomerate, FO Gr. crassiformis
FO Ga. uvula, FO Gs. extremus,
LO Gr. conomiozea
FO Gg. umbilicata, FO Ge. aequilateralis
FO Gr. sphericomiozea
FO Gr. conomiozea, LO Gt. nepenthes,
FO Gr. pseudomiocenica


Depth 1208
mbsf mag strat
age
0.4 0.007
16.1 0.292
38.2 0.694
43.7 0.797
53.2 1.005
60.5 1.182
66.7 1.331
76.2 1.559
79.6 1.640
83.6 1.736
89.7 1.878


Chron
Site 1208


Cln
Cln
Cln
Clr.lr
Clr.ln
Clr.2r
Clr.2r
Clr.2r
Clr.2r
Clr.2r
C2n


92.3 1.938 C2n
95.2 2.014 C2r.lr


98.5
101.9
108.1
111.4
121.5
123.0
125.8
132.5

135.3
137.2
142.2
146.2
147.4
150.0
151.5
158.2
159.6


2.103
2.171
2.316
2.393
2.632
2.670
2.740
2.907

2.978
3.250
3.154
3.276
3.318
3.391
3.433
3.631
3.740


C2r.lr
C2r.2r
C2r.2r
C2r.2r
C2An. ln
C2An. ln
C2An. ln
C2An. ln

C2An. ln
C2An. ln
C2An.2n
C2An.2r
C2An.2r
C2An.3n
C2An.3n
C2Ar
C2Ar


161.0 3.739 C2Ar
162.4 3.793 C2Ar


165.0
172.0
178.7

182.9
186.5
190.8


3.894
4.165
4.360

4.499
4.669
4.879


Berggren et
al 95ab age
(Ma)


1208
tuned
ae


0.65 Cln


1.004
1.197
1.34
1.562
1.653
1.742
2.41 1.882

1.930
3.35C2An.2n 2.015


2.116
2.164
2.282
2.382
2.621
2.654
2.714
2.902


2.966
3.024
3.147
3.301
3.338
4.5 Nunivak 3.427
5.2 E.Gilbert 3.457
3.637
3.682


3.58 G/G
boundary


C2Ar
C2Ar
C3n.lr

C3n.2n
C3n.2r
C3n.3n


3.708
3.761

3.871
4.187
4.413


4.54
5.6 C3r 4.696
4.2 Cochiti 4.873


ATNTS
Age











Table 3-5. Miocene foraminifer datums, with depths, correlative polarity chron, tuned age and
compared to Berggren et al. (1995a, b) and ATNTS 2004 (Lourens et al., 2004) from
ODP Legs 138 and 111.


Datum


FO Gr. tumida, LO Gs. kennetti
FO Gs. bollii
FO Gs. kennetti
FO Gd. hexagona
FO N. dutertrei, FO Gs. conglobatus
FO Ss. paenedehiscens, LO Gr.
merotumida
FO Ge. pseudobesa, FO Gr. margaritae,
FO Ss. kochi, FO Gr. plesiotumida,
FO N. humerosa, FO Gr. scitula
LO Gr. miotumida c.f.
FO Gr. cibaoensis, FO N. acostaensis,
FO Gr. miotumida c.f.
LO Gq. baroemoenensis
FO Gr. juanai
FO B. praedigitata, FO Gs. obliquus,
FO N. pachyderma dextrall), (sinistral)
FO Gs. ruber, FO Gt. apertura, LO Gq.
dehiscens
FO Gr. merotumida
LO Gr. praemenardii
FO N. "dupac"
LO Gt. druryi
LO Ss. disjuncta
FO Gt. decoraperta
FO Gr. miozea
LO Gr. mayeri
LO N. continuosa
LO Cs. parvulus, LO Gr. panda

FO Gt. nepenthes, FO Gr. mayeri, FO Gr.
menardii
FO Gt. drurvi


Depth 1208 mag
(mbsf) strat age
(Ma)


203.1
209.4
211.3
214.3
217.3
224.5

227.3


233.6
240.2

245.9
251.1
255.6

263.1

270.5
272.6
276.8
282.3
284.3
285.0
291.9
294.1
295.2
296.8

298.9

299.7


5.354
5.608
5.685
5.806
5.934
6.342


Chron Berggren
Site 1208 et al.
(1995b)
C3r 5.6 C3r
C3r
C3r
C3r
C3An.ln
C3An.2n


6.434 C3An.2n 6.0 C3An


C3Ar
C3Bn

C4n.2n
C4r.lr
C4r.lr


7.8 C4n.2n


9.4 C4Ar.ln


C5n. n
C5n.2n
C5n.2n
C5r.lr
C5r. lr
C5r.lr
C5r.3r
C5An.lr
C5An.lr
C5An.lr


12.3 C5An.2n

12.3


11.8
C5r.3r
11.8
C5r.3r


1208
tuned
age
5.327
5.591
5.675
5.816
5.961


ATNTS
2004

5.57




6.2


11.49


11.63


























Northern High
1207


Central High


1208 1as" -0













1 5 and Southern High
S121 He3'os 1 asin




r. -- _i 1209



306 and S -
o1214 j 305 and Soulhrn High
1213 1211


Opn Rise sea-ounts
Olin Rise seamounts


Pacific Ocean


3000 mbsl
4000 misI
5000 mbsl
SQOW mtSI


155'E 160 165' 170'


Figure 3-1. Bathymetric map showing the location of Shatsky Rise in the Pacific Ocean and a
larger map of Shatsky Rise showing the Sites drilled on ODP Leg 198 including Site
1208 on the Central High of the Rise (after Bown, 2005).


40 -
N


o


ji















0 I



LI




n--
UL
20 :'
E:
77


I : I I


1208A

LII


u'


40 k


60


GLL
B,
Ii


- -pi


'2
I-






,,;,,-


100 I ~- I l
0 50 100 150 200 250 300 350 -80
iL n.. Ill iI.,i i l


Blake

Iceland Basin

















n- n


0I


'It
LtI


U
,-

E-


LI
,d


I I
0 40 80
Inclination ( )


Chrons
(CK92/95)


Figure 3-2. Inclination, declination and MAD values plotted against meters below sea floor. Gray

line indicates the AF demagnetization data from the shipboard pass-through
magnetometer at the 20 mT demagnetization step. Open squares indicate data from

discrete samples. The polarity interpretation is shown in black (normal polarity) and
white (reverse polarity) and chrons are labeled according to Cande and Kent (1992,

1995). Excursions are labeled according to Channell et al. (2002) and Singer et al.
(1999). Ages for excursions are calculated from the astronomic age model for Site

1208.


0 2 4 6 8 10 12
MAD (o)











1208A


I''








-3ii


140 V


- -


160 rL
U-






ISO
1-0


'-' i


lj
1:I i

i_-i


-


-9P
u-t


3950


44-05


4938


200 ''--
0 50 100 150 200 250 300 350 -80 -40 0 40 80
Declination (o) Inclination (0)


Chrons


C2r.2r


C2An.ln


C2An.lr

C2An.2n
C2An.2r

C2An.3n


C2Ar


C3n.ln

C3n.lr

C3n.2n
C3n.2r
C3n.3n
C3n.3r

C3n.4n


-j






i































0 2 4 6 8 10 1214
MAD


Figure 3-3. Inclination, declination and MAD values plotted against meters below sea floor. Gray
line indicates data from the shipboard pass-through magnetometer at the 20 mT
demagnetization step. Open squares indicate data from discrete samples. The polarity
interpretation is shown in black (normal polarity) and white (reverse polarity) and
chrons are labeled according to Cande and Kent (1992, 1995). Excursions are labeled
according to Channell et al. (2002) and Singer et al. (1999). Ages for excursions are
calculated from the astronomic age model for Site 1208.












1208


200









220









240

-D

-0




260









280









300


100 200 300
Declination (0)


4- -
















I -


Th----


E--C










-Qr







-F3









i-I_


-40 0 40
Inclination (0)


I-
___ r


~711 --



--~-



--
7---_
c~i



~F-o


Chrons




C3r




C3An.ln _

C3An.2n




C3Ar


C3Br


C4n.2n I


C4r

C4An


C4Ar


C5n.2n


C5r


f C5An

C5Ar


0 2 4 6
MAD ()


Figure 3-4. Inclination, declination and MAD values plotted against meters below sea floor. Gray
line indicates data from the shipboard pass-through magnetometer. Open squares
indicate data from discrete samples. The polarity interpretation is shown in black
(normal polarity) and white (reverse polarity) and chrons are labeled according to
Cande and Kent (1992, 1995). Gray bar indicates indeterminate polarity.


-D---


I I I I I I II ~I I I


_P














N/Up


\S 1208A-2H-3
I 20f A- 191 -5 imbsf 7,95
mbsf = 173 40 Lo rear 0m I
S Lo Treat 0 mT Hi Treat 60 m
Hi Treat 60 mT


1208A-18H-2
mbsf= 158 57
Lo Treat 0 mT
Hi Treat 60T T





5/Un S/Dn
5/Dn

N/Up N/Up N/Up




W E
1208A-I01-5
nmbsf=7 75
Lo Treat 0 mT
Hi Trcar 60 mT
1208A-6H-5
m tmbsf=4894
W E E o Treal 0 mT
Iii Treal 60 mT

1208A-IOH-2 W
o mbsf 84 07
Lo treat i Il
Hi Ireat 60 rm
S/On S/Dn S/Dn



Figure 3-5. Orthogonal projections showing AF demagnetization data from discrete samples.

Open circles represent the vector end point projections on the vertical plane and

closed circles represent vector end point projections on the horizontal plane.
















C 2I,






0 U ,1 i I, 1 2Ix)

0 2 4 6 g 10 12

b)
141






E





1 2 3 4 5 6
Age NkL,.


Figure 3-6. a) Interval sedimentation rates (black line) and age versus depth (red line) calculated
from the magnetostratigraphic data. b) interval sedimentation rates calculated for the
tuned age model for the 1-6 Ma interval.
tuned age model for the 1-6 Ma interval.












80


4_-1
7000





4
7! a 2

-4
1000


4-

.-2_
-4__
-6_
2000 2200 2400 Age (ka) 2600 2800


5400
Age (ka)


5800


Figure 3-7. Reflectance (L*) data (black line) tuned to the astronomic solution for obliquity from

Laskar et al. (2004). Lower plots shows the output of a gaussian filter centered on the

obliquity frequency (0.024), applied to the reflectance (L*) data.


1200 1400 Age (ka) 10 1800
Age (ka)


lJ 5


-22.5
-22





2000


24 5
23.5 -
23 -
22.5
22


24. 5
24
23.5
23


6000


5000


Z/"ijrJ\~\/~\/2rvSrV"Lr~J?/\/J













0-


10-


20-


30

40-


50-


60


70-


80


90-


100 -

110-

120-


130-


140-


tt 150-
-o-
2
E
160-
I--
E 170-


180-


190


EPOCH AGE


(1)

C




0
4-J
O-
Ll









-1 77--


a)

C

(1)
U

0
U -


200 -I15.3_


NANNO DATUMS
RECOVERY PALEOMAG DEPTH (mbsf) AGE (Ma)
1-H

2-H
-14.24) FO hulyi (0.258
3-H
Cn

4-H [.3030) LO P.lacra 55 1)

5-H
(43 11) FO G omg (0 784)
6-H

7-H

Ci r
8-H

9-H

10-H
C2n 1S790) FOG cantbromrco (.141>

11-H
1- 0.16) LO a owerr (2.146)
12-H C2r

13-H
1- 16.40) LO D pentrradiatus 2.499)
S19.08) LO ?urculus (2.556,)
14-H
1- 28A.7 LO l nma le 12.6021
15-H
C2An
16-H

17H 3n

18-H
C2Ar (16390) LO Large Rersu.lofi-eneaa (3.833)

19-H 3 ,

20-H
On 2
21 -X 118 7.901 FO C .cta s(4.75)

22-X
4n


FORAM DATUMS
DEPTH (mbsf) AGE (Ma)
04) LO G. crass (007)



(1611) LO C rosaens (0 292)





(38 2) LO Gs buoloideu (0.694)
(43.7) LO &praedigitator (07971

(53.21)LO ( wood (1.004)

(6051) LQ 6 o (1.197)
(66.7) LO Gs obtiquus (1.34)

(76.2) LO N. acotaensl (1.562)
(79.6) LO humeorsa 1.653)
(8315) LO Cc decoraperra 1.742)
(89.7) LO Gr puncticu jat, FO Ga lnedus FO Gs.elongatu (1.882)
(92.13 FO Gr hirHta (1.93)
(98ss5 F Gca errk LO G, imbr, LO prrrAmo (2.116)
S(101.9) FO digitata, FO P.bl lloculata 2.164)
(106.1) LO Gq.venez~elano 2.2B2)
(1114 LO 6r aperira, LO sposenedehsrrens 12.382

(121 51 FO GesiphoniferaLON dupac (2.621)
(1233 LO Gr pseudomioremnra (2.654)
(125.81 FO Grbermudezi, LOGr.juano 2.714)

(132.5) LO Gr piesieumida. LO G. exremus, LOGs sriob2.902)
(135.3) FO Gr. lraba (2.W6)
(137.2) LO Gq.rongqoerar, LO Gr. phericmrozea (3.024
(142.2) FO P. prm6fis (3.147)
- ; -, . ,
101
Is.

S' ." 37 .' 682
(I 65.) FO Gs. buosdeui (3.871)

(1720) FO Gq ci-ngomeroa, FOC 6ir crssfotr (4.187)

(178.71 FO Ga. uva, FO Gs. extremes, LO Gr. on(omroze (4.4131
(182 9) FO G Umbircara, FO Ge aoequiaeoras (4,54)
1186.53 FO r.sptherWorfozea (4.696)
8 FO Gr, c ;nomze, LO G nepenrhes 14. 31
FOGr) pseadomeionr ca 87


Figure 3-8. Plio-Pleistocene planktonic foraminifer and calcareous nannofossil datums, core
recovery, and magnetostratigraphy, plotted against meters below the sea floor. Ages
in bold are astronomically calibrated ages from this study. Ticks indicate position of
samples taken for foraminifer analysis. Depths in mbsf of datums are given in
parentheses before the datum.

















3r K (207.00) LD D quinqueramus (5.472)


235.52) FO Amaurollthus (6.941
:36n
f3r=3r

L4n
(250,80) FO D buiquer mu- %B8075)
F1D) b2erregg 0niB.0751
C4r 2n
4An

4Ar 2 (265 94) LO D hoar us 9 586
Sr (270.10) FO C calyculI (9.792)

--n (274.20) FO l homatus (10 125)
(279.70) FO C cools (10,699)

(285 06) LO C. mopeloa CuK (11 nOg)

-3
LO C premoacianyrei(13.19)
.5An 1 95"41)LO C.torldants 1,19)

(303.10) FO R.peudoumbdicus ,7npm) (13.70)

Undef.
(314.17) LO 3. heromrphus (13.52)
'- I ;:', ....... ,


(203.1) FO Gr. tumida, LO Gs kenner (5.327)



S214.3} FO (d. heagona (5.816)
217.3) FO NdutrelrP, FO G. congobaus (5..961)


*^ r :,16,494)

(1233.6) LOGr.mrroturmdat (7.0)

(1240.2) FOGt r. i entwiFOl oostcnsrs4

( 245.9) LO Gc. boeraemonenris (7.9)

S(251.1) FO Gr.o jun(B.1)
., .r .. ,,,.. (8.7)


(263.1) FO G. ruber, FO Gt cpetura, LO G, deIs:ens (9.4)

= 1270.5) FO Gr.merotumdo (9.9)
(272.6) LO Gr raemenardt (10.0)
(276.8) FON dupac (0.21




1291.9) LO Gr. moe (i11 )

,r sl. L4,1 I )
- ', .. ,. ., .... .



3 [310.1) FOPt .glorneros Il.( 11)


.. .. ,l .r


Figure 3-9. Miocene planktonic foraminifer and calcareous nannofossil datums core recovery,

and magnetostratigraphy, plotted against meters below the sea floor (after Venti,

2006). Ages in bold are astronomically calibrated ages from this study. Ticks indicate

position of samples taken for foraminifer analysis. Depths in mbsf of datums are

given in parentheses before the datum.










NANNO ZONES
M71 B73,75,0B80
NN21 CN15 F Gf
CN14 Fa O LG.,mo ,"''.7':,"
FO G.omeaa (0.7841


7- a I CN9
Br- b b
r Z- FO Amouraith us (6.941)
C4n 2 1 a a
C4r -- FO qunqurmus (8.075)
9 C4An C NN10 CN8 a
C4Ar 2
[ -LOD.hamotus (9.586)
10- C5n 0 NN9 CN7
SN FO D.hamatus (10.125)
11 NNFO Ccoaitus (10.699)
C5r NN7 b
1 r LO Cfloridanus 13.19)
12- C5Arn CN5
C5Ar NN6 a
13 '
B --- FOReticuIofenestra(>7pm)
14- -
C5ADn NN5 CN4
15- c5Bn ---- 1 FO C.miopeCNgicus



Figure 3-10. Calcareous nannofossil biostratigraphy including the zonations of Martini (1971)
and Okada and Bukry (1980) modified from Bukry (1973, 1975). Ages in bold are
astronomically calibrated ages from this study.













W A -q:
C2 2, 2
o~all i I I ll s~. il j
-i- ^ | ^ I
-^ C


1 I s| I | I I |


ig
.0)
*5cl


.* 0 r Q "3-
E
C
ap.joui~jdua; V*<3U~ pjoyf^L
inn ^


."

aouo6DIU *J93
13


2


saltuadau 'D


vounq!9


N


u 9 a----'-------- i
LO



o 1 Ln M -I
Sa c l-c

m & | j_





* z z z z z zzz z z






am 06 .1 El

SI i uLpuBz UD EUOIPJO/ u Pl j uePW 6u,-i
8
2 jeueooiy eU0oo!j/!A
I 1 i ii i i
(5 ~pZu~saW L!o~l l~lB~aI u!Se
4B
CII~al lPV


S- -I I '


o M CN 9 in wD N 00 ) CM (4 W C
(eU)3E
Figure 3-11. A proposed biostratigraphy for the mid-latitude North Pacific uses 16 planktic
foraminifer datums to divide the late Neogene into 15 biozones. The new stratigraphy
is integrated into the Geomagnetic Polarity Timescale and compared to pre-existing
planktic foraminifer zonal schemes for temperate and tropical region, as well as to
tropical calcareous nannofossil zonations (after Venti, 2006). Abbreviations for
zonations are as follows: B69: Blow (1969) modified by Kennett and Srinivasan
(1981a, 1981b) BKSA95: Berggren et al. (1995b) J85: Jenkins (1985) SK81:
Srinivasan and Kennett (1981a) M71: Martini (1971) B73,75: Bukry (1973, 1975),
OB80 Okada and Bukry, (1980): Ages in bold are astronomically calibrated ages
from this study.


I..
Cih
WUW
MM_


NI
OnC


DC
oU
Z

I-


0
IX


U









CHAPTER 4
PALEOINTENSITY-ASSISTED CHRONOSTRATIGRAPHY OF DETRITAL LAYERS ON
THE EIRIK DRIFT (NORTH ATLANTIC) SINCE MARINE ISOTOPE STAGE 11

Introduction

The Eirik Drift drapes the top of the underlying Eirik Ridge located off the southern tip of

Greenland (McCave and Tucholke, 1986). Magnetic anomalies have not been identified directly

beneath the Eirik Ridge, although the adjacent oceanic crust in both the Irminger Basin and

Labrador Sea is associated with marine magnetic anomaly 24 of Paleocene-Eocene boundary age

(Srivastava and Tapscott, 1986). The Eirik drift is 800 km long and has been constructed by the

interaction of the southwestward flowing Western Boundary Undercurrent (WBUC) and

basement topography (Chough and Hesse, 1985). The WBUC carries water masses originating

from the Norwegian and Greenland Seas that enter the North Atlantic over the Iceland-Scotland

Ridge and Denmark Strait (McCave and Tucholke, 1986; Lucotte and Hillaire-Marcel, 1994).

The WBUC moves over, and constructs the Eirik Drift and then follows bathymetric contours

around the Labrador Basin (McCave and Tucholke, 1986).

Drilling on the Eirik Drift includes Site 646 (ODP Leg 105), and piston and gravity cores

collected during cruises by the CSS Hudson in 1990, the Marion Dufresne in 1999 and the R/V

Knorr in 2002. Seismic records used to extrapolate the sequence recovered at Site 646 indicate

that the drift has been constructed since the middle to early Pliocene (Arthur et al., 1989).

Although sedimentation on the drift sequence was more or less continuous during the Late

Pliocene and Pleistocene, sedimentation rates vary considerably with glacial/interglacial

conditions and with location on the drift.

Piston cores HU90-013-012 (water depth: 2830 m) and HU90-013-013 (water depth: 3380

m) (Figure 4-1, Table 4-1), collected in 1990 during a cruise of the CSSHudson, record the last

glacial cycle at differing water depths on the Eirik Drift (Hillaire-Marcel et al., 1994). Core









HU90-013-013 shows high sedimentation rates in the Holocene while Core HU90-013-012 has

very low Holocene sedimentation rates due to winnowing by the WBUC (Stoner et al., 1995a,

1996). Increases in magnetic concentration and grain size during the early Holocene and at the

MIS 6/5e transition in HU90-013-013, were attributed to detrital influx associated with retreat of

the Greenland Ice sheet (Stoner et al., 1995b). In core HU90-013-013, four discrete detrital

layers were identified within MIS 2 and 3 based on their magnetic properties (coarse magnetic

grain size) and relatively high percent carbonate values. Stoner et al. (1996) correlated three of

these detrital layers with Heinrich events 1, 2 and 4. Stoner et al., (1998) revised the chronology

for core HU90-013-013 by correlation to SPECMAP (Martinson et al., 1987) and refined the

ages and correlation of the detrital layers to North Atlantic detrital layers.

We present data from three jumbo piston cores (JPC15, JPC18, JPC19) collected on the

Eirik Drift in the summer of 2002 during Cruise KN166-14 of the RVKnorr, and from Core

MD99-2227 collected during the 1999 Images campaign (Figure 4-1). JPC15 was taken on the

upper slope of the ridge at a water depth of 2230 m. Core JPC19 was collected from the crest of

the ridge at a water depth of 3184 m, and Core JPC18 from the southern flank of the ridge at a

water depth of 3435 m. Core MD99-2227 was collected from the western toe of the drift at 3460

m water depth. The recovered sediments are mostly dark gray bioturbated silty clays, with clayey

silt and sandy mud, and occasional gray nannofossil/foraminifer rich clayey silt layers (see

Turon, Hillaire-Marcel et al., 1999, for a lithologic description of MD99-2227).

Methods

U-channel samples (2x2 cm square cross-section and 150 cm in length) were collected

from the center of the split face of piston core sections. These samples were measured on a 2G-

Enterprises pass-through cryogenic magnetometer at the University of Florida. Natural remanent

magnetization (NRM) was demagnetized step-wise using alternating fields (AF) in 5 mT









increments for 0-60 mT peak fields, and in 10 mT increments for 60 mT-100 mT peak fields.

Volume susceptibility was then measured using a susceptibility track specifically designed for u-

channels (Thomas et al., 2003) that has a measurement resolution of a few centimeters.

Anhysteretic remanent magnetization (ARM) was applied using an AF field of 100 mT and a

bias DC field of 50 pT. Isothermal remanent magnetization (IRM) was imparted using a 0.5 T

DC field. Both artificial remanences were demagnetized with the same AF steps used to

demagnetize NRM. Principal components were calculated from the NRM data using the method

of Kirschvink (1980) applied to the 20-80 mT interval. Relative paleointensity proxies were

generated by normalizing the NRM data by both ARM or IRM, demagnetized at a common peak

field. A mean of nine normalized remanence values, in the 20-60 mT peak field range, was used

to generate the relative paleointensity proxies. ARM and susceptibility data were also used to

ascertain magnetic grain size changes that help define detrital layers. The parameter karm

(anhysteretic susceptibility), obtained by normalizing ARM intensity by the strength of the dc

field used to acquire the ARM, was divided by volume susceptibility, to determine kam/k, a

proxy for magnetite grain size.

On completion of the magnetic measurements on the u-channel samples, X-radiographs

were taken across detrital layers, identified by u-channel magnetic measurements and carbonate

analyses, to provide a picture of the internal structure of these layers and identify the presence or

absence of traction structures. Discrete toothpick-sized samples, collected at 1-cm intervals

across detrital layers, were used for smear slide observation (Table 4-2) and for measurement of

magnetic hysteresis parameters using a Princeton Measurements Corp. vibrating sample

magnetometer (VSM). Magnetic hysteresis parameters provide a means of estimating magnetite

grain size, and therefore of recognizing grading in detrital layers.









Cores were sub-sampled for oxygen isotope analysis at 5-cm spacing. Samples from Core

MD99-2227 were analyzed at GEOTOP (Montreal) while samples from the KN166-14 cores

were analyzed in the stable isotope laboratory at Rutgers University. For all the cores,

foraminifer shells of the planktonic species Neogloboquadrinapachyderma (left coiling) were

picked in the 150-250 [tm fraction for the isotopic analyses. Planktonic foraminifer species were

used for the isotopic analyses due to the small amount of benthos present in the cores. For Core

MD99-2227, samples were collected at 5 cm intervals for carbonate analyses using an elemental

analyzer.

Age models for the piston cores were constructed by matching relative geomagnetic

paleointensity records and planktic 6180 records to target curves, with the location of magnetic

excursions (Laschamp and Iceland Basin) providing additional age constraints. The combination

of paleointensity records and oxygen isotope data provide enhanced temporal resolution

compared to using either dataset independently.

NRM and Normalized Remanence Record

The natural remanent magnetization (NRM) data for all four cores are shown as

component inclination, corrected component declination, and maximum angular deviation

(MAD) values (Figure 4-2). Cores were not oriented during collection, and therefore declination

data were corrected by aligning the mean declination of each core to North. Twisting within

cores during the coring process is indicated by anomalous declination changes in Core JPC 18

(114.5-189 cm) (Figure 4-2). Core MD99-2227 is affected by stretching in the upper 7 meters

that has significantly affected the magnetization directions (Figure 4-2).









Polarity Excursions

Brief polarity excursions are a characteristic of the geomagnetic field, at least during the

last -2 Myr, and excursions of known age provide useful stratigraphic markers. Component

magnetizations from u-channels indicate directional excursions at 9.3 meters below seafloor

(mbsf) in Core JPC15, at 13.4 mbsf in Core JPC18, and at 18.7 mbsf in Core JPC19 (Figures 4-2

and 4-3). For Core JPC15, the observed excursion is correlated to the Laschamp excursion (-41

ka). For Cores JPC18 and JPC19, the observed excursion is correlated to the Iceland Basin

excursion (-185 ka). Orthogonal projections of alternating field demagnetization data from

intervals recording the Iceland Basin excursion in Cores JPC18 and JPC19 (Figure 4-3) indicate

that the excursions are unambiguously recorded by u-channel samples and by discrete samples

collected alongside the u-channel trough.

Relative Paleointensity

It is generally accepted that the generation of useful paleointensity proxies requires that the

sediments contain magnetite as the only NRM carrier. Also the sediment should have a narrow

range of magnetite concentration, as indicated by magnetic concentration parameters varying by

less than an order of magnitude, and have restricted magnetite grain-size in the few micron grain-

size range, corresponding to pseudo-single domain grains (Tauxe, 1993). There is no evidence

from demagnetization characteristics of NRM, or from hysteresis parameters, for high-coercivity

magnetic minerals such as hematite or pyrrhotite. Using plots of anhysteretic susceptibility

against susceptibility, and the calibration of King et al. (1983), we estimate that these sediments

generally have magnetite grain sizes in the 1-10 [tm range (Figure 4-4). Records of ARM, IRM

and susceptibility (Figure 4-5) show that the concentration parameters generally vary within an

order of magnitude, the limit deemed suitable for determination of relative paleointensity proxies









(Tauxe, 1993). The exception is within the coarser-grained intervals in the early part of

interglacials, where the concentration parameters vary by more than an order of magnitude.

NRM measured on u-channel samples was normalized using both ARM and IRM,

demagnetized at the same peak fields as the NRM. To generate the paleointensity proxies, a

mean of nine demagnetization steps in the 20-60 mT interval were used to calculate mean

NRM/ARM and mean NRM/IRM. Although the two proxies are generally consistent with each

other, mean NRM/ARM has the lower standard deviations and was therefore chosen as the

preferred paleointensity proxy.

Chronology

To construct age models for the four cores in this study, we correlate the planktonic

oxygen isotope records to the benthic oxygen isotope stack (Lisiecki and Raymo, 2004). We then

adjust this correlation to optimize the fit of the relative paleointensity records to the

paleointensity record from ODP Site 983 (Channell et al., 1997; Channell, 1999). Following

Stoner et al. (2003), the paleointensity and oxygen isotope data from ODP Site 1089 were used

to improve the age model for ODP Site 983 particularly in the MIS 3-4 interval. For the Eirik

Drift cores, a combination of oxygen isotope data and relative paleointensity data can produce a

higher-resolution age model than would be possible using either data set independently.

The magnetic excursion recorded at 18.7 mbsf in JPC19 (Figures 4-2 and 4-3) is

interpreted as the Iceland Basin excursion (Channell et al., 1997; Channell, 1999). It lies in a

prominent paleointensity low at 185 ka in JPC19 (Figure 4-6), consistent with the expected age

of this excursion. According to the age model, Core JPC19, from the crest of the drift at a water

depth of 3184 m, has an age at its base of 300 kyrs with a mean sedimentation rate of 10.5

cm/kyr.









In Core JPC18, from southern flank of the Eirik ridge at a water depth of 3435 m,

sediments coeval with interglacial periods are apparently missing, as shown by the lack of

Holocene oxygen isotope values (Figure 4-7). MIS 5e is also absent in the record, because

oxygen isotope values in this interval are too high for full interglacial values. The polarity

excursion observed at 13.45 mbsf (Figures 4-2 and 4-3) is identified as the Iceland Basin

excursion and it occupies a distinct paleointensity low at 185 ka (Figure 4-7), an age consistent

with the observation of this excursion elsewhere. The overall mean sedimentation rate in Core

JPC18 is 9 cm/kyr.

Core JPC15 was taken on the upper slope of Eirik ridge at a water depth of 2230 m. The

polarity excursion observed at 9.3 mbsf in Core JPC15 (Figures 4-2 and 4-3) occurs within a

prominent paleointensity low at 40 ka (Figure 4-8) and is therefore interpreted as the

Laschamp excursion. The base of JPC 15 has an age of 160 ka and the mean sedimentation rate is

15 cm/kyr (Figure 4-8).

Core MD99-2227 shows significant stretching in the upper part of the core, however, the

correlation to the calibrated ODP Site 983 paleointensity record is possible in the lower part

(Figure 4-9). The paleointensity correlation is consistent with the correlation of the planktic

oxygen isotope record to the benthic oxygen isotope stack of Lisiecki and Raymo (2005). These

correlations give a basal age for Core MD99-2227 of 430 ka, and mean sedimentation rates of 10

cm/kyr (Figure 4-9).

Detrital Layer Stratigraphy

The ratio of anhysteretic susceptibility to susceptibility (karm/k) has been shown to be a

useful magnetite grain size proxy (e.g. King et al., 1983; Tauxe, 1993). Although the plots of kam

versus k of each core (Figure 4-4) indicate magnetic grain sizes within a restricted (few micron)

range, the karm/k data plotted versus age (Figure 4-10) indicate distinct broad intervals of low









values of kam/k that coincide with the early Holocene (when recorded), with MIS 5e, and with

the early parts of MIS 7, 9 and 11 (shaded in Figure 4-10). Low values of kam/k indicate

relatively coarse magnetite grain sizes in these intervals. Although Core JPC18 is missing part of

the Holocene, and almost the entire MIS 5e, the intervals of low values of kam/k appear to be

partially recorded.

Volume magnetic susceptibility data measured on u-channel samples from Cores JPC 19

and MD99-2227 show an increase in magnetic concentration in the early Holocene, MIS 5e, and

in the early parts of MIS 7, 9 and 11 (Figure 4-10). These intervals of high magnetic

concentration coincide with the intervals of low values of kam/k (Figure 4-10) that indicate

relatively coarse magnetite grain sizes.

In Core JPC15, high sedimentation rates between 20-60 ka (500-1200 cm) allow the

identification of millennial-scale cycles in volume magnetic susceptibility (Figure 4-5). These

appear to mimic the D/O cycles the Greenland Ice Core (GISP) oxygen isotope record, and are

reminiscent of susceptibility cycles identified by Kissel et al. (1999) in cores along the path of

North Atlantic Deep Water (NADW), and attributed to changes in the strength of bottom

currents. The depth of the WBUC, that varies in response to the relative outflows of water

masses from the Greenland and Norwegian Seas, could also be account for the variations.

In addition to these broad decimeter-scale intervals defined by kar/k and k values, a total

of seventeen cm-scale layers with magnetic properties and percent carbonate values significantly

different from the surrounding sediments have been identified in MD99-2227 (Figure 4-11).

These layers have been labeled according to marine isotope stage and their detrital carbonate

(DC) content. For example, 6LDC indicates a low detrital carbonate (LDC) layer within MIS 6

(Table 4-2).









Eight of the seventeen cm-scale layers are designated detrital carbonate layers (DC) on the

basis of their high detrital carbonate contents. Four of these layers (3DC, 7DCa, 8DC, 11DC) are

recognized by coarser grained magnetic material (compared to the background sediment), as

indicated by low karm/k values (Figure 4-11). One of these DC layers (7DCa) shows a peak in

magnetic susceptibility while the other seven DC layers do not. Two DC layers (5DC, 9DC)

show finer-grained magnetic material (compared to background sediment), and two DC layers

(7DCb, 2DC) are not differentiated by magnetic grain size from the background sediment but all

DC layers coincide with highs in percent carbonate and six show peaks in GRA bulk density

(Figure 4-11). All DC layers are light in color, do not show a sharp base, and appear to show

some bioturbation. The X-radiographs of these layers confirm a high concentration of IRD, but

no laminae or evidence for traction (Figure 4-12). Smear slides indicate a high percentage of

coarse detrital carbonate material in these layers (Table 4-2).

Nine of the seventeen cm-scale detrital layers are designated low detrital carbonate (LDC)

layers (Figure 4-11). These do not feature an increase in percent carbonate, but show a peak in

magnetic susceptibility, a low in karm/k, and an increase in GRA bulk density. These LDC layers

occur within MIS 1, 2, 5, 6, 7, 9 and 11 and show sharp bases, bioturbated tops and are 4-18 cm

thick (Figure 4-11, Table 4-2). The X-radiographs indicate a sharp base and laminae within the

layers (Figure 4-12), some of the laminae are inclined and indicative of traction, implying rapid

deposition from turbidity currents or contourites.

Toothpick-sized samples collected at 1-cm intervals through detrital layers were used to

determine magnetic hysteresis parameters that can be used as a means of assessing the grain size

of magnetite (Day et al., 1977). All but one of the detrital layers exhibit hysteresis parameters

that fall within the pseudo-single domain (PSD) grain size range (Figure 4-13). The detrital









carbonate layer identified in MIS2 (2DC) shows coarse multi-domain magnetite that is

anomalous compared to all other detrital layers (Figure 4-13). For five of the nine LDC layers,

we see evidence for progressive change in hysteresis parameters through the detrital layer

indicative of grading, fining upward from the base of the layer. Bioturbation of the detrital layer

into the overlying sediment could also cause the layer to appear graded. However, the presence

of distinct laminae within the LDC layers shows that no bioturbation of the layer has occurred.

None of the DC layers show this "grading" in hysteresis parameters. The presence of grading in

the LDC layers indicates a turbiditic rather than a contourite origin for these layers.

Smear slides indicate that LDC layers contain little clay and significant amounts of silt-

sized opaque grains, green hornblende and quartz. Trace amounts of detrital carbonate are

present in LDC layers and throughout the rest of the core, whereas the percentage of detrital

carbonate in the DC layers exceeds 10% (Table 4-2).

Discussion

Sedimentation rates on the Eirik Drift have been shown to be greatly affected by changes

in the strength and bathymetry of the Western Boundary Undercurrent (WBUC) that is thought

to be switched off during glacials and active during interglacials (Hillaire-Marcel et al., 1994;

Hillaire-Marcel and Bilodeau, 2000). The core of this current is thought to occupy water depths

between 2500 and 3000 meters (Hillaire-Marcel et al., 1994), resulting in winnowing and almost

complete removal of Holocene and MIS 5e sediment from these depths. Cores from outside the

influence of the flow would be expected to have interglacial sedimentation rates comparable to,

or higher than, glacial sedimentation rates.

When combined with previous studies carried out on the drift, the new results indicate that

both water depth and position on the drift influence interval sedimentation rates. Although the

site of Core JPC18 is located -450 meters below the supposed core of the WBUC, sediment of









Holocene and MIS 5e age is missing at this site. This implies that the WBUC is active at deeper

water depths than previously supposed on the southern side of the Eirik ridge (Figure 4-1). This

may be consistent with a deep branch of the WBUC, with a gyre in the outer Labrador Sea that

feeds the Gloria Drift (Figure 4-1).

Cores HU90-013-013 (water depth 3471 m), JPC19 (water depth 3184 m) and MD99-2227

have relatively high Holocene sedimentation rates of 35 cm/kyr, -13 cm/kyr, and 10 cm/kyr

respectively. Sedimentation rates in cores MD99-2227 and JPC19 appear to be low at the onset

of deglaciation and then increase. This may be due to increased winnowing by the WBUC at the

onset of the deglaciation, offset by increased detrital input as the deglaciation proceeds.

Core HU90-013-012 at 2830 meters water depth lies within the influence of the WBUC

and has very low sedimentation rates in the Holocene (Stoner et al., 1995a, 1996). Higher up the

slope, Core JPC15 at a water depth of 2230 meters has low sedimentation rates in the Holocene

and MIS 5e, although the site supposedly lies outside the main influence of the WBUC. Hillaire-

Marcel et al. (1994) noted that, in core HU90-013-06 at even shallower water depths (1105 m)

on the Eirik ridge, active bottom currents also resulted in very low Holocene sedimentation rates.

Hillaire-Marcel et al. (1994) interpreted DC and LDC layers deposited during the last

glacial cycle at Orphan Knoll, on the western side of the Northwest Atlantic Mid-Ocean Channel

(NAMOC), as being related to ice advances of the Laurentide Ice Sheet that triggered turbiditic

flows down the NAMOC (Figure 4-1). Sediment suspended by these flows is thought to have

deposited cm-scale sandy mud beds rich in detrital carbonate (DC layers) at Orphan Knoll. Not

all the detrital layers observed at Orphan Knoll are recognized on Eirik Drift, although two LDC

layers and one DC layer in Core HU90-0130-013 (Figure 4-1) were considered coeval with

Orphan Knoll detrital layers (Stoner et al., 1996).









The cm-scale detrital layers identified in core MD99-2227 extend the record of detrital

layers beyond the last glacial cycle. Detrital layers on Eirik Drift occur during both glacial and

interglacial conditions. However, the layers occurring in the interglacials are close to the

Terminations in the Holocene, MIS 5, 7 and 11. It is only in MIS 9 that the DC layer appears to

occur in the later part of the interglacial implying that the Laurentide Ice Sheet was present

throughout MIS 9.

Detrital layer 1LDC with an age of 13 ka in MD99-2227 (Table 4-3) is tentatively

correlated to DCO of Stoner et al. (1998). Layer 2LDC has an age of 18 ka and is correlated to

DC1 (16 ka) from Orphan Knoll (Stoner et al., 1998) and with H1 of Bond et al., (1999) from the

central Atlantic. The DC layer 2DC correlates with DC2 of Stoner et al. (1998) and with H2

(Bond et al., 1999). The detrital layer labeled 3DC (39 ka) is correlated to DC4 from Orphan

Knoll and to H4 (38 ka). As discussed above, the characteristics of LDC layers implies

deposition by turbidity currents (derived from the Greenland Slope). If so, this turbiditic activity

is sometimes coeval with Heinrich layers of the central Atlantic and with detrital events at

Orphan Knoll.

The ages of layers designated 2LDC, 2DC and 3DC in this study are consistent with ages

for Heinrich events H1, H2, and H4 (Table 4-3). No identifiable events that coeval with Heinrich

events H3, H5 or H6 are found. Hiscott et al. (2001) identified Heinrich-like detrital layers in

core MD95-2025 from near Orphan Knoll back to MIS 9. Two detrital carbonate layers within

early MIS 5 at Orphan Knoll (H8 and H9 of Hiscott et al., 2001) appear to be coeval with DC

events identified on Eirik Drift, implying that instabilities of the Laurentide Ice Sheet are

recorded at both sites. Detrital carbonate layers within MIS 7 and MIS 9 at Orphan Knoll (H10

and H13 of Hiscott et al., 2001) are coeval with a LDC layers (7LDC and 10 LDC) identified on









Eirik Drift (Table 4-3), implying that the LIS instabilities that triggered the detrital carbonate

layers at Orphan Knoll were coeval with instabilities on the Greenland slope that triggered the

LDC layers on Eirik Drift. Such conclusions are highly dependent on the resolution of

stratigraphic correlation. While stratigraphic correlation of detrital layers from the Orphan Knoll

to the central Atlantic for the last glacial cycle is rather well constrained (Bond et al., 1999;

Stoner et al., 1996, 2000), the correlations beyond the last glacial cycle are considerably more

speculative (e.g. Hiscott et al., 2001; van Kreveld et al., 1996) due to lack of stratigraphic

resolution that inhibits unequivocal correlation of detrital layers.

Conclusions

Piston cores collected from Eirik Drift have produced records of relative paleointensity and

of the Laschamp and Iceland Basin polarity excursions that augment oxygen isotope data for

generating age models. Magnetic data from cores JPC19 and MD99-2227 show broad intervals

of increased magnetic grain size and concentration during MIS 5e and at the MIS 2/1 transition,

consistent with observations from Core HU90-013-013 (Stoner et al., 1995b). Core MD99-2227

also shows a similar increase in magnetic grain size and concentration at the onset of interglacial

MIS 7, 9 and 11, implying that retreat of the Greenland Ice Sheet produced a characteristic

detrital signal at the onset of all interglacial stages over the last 400 kyr.

Seventeen cm-scale detrital carbonate and low detrital carbonate layers are identified in

MD99-2227 (Figure 4-11, Table 4-2). They occur in both glacial and interglacial stages. The

detrital layers can be subdivided into two classes. Detrital carbonate (DC) layers are composed

of carbonate-rich IRD. They usually, but not always, carry a magnetic signal indicating high

magnetic concentration and increased magnetic grain size relative to background sediment. Low

detrital carbonate (LDC) layers have <10% detrital carbonate, usually show evidence (from

magnetic hysteresis ratios) for fining-upward grading, and X-radiograph evidence for traction.









These layers are also usually marked by high magnetic concentration and increased magnetic

grain size relative to background sediment.

Based on the differences between DC and LDC layers, we interpret the former as Hudson

Strait derived detrital layers, and the latter as layers dominated by material from turbidites

derived from the Greenland slope. 1LDC, 2LDC, 2DC and 3DC are correlative with detrital

layers observed at Orphan Knoll (Stoner et al., 1996) (Table 4-3). Three of them (1LDC, 2DC

and 3DC) are coeval with central Atlantic Heinrich layers H1, H2 and H4 (Bond et al., 1999).

Beyond the last glacial cycle, the correlation of detrital layers from Eirik Drift (this paper) to

Orphan Knoll (Hiscott et al., 2001) and to the central Atlantic (van Kreveld et al., 1996) is

limited by the imprecision of stratigraphic correlation (Table 4-3). Nonetheless, as illustrated

here, the use of paleointensity-assisted chronostratigraphy, the combination of relative

paleointensity with standard oxygen isotope stratigraphy, improves stratigraphic correlations

across the northern North Atlantic Ocean (and beyond), and thereby facilitates the interpretation

of detrital layers in terms of their correlation, aerial extent and provenance.









Table 4-1. Core, latitude, longitude, water depth and base age of the core.


Core

JPC15
JPC18
JPC19
MD99-2227


Latitude Longitude Water
depth
-45.57 58.20 2230
-47.13 57.19 3435
-47.60 57.58 3184
-48.22 58.12 3460


Base age
(kyr)
150
300
250
430











Table 4-2. DC and LDC layer properties in Core MD99-2227.


event Thick- depth Age MIS Name k
ness (cm) (ka) peak
(cm)


5 440.22 13.04
14 616.85 18.2
15 663 21.4
21 858.7 39.1
6 1872.3 111.68
16 2019.4 129.34
14 2192.9 152.17
16 2505.4 191.58
12 2700 214.98
6 2872.3 233.42
11 3083.7 266.57
17 3229.2 289.89
5 3536.4 335.6
18 4008.2 391.03
4 4084.2 403.48
7 4133.7 409.57
7 4240 421.43


1/2 1LDC yes
2 2LDC yes
2 2DC no
3 3DC no
5 5LDC yes
5 5DC no
6 6LDC yes
7 7DCa yes
7 7DCb no
7 7LDC yes
8 8DC no
9 9DC no
9/10 9LDC yes
11 11LDCa yes
11 11LDCb yes
11 11LDCc yes
11 11DC no


kan/k % GRAPE % X-Ray Sharp
carb density detr. base
carb.


coarse low peak
coarse low peak
high peak
coarse high peak
coarse low peak
fine high peak
coarse low peak
coarse high
high peak
coarse low peak
coarse high peak
fine high peak
coarse low peak
coarse low peak
coarse low peak
coarse low peak
coarse high peak


10 traction yes
trace traction yes
15 no
40 no
trace yes
70 no
trace traction yes
20 no
25 no
trace traction yes
20 no
30 no
trace traction yes
trace traction yes
5 traction yes
10 traction yes
50 IRD rich no


Grading


yes
yes
no
no
no
no
yes
no
no
no
no
no
no
no
yes
yes
no











Table 4-3. Detrital Layers from other studies considered to be correlative to detrital layers
identified on Eirik drift.


Event Name


H-layers Bond
et al. (1999)
(age ka)

H1 (16.8)
H2 (24)
H4 (38)


Hiscott et al
(2001)
(age ka)
H1(11-12)

H2 (18-22)
H4 (39-42)
H8(92-108)
H9(121-126)


Van Kreveld
et al (1996)
(age ka)
hl (15)

h2 (21)
h4 (40-43)

h7 (128-131)


depth
(cm)


MD99-2227 Stoner et al.
Age (ka) (1998)
(age ka)
440.22 13.04 DCO (12)
616.85 18.2 LDC1 (18)
663 21.4 LDC3 (21)
858.7 39.1 DC4 (36)
1872.3 111.68
2019.4 129.34
2192.9 152.17
2505.4 191.58
2700 214.98
2872.3 233.42
3083.7 266.57
3229.2 289.89
3536.4 335.6
4008.2 391.03
4084.2 403.48
4133.7 409.57
4240 421.43


h12 (189)


H10(231-240)



H13(335-340)


1 1LDC
2 2LDC
3 2DC
4 3DC
5 5LDC
6 5DC
7 6LDC
8 7DCa
9 7DCb
10 7LDC
11 8DC
12 9DC
13 9LDC
14 11LDCa
15 11LDCb
16 11LDCc
17 11DC













48 W


"' 58 N 58N
4"Y -- .^ ... -'K V '----c D rif
Drift
--0- *. JPCI


CANADA \
T VV ^'MD-2024 \ *
,. ,JPC18
Orphan 57 N 57 N
'.- oll 48 W 46 W












5r7



Figure 4-1. Location map showing the Labrador Sea from Hillaire-Marcel and Bilodeau (2000)
and the location of piston cores JPC15, JPC18, JPC19, and MD99-2227. Black
arrows indicate the path of the Western Boundary Undercurrent. NAMOC: Northwest
Atlantic Mid-Ocean Channel.













s 40
4 o Laschamp JPC15
c excursion
= 40
-00 350 C

;:w ir 3To 9
V -200 | 4
150








4000 4
0 500 ]000 I _500 2000


.Pcc 4aIceland Basin
0 excursion c

0 500 000 5et (( n) on 2000
: 250 1 B
-80 excurs f JPCy I50

GO a C
250


050
Depth (cm)



o 4t n clnio, c t c d t ad m u a

-JPCI Iceland as
1 4A tt01"AuJ.r A. 5
20 I







o 500 1000 1500 2000
Depth (cm)

















S MD99-2227
JPC19 Iceland Basin
-. 40 excursion |

33M












34T U 2r 200;;
:" i i : I
|25 4 k a









0 500 1000 20) 2500 2000 00
Depth (cm)










Figure 4-2. Component inclination, corrected component declination and maximum angular'
deviation (MAD) MD99-2227


'; I S I 1 ; 1 i3 ; 200150
i^tlA i *; i. : '2100





0 W I"0 1500 200. 2300 30" .500 4W0
Depth (cm)



Figure 4-2. Component inclination, corrected component declination and maximum angular
deviation (MAD) values for cores JPC15, JPC19, JPC18 and MD99-2227.












Laschamp Iceland Basin
-80 ----~ --I I -
. JPCI5 JPC18


-40 r 4 *
900 910 920 930 0 95(1 1330 130 1350 1360 1370 1380

Wm









900 910 920 930 940 9501 1330 1-40 1390 1360 1370 1390
Depth (cm) Depth (cm)


Iceland Basin Excursion


Discrete sample (lcm3) Discrete sample (8cm3)
JPC18 NUp

section 7


Deconvolved
u-channel data


U-channel
data


Depth- 135 Inm
Lo treat= 0 mT
Hi real= I0W mT


Deplh=18.73m
LO treat= 0 mT
Hi trat= 100 mT






S/Dn


NNUp



I Depth-18.74m
SL treat- 0 mT
Hi treat- 500 mT


-------B


Figure 4-3. a). Component inclination, declination and maximum angular deviation (MAD)
values recording Laschamp and Iceland Basin polarity excursions from piston cores
JPC15, JPC18 and JPC19. Key: U-channel data (closed circles), deconvolved u-
channel data (open squares-dashed line) using the method of Guyodo et al. (2003), 8-
cm3 discrete sample cubes (open squares) and 1-cm3 cubes (diamonds).3b).
Orthogonal projections from the Iceland Basin excursion from cores JPC19 and
JPC18, from u-channel data, deconvolved u-channel data, and discrete samples. Open
circles represent the vector end point projection on the vertical plane. Closed circles
represent the vector end point projection on the horizontal plane.


1855 1860 1865 1870 1875
Depth (cm)


JPC19
section 4


Depth 18.75m
Lo treat= 0 mT
Hi treat= 00 rnT











MD99-2227


0 2 4 6 8 10 12 14
Volume susceptibility (10-3 SI units)
JPC 15


JPC19


4-
l20-25




22


U L



0
0 1 2 3 4 5 6
V,.lueni, Iic pi-II.t, (10-3 SI units)
jI'CIS


Vo1 2 3 4 5 6
Volume susceptibility I (I SI units)


7 0 1 2 3 4
Volume susceptibility (10-3 S1 units)


Figure 4-4. Anhysteretic susceptibility (karm) plotted against volume susceptibility (k) for JPC 18,
JPC19, JPC15 and MD99-2227. Diamonds indicate 'background' sediment, red
squares indicate coarse decimeter-scale interglacial intervals, and blue circles indicate
cm-scale detrital layers. Black lines indicate magnetic grain-size boundaries placed
using the calibration of King et al. (1983).











S25
-JPC19


E II I 5



2P
A vi A i 5 2'i "1
-" '* / '. ^ '.' 1. '



_"r 'p"' '. "


0 30
00

04 a
3- -1 ..






3 --











J S, 1000 1510 2BOO 2SDO
-','









MD99-2227 20
." / / : "" A *,',/,. 0 i
*, I "10 ^ '





10,1


4.4

0 I
0
0.
0 500 1000 1500 2000 2500
Depth (cm)

I i i-' 25

FMD9,9-2227 .



JPC19. Orange-IRM, green-ARM, red-NRM, blue-volume susceptibility.


0_ '0.6
-0.4 $
0.2 2

E 6 -
00


O 10 0 2000 3000 4000
Depth (cm)
Figure 4-5. NRM, ARM, IRM and volume susceptibility for MD99-2227, JPC15, JPC18 and
JPC 19. Orange-IRM, green-ARM, red-NRM, blue-volume susceptibility.























2 -


2.5 .5

4 3
5 isiecki and Raymo (2005).5
5 -4

4.5
30 5
r= 25 5.5
E 20'
10

0 50 100 150 200 250 300

Age (ka)
Figure 4-6. JPC19: Relative paleointensity record correlated to that from ODP Site 983
(Channell et al., 1997; Channell, 1999). Lower plot: planktic 180O data from JPC19
correlated to the benthic 6180 stack of Lisiecki and Raymo (2005). Interval
sedimentation rates are shown in orange.












































40 80 120


S1 I I20 I I I I20
160 200 240 280 320


Age (ka)


. JPC18: Relative paleointensity data correlated to ODP Site 983 (Channell et al.,
1997; Channell, 1999). Lower plot shows planktic 6180 data correlated to the benthic
6180 stack of Lisiecki and Raymo (2005). Interval sedimentation rates are shown in
orange.


3.5
4
4.5

5_




Z 12 2


Y 4

0
C)v


Figure 4-7























JPCIS -3 0
2.5 4
3 4.5
S 3.5
s 3Lisiecki and Raymo (2005)



5
5.5 0

0 20 40 60 80 100 120 140 160
Age (ka)



Figure 4-8. JPC15: Relative paleointensity data correlated to ODP Site 983 (Channell et al.,
1997; Channell, 1999). Lower plot shows planktic 18O data correlated to the benthic
680 stack of Lisiecki and Raymo, (2005). Interval sedimentation rates are shown in
orange.













MD99-2227




i~ft


3.5
4 4*


3.5 Lisieki and Rayrno i2O'i 5
4



25
5.5- 20 _

0 10
0 100 200 300 400
Age (ka)



Figure 4-9. MD99-2227: Relative paleointensity data correlated to ODP Site 983 (Channell et al.,
1997; Channell, 1999). Lower plot shows planktic 6180 data correlated to the benthic
180 stack of Lisiecki and Raymo (2005). Black bar indicates stretched interval due
to coring. Interval sedimentation rates are shown in orange.




Full Text

PAGE 1

1MAGNETIC STRATIGRAPHY AND ENVIRONMENTAL MAGNETISM OF OCEANICSEDIMENTSByHELEN F. EVANSA DISSERTATION PRESENTED TO THE GRADUATE SCHOOLOF THE UNIVERSITY OF FLORIDA IN PARTIAL FULFILLMENTOF THE REQUIREMENTS FOR THE DEGREE OFDOCTOR OF PHILOSOPHYUNIVERSITY OF FLORIDA2006

PAGE 2

2Copyright 2006byHelen F. Evans

PAGE 3

3For Jane and Eryl

PAGE 4

4ACKNOWLEDGMENTSI would like to thank my advisor, Jim Channell, for giving me the opportunity to study inFlorida and for all his help and support over the last 7 years. I also thank my committee membersEllen Martin, Neil Opdyke, John Jaeger and Bo Gustafson for agreeing to supervise my researchover the last five years and for all their help, academic and otherwise. I also had the pleasure ofcollaborating and interacting with a number of talented individuals without whose help I wouldnot have been able to accomplish this work: Gary Acton, Paul Bown, Yohan Guyodo, SeanHiggins, Claude Hillaire-Marcel, Dave Hodell, Kainian Huang, Mark Leckie, Ulla Rohl, JosephStoner, Ray Thomas, Thomas Westerhold and many others. My research was made possible bythe technical and financial support of several organizations including the National ScienceFoundation, the JOIDES U.S. Sciences Support Advisory Committee (USSAC), and theGraduate Student Council. Support was also provided by the Institute for Rock Magnetism, theCollege of Liberal Arts and Sciences, the Graduate School, the McLaughlin family, and thedepartment of Geological Sciences at the University of Florida.I thank all those who gave me their moral support during my years of study, which oftenconsisted of many hours of festivities in the numerous bars and restaurants in Gainesville, SanFrancisco and further afield. I thank in particular Gillian Rosen, Joe Graves, Joann and JasonHochstein, Howie Scher, George and Katherin Kamenov, Cara Gentry, Jen Mays, Steve Volpe,Phil DAmo, Bricky Way, Kendall Fountain, Adi Gilli, Simon Nielsen, Dave Hodell, MikeRosenmeier, William Kenney, Yohan Guyodo and Victoria Meija. Finally I would like toexpress my gratitude to my family without whose support I would not have been able tocomplete this work. I thank my late mother Eryl, my father Terry and my brother Michael. I alsothank Diane and John Thomas, Judy and Robin Ganz, and Molly and Reg Beynon.

PAGE 5

5TABLE OF CONTENTSpage ACKNOWLEDGMENTS...........................................................................................................4LIST OF TABLES......................................................................................................................7LIST OF FIGURES....................................................................................................................8ABSTRACT.............................................................................................................................12CHAPTER1INTRODUCTION.............................................................................................................142LATE MIOCENE-HOLOCENE MAGNETIC POLARITY STRATIGRAPHY ANDASTROCHRONOLOGY FROM ODP LEG 198-SHATSKY RISE...................................18Introduction.......................................................................................................................18Methods.............................................................................................................................19Magnetostratigraphic Interpretation....................................................................................20Astrochronology................................................................................................................23Discussion.........................................................................................................................24Conclusions.......................................................................................................................273INTEGRATED NEOGENE MAGNETIC, CYCLE AND BIOSTRATIGRAPHYFROM ODP SITE 1208 (SHATSKY RISE, PACIFIC OCEAN)........................................48Introduction.......................................................................................................................48Site Location and Lithology...............................................................................................51Magnetic Stratigraphy........................................................................................................51Cycle Stratigraphy.............................................................................................................53Calcareous Nannofossils....................................................................................................54Planktonic Foraminifera.....................................................................................................56Conclusions.......................................................................................................................584PALEOINTENSITY-ASSISTED CHRONOSTRATIGRAPHY OF DETRITALLAYERS ON THE EIRIK DRIFT (NORTH ATLANTIC) SINCE MARINE ISOTOPESTAGE 11.........................................................................................................................75Introduction.......................................................................................................................75Methods.............................................................................................................................76NRM and Normalized Remanence Record.........................................................................78Polarity Excursions.....................................................................................................79Relative Paleointensity................................................................................................79Chronology........................................................................................................................80Detrital Layer Stratigraphy.................................................................................................81

PAGE 6

6Discussion.........................................................................................................................84Conclusions.......................................................................................................................875RELATIVE PALEOINTENSITY STACK FOR THE LAST 85 KYR ON A REVISEDGISP CHRONOLOGY, AND ENVIRONMENTAL MAGNETISM OF THEGARDAR DRIFT............................................................................................................105Introduction.....................................................................................................................105Site Locations..................................................................................................................107Methods...........................................................................................................................108Directional Magnetic Data...............................................................................................110Normalized Remanence...................................................................................................111Stable Isotope Data and Age Models................................................................................112Bulk Magnetic and Physical Parameters...........................................................................112Relative Paleointensity Stack...........................................................................................113Environmental Magnetism...............................................................................................115Conclusions.....................................................................................................................1186RELATIVE GEOMAGNETIC PALEOINTENSITY IN THE GAUSS AND GILBERTCHRONS FROM IODP SITE U1313 (NORTH ATLANTIC).........................................136Introduction.....................................................................................................................136Methods...........................................................................................................................138Results.............................................................................................................................139Discussion.......................................................................................................................1427ODP SITE 1092 REVISED COMPOSITE DEPTH SECTION HAS IMPLICATIONSFOR UPPER MIOCENE "CRYPTOCHRONS"...............................................................161Introduction.....................................................................................................................161Revised Composite Depths (rmcd)...................................................................................162Implications for Magnetic Stratigraphy............................................................................1638ASTRONOMICAL AGES FOR MIOCENE POLARITY CHRONS C4AR-C5R (9.3-11.2 MA), AND FOR THREE EXCURSION CHRONS WITHIN C5N.2N.....................171Introduction.....................................................................................................................171Methods and Results........................................................................................................173Comparison with Other Timescales..................................................................................175Excursion Chrons.............................................................................................................1789CONCLUSIONS AND FUTURE WORK.......................................................................189LIST OF REFERENCES........................................................................................................191BIOGRAPHICAL SKETCH...................................................................................................204

PAGE 7

7LIST OF TABLESTable page 2-1Latitude, longitude, water depth....................................................................................282-2Magnetostratigraphic age model....................................................................................292-3Comparison of astrochronological age models...............................................................302-4Astrochronological ages for Leg 198.............................................................................313-1Depths of reversal boundaries from ODP Site 1208.......................................................593-2Astronomically calibrated ages for reversal boundaries from ODP Site 1208.................603-3Nannofossil datums for ODP Site 1208.........................................................................613-4Plio-Pleistocene foraminfer datums...............................................................................623-5Miocene foraminifer datums..........................................................................................634-1Core, latitude, longitude, water depth and base age of the core......................................894-2DC and LDC layer properties in Core MD99-2227........................................................904-3Detrital Layers from other studies considered to be correlative to detrital layersidentified on Eirik drift..................................................................................................915-1Summary of the cores used in this study and the eleven cores used in the relativepaleointensity stack.....................................................................................................1206-1Depth of polarity chrons from IODP Site U1313.........................................................1466-2Polarity reversal ages determined at Site U1313..........................................................1477-1Adjusted depths of core tops from ODP site 1092........................................................1667-2Position of the polarity zone boundaries at site 1092....................................................1678-1Astronomical ages from recent timescales compared with those inferred at ODP Site1092............................................................................................................................181

PAGE 8

8LIST OF FIGURESFigure page 2-1Bathymetric map of Shatsky Rise. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 2-2Representative orthogonal projections of AF demagnetization data...............................332-3Site 1207 component inclination and declination from discrete samples for 0-80meters...........................................................................................................................342-4Site 1207 component inclination and declination from discrete samples for 80-160meters...........................................................................................................................352-5Interval sedimentation rates and age versus depth..........................................................362-6Site 1209 component inclination and declination from discrete samples........................372-7Site 1210 component inclination and declination from discrete samples........................382-8Site 1211 component inclination and declination from discrete samples........................392-9Site 1212 component inclination and declination from discrete samples........................402-10Power spectra................................................................................................................412-11The astronomical solution for obliquity compared with tuned L* reflectance datafrom Site 1207...............................................................................................................422-12The astronomical solution for obliquity compared with tuned L* reflectance datafrom Site 1208...............................................................................................................432-13The astronomical solution for obliquity compared with tuned L* reflectance datafrom Site 1209...............................................................................................................442-14The astronomical solution for obliquity compared with tuned L* reflectance datafrom Site 1210...............................................................................................................452-15The astronomical solution for obliquity compared with tuned L* reflectance datafrom Site 1211...............................................................................................................462-16Cross-spectral analysis..................................................................................................473-1Bathymetric map showing the location of Shatsky Rise in the Pacific Ocean.................643-2Inclination, declination and MAD values plotted against meters below sea floor...........653-3Inclination, declination and MAD values.......................................................................66

PAGE 9

93-4Inclination, declination and MAD values.......................................................................673-5Orthogonal projections showing AF demagnetization data.............................................683-6Interval sedimentation rates...........................................................................................693-7Reflectance (L*) data....................................................................................................703-8Plio-Pleistocene planktonic foraminifer and calcareous nannofossil datums...................713-9Miocene planktonic foraminifer and calcareous nannofossil datums..............................723-10Calcareous nannofossil biostratigraphy..........................................................................733-11A proposed biostratigraphy for the mid-latitude North Pacific.......................................744-1Location map showing the Labrador Sea.......................................................................924-2Component inclination, corrected component declination and maximum angulardeviation.......................................................................................................................934-3Component inclination, declination and maximum angular deviation (MAD) valuesrecording Laschamp and Iceland Basin polarity excursions...........................................944-4Anhysteretic susceptibility (karm) plotted against volume susceptibility (k)....................954-5NRM, ARM, IRM and volume susceptibility.................................................................964-6JPC19: Relative paleointensity record correlated to that from ODP Site 983..................974-7JPC18: Relative paleointensity data correlated to ODP Site 983....................................984-8JPC15: Relative paleointensity data correlated to ODP Site 983....................................994-9MD99-2227: Relative paleointensity data correlated to ODP Site 983.........................1004-10karm/k and magnetic susceptibility versus age...............................................................1014-11Core MD99-2227: karm/k, magnetic susceptibility, bulk (GRAPE) density...................1024-12Photographs and X-radiographs of three detrital..........................................................1034-13Hysteresis ratios Mr/Ms plotted versus Hcr/Hc............................................................1045-1Location map for cores analyzed in this study..............................................................1215-2Correlation of the magnetic susceptibility records........................................................1225-3Orthogonal projections of alternating field demagnetization data.................................123

PAGE 10

105-4Component inclination, declination and maximum angular deviation (MAD) values...1245-5Plot of anhysteretic susceptibility (karm) versus volume susceptibility (k).....................1255-6Paleointensity proxies..................................................................................................1265-7Core JPC13 benthic oxygen isotope record..................................................................1275-8Relative paleointensity records from Cores JPC2, JPC5 correlated to Core JPC13.......1285-9Interval sedimentation rates for Cores JPC2, JPC5 and JPC13.....................................1295-10Core JPC13: GRA bulk density ..................................................................................1305-11Anhysteretic susceptibility divided by volume magnetic susceptibility........................1315-12Eleven relative paleointensity records from the North Atlantic Ocean..........................1325-13The new relative paleointensity stack..........................................................................1335-14Comparison of the EHC06 paleointensity stack to 36Cl flux.........................................1345-15Comparison of the EHC06 paleointensity stack...........................................................1356-1Location map for IODP Site U1313.............................................................................1486-2Magnetic polarity stratigraphy from IODP Site U1313 in the 120-200 mcd interval.....1496-3Magnetic polarity stratigraphy from IODP Site U1313 in the 200-280 mcd interval....1506-4Vector end-point projections of AF demagnetization data............................................1516-5Interval sedimentation rates.........................................................................................1526-6Gauss Chronozone at Site U1313................................................................................1536-7The magnetic grain size proxy, anhysteretic susceptibility divided by susceptibility....1546-8Later part of the Gilbert Chronozone at Site U1313.....................................................1556-9Relative paleointensity records from IODP Site U1313...............................................1566-10Volume magnetic susceptibility from u-channel samples.............................................1576-11Volume magnetic susceptibility from u-channel samples and L* reflectance datameasured shipboard.....................................................................................................1586-12Mean volume magnetic susceptibility..........................................................................1596-13Output of a gaussian filter centered on a period of 41 kyr............................................160

PAGE 11

117-1Fe intensity (XRF) data plotted as a five-point moving average...................................1687-2Inclination of the characteristic magnetization component...........................................1697-3Site 1092.....................................................................................................................1708-1Magnetic component inclination for the C4Ar.1n-C5r.1n interval................................1828-2Oxygen isotope records from the C4An-C5r.1n interval at ODP Site 1092..................1838-3Power spectrum generated from the oxygen isotope stack in the depth domain............1848-4Upper plot shows the correlation of the filtered (filter centered at 0.0244 0.0073kyr1) oxygen isotope stack to the astronomical solution for obliquity...........................1858-5Interval sedimentation rates for the C4Ar.1n-C5r.1n interval.......................................1868-6Comparison of the age estimates of polarity chrons at ODP Site 1092.........................1878-7The Site 1092 relative paleointensity record for C5n.2n...............................................188

PAGE 12

12Abstract of Dissertation Presented to the Graduate Schoolof the University of Florida in Partial Fulfillment of theRequirements for the Degree of Doctor of PhilosophyMAGNETIC STRATIGRAPHY AND ENVIRONMENTAL MAGNETISM OF OCEANICSEDIMENTSByHelen F. EvansDecember 2006Chair: James E. T. ChannellMajor: Geology This dissertation presents the results of chronostratigraphic studies on marine sediment cores from three Oceans. Using a combination of magnetic stratigraphy, biostratigraphy and cycle stratigraphy it is possible to produce chronostratigraphies that exceed the resolution of any individual technique. In the North Atlantic, environmental magnetic records from Eirik Drift, south of Greenland, record detrital signals related to the melting of the Greenland and Laurentide Ice Sheets. The detrital layer stratigraphy has been placed in a paleointensity-assisted chronostratigraphic template, based on paleointensity and stable isotope data, to enhance correlation of detrital layers across the North Atlantic region. In the central Atlantic, on Gardar Drift, correlation of a benthic oxygen isotope record to the Greenland and Vostok Ice cores has placed cores from the drift on a revised GISP chronology. A stack of relative paleointensity records was developed and placed on the revised GISP chronology. In marine isotope stage 3, a benthic isotope record appears to record changes in bottom water temperature that are coeval with magnetic grain size changes. IODP Site U1313 from the North Atlantic produced a high-resolution polarity stratigraphy and relative paleointensity record between 2.5 and 6.0 Ma. This is one of a handful

PAGE 13

13of paleointensity records for this interval. Cycles in magnetic susceptibility allowed age-calibration by correlation to a benthic oxygen isotope stack.Sediment cores from the Pacific Ocean produced excellent magnetic stratigraphies, andcycles in the sediment allowed astronomic calibration of reversal boundaries. Based on thecorrelation of planktonic foraminifer datums to the magnetic stratigraphy at ODP Site 1208, anew planktonic foraminifer zonation for the northwest Pacific Ocean can be precisely correlatedto polarity chrons and astronomically calibrated ages. Numerous paleomagnetic excursions aretentatively identified for the first time in Pacific sediments.Oxygen isotope records from the Late Miocene (9.3-11.2 Ma) at ODP Site 1092 (SouthAtlantic) allowed astronomic calibration of ages of reversal boundaries and three polarityexcursions within Chron 5. This is the first time astronomically calibrated ages have beenassigned to these polarity excursion chrons and indicate a duration for the excursions of 3-4 kyrs.

PAGE 14

14CHAPTER 1INTRODUCTIONStratigraphy is a fundamental part of Geology. Earth processes unfold over a great range oftime scales from millions of years to minutes and seconds. One of the challenges in stratigraphyis to be able to assign dates to events in the geologic record. The geologic timescale is the meansby which we can understand the history of the Earth and magnetic reversal stratigraphy providesthe central framework for the geologic timescale to which other dating techniques(biostratigraphic, radiometric, orbital) can be tied. This is because magnetic reversals are, ongeologic timescales, globally synchronous, environmentally independent events.The geomagnetic timescale of Heirtzler et al. (1968) was one of the foundations of theplate tectonic revolution. They proposed a geomagnetic polarity timescale for the LateCretaceous to Recent based on a few long magnetic anomaly profiles. The evolution of thepolarity timescale since 1968 has involved two types of revisions: adjustments of the relativespacing of some anomalies and calibration of the polarity sequence in time (Cande and Kent,1992). Over the past forty years the pattern of normal and reversed polarities has beenextensively studied and most of its large-scale features for the past 200 million years are nowwell understood (Gradstein et al., 2005).Classic magnetic polarity reversal stratigraphy lacks the resolution necessary for thehigh-resolution (millennial-scale) climate studies being conducted today. This led to thedevelopment of high-resolution cryogenic magnetometers capable of measuring whole-coresamples or u-channel samples. This in turn led to the development of "composite sections" formarine sediment cores whereby multiple cores were taken at a single site and spliced together toprovide a complete stratigraphic section (Hagelberg et al., 1995).

PAGE 15

15Changes in the intensity of the Earth's magnetic field occur over much shorter timescalesthan polarity reversals. These changes can be measured in sedimentary cores to produce recordsof relative geomagnetic paleointensity. This is done by normalizing the natural remanentmagnetization by an artificial remanence to remove intensity changes due to changes inconcentration of magnetic material in the core. Records of relative paleointensity have beenshown to be globally correlative on millennial timescales for the last glacial cycle (Laj et al.2004).In attempting to understand the time-depth relationship in marine sediment cores andtherefore understand more about the Earth's climate and evolution my work covers three Oceans,the South Atlantic, the North Atlantic and the Pacific. Below is a summary of the work presentedin this dissertation. The nature of this work is collaborative and, as such, data provided by mycolleagues is included in this dissertation. Their contribution is acknowledged and clearlydetailed in the following summary.In Chapter 2 magnetostratigraphic and cyclostratigraphic results are presented for the 0-12 Ma interval from sites drilled during ODP Leg 198 to Shatsky Rise. Cyclic alternations in thepercentage of calcium carbonate, as shown by color reflectance data and gamma ray attenuationbulk density measured on the sediments, allowed astronomic calibration of the magneticstratigraphy from six ODP Sites. This chapter was published in the Scientific Results Volume forOcean Drilling Program (ODP) Leg 198 (Evans et al., 2005). Chapter 3 is a continuation of thiswork and has produced an integrated magnetobioand cyclostratigraphy from ODP Site 1208for the 1-12 Ma interval. Biostratigraphic data included in this chapter were provided byNicholas Venti, Mark Leckie (U. Massachusetts, foraminifer) and Paul Bown (UniversityCollege London, nannofossils).

PAGE 16

16In Chapters 4 and 5, I used sedimentary relative paleointensity records to correlatebetween cores collected on drift deposits in the North Atlantic. In Chapter 4, I present a study ofsediments from the Eirik drift for the 0-400 ka interval. Detrital layers identified within fourcores are placed in a paleointensity assisted chronostratigraphic framework. Environmentalmagnetic records from climatically sensitive regions such as the North Atlantic can provideinformation about changes in the strength of bottom currents and ice sheet dynamics both ofwhich are climatically sensitive. Oxygen isotope data used in this chapter were provided by JimWright and Lauren Nietzke (Rutgers University) and Claude Hillaire-Marcel (GEOTOP) (CoreMD99-2227). This chapter is under review in the journal Geophysics, Geochemistry andGeosystems. In Chapter 5 cores from the Gardar Drift provide records of changes in magneticgrain size over glacial/interglacial and stadial/interstadial cycles for the last 130 ka. Thesechanges are interpreted as changes in the speed of bottom currents forming the drift deposits overglacial/interglacial cycles and stadial/interstadial cycles. David Hodell (UF) provided oxygenisotope data in Chapter 5.In Chapter 6 a paleomagnetic study of IODP Site U1313 from the North Atlantic ispresented. The magnetic stratigraphy spans the interval from 2.5-6.3 Ma including the Gauss andGilbert chronozones. A relative paleointensity record for the Gauss and Gilbert chrons, is one ofonly a handful of such records for this time interval. Cycles in magnetic susceptibility haveallowed astronomic calibration of the ages of reversal boundaries.In 2001, my MS thesis consisted of a paleomagnetic study of ODP Site 1092 from theSouth Atlantic. Chapter 7 presents a revision of the composite depth scale from ODP Site 1092.X-Ray fluorescence (XRF) scanning data were provided by Thomas Westerhold (UniversityBremen). This chapter was published in Geophysical Journal International (Evans et al. 2004).

PAGE 17

17In Chapter 8 we use cyclic alternations in a stack of three oxygen isotope records (Paulsen et al.,in press) from ODP Site 1092 in the South Atlantic to astronomically tune the magneticstratigraphy from 9.3-11.2 Ma. This includes the long normal polarity chron C5n.2n and threeshort reverse polarity intervals within it identified by Evans and Channell (2003). It also includesa critical age tie-point from the Cande and Kent (1995) Geomagnetic Polarity Timescale (GPTS)at the base of C5n.2n. This chapter is under review at Earth and Planetary Science Letters.

PAGE 18

18CHAPTER 2LATE MIOCENE-HOLOCENE MAGNETIC POLARITY STRATIGRAPHY ANDASTROCHRONOLOGY FROM ODP LEG 198-SHATSKY RISEIntroductionShatsky Rise is a medium-sized large igneous province in the west-Central Pacific Ocean(Figure 2-1) and is possibly the oldest existing oceanic plateau. The rise consists of threeprominent topographic highs. Sites 1209, 1210, 1211 and 1212 were cored on the Southern High(Bralower, Premoli Silva, Malone et al., 2002). Eight sites on the Southern High of the rise weredrilled during Deep Sea Drilling Project (DSDP) and earlier Ocean Drilling Program (ODP) legs(Sites 47, 48, 49, 50, 305, 306, 577, and 810). Of these, ODP Sites 577 and 810 providedinterpretable Neogene magnetic stratigraphies.Sites 1207 and 1208, from the Northern and Central Highs, provided unexpectedlyexpanded late Miocene (12.5 Ma) to Holocene sequences. These locations had not been coredduring previous DSDP/ODP expeditions. The initial age model for all of the sites was based oncorrelation of the sequence of polarity zones to the geomagnetic polarity timescale (GPTS)(Cande and Kent, 1992, 1995). Mean sedimentation rates at the five sites vary from 1to 4cm/k.y. Latitude and longitude of the sites and basal ages of the Neogene sediments are given inTable 2-1. Neogene sediments at the sites consisted mostly of light gray to pale orangenannofossil oozes with varying amounts of clay, radiolarians, and diatoms. Magneticsusceptibility is low (< 2 x10-5 SI) at all the sites and shows a decreasing trend from theQuaternary to the late Miocene. Composite sections were constructed shipboard for four of thesites (1209, 1210, 1211, and 1212) using multi-sensor track (MST) data including magneticsusceptibility, gamma ray attenuation (GRA) bulk density, and reflectance data. Sites 1207 and1208 were not double-cored, and depths at these sites are in meters below sea floor (mbsf). The

PAGE 19

19magnetic stratigraphy from the six sites (1207, 1208, 1209, 1210, 1211, and 1212), was based onshipboard pass-through magnetometer measurements and discrete samples measured post-cruise.Sediments from five of the sites (1207, 1208, 1209, 1210, and 1211) showed a prominentcyclicity in reflectance data for parts of the sections, and this is the basis for the construction ofan astronomically tuned age model for the 0to 8-Ma interval. The astronomically calibratedpolarity timescale has been well established for the 0to 6-Ma interval (Shackleton et al., 1990,1995; Hilgen 1991a, 1991b). Hilgen (1991a, 1991b) produced his astronomically calibratedpolarity timescale for the 2to 5.23-Ma interval using sapropel occurrences and carbonatecontent in Mediterranean sections. These polarity chron ages were incorporated into the GPTS ofCande and Kent (1995).In this study we produced an astronomically calibrated magnetic reversal stratigraphy forthe 0to 8-Ma interval. This is in good agreement with Hilgen (1991a, 1991b) and Shackleton etal., (1995) in the 0to 6-Ma interval. In the 6-to 8-Ma interval, polarity chron ages are in betteragreement with the Shackleton et al. (1995) timescale, differing by up to ~200 k.y. from that ofHilgen et al. (1995) and the ATNTS 2004 of Lourens et al. (2004). This chapter was published inthe Scientific Results Volume for ODP Leg 198 (Evans et al., 2005).MethodsTwo types of paleomagnetic measurements were made on sediments collected during ODPLeg 198; pass-through measurements on half-cores and discrete sample measurements. Discretesample cubes (2cm x 2cm) were collected during Leg 198 to augment measurements using theshipboard pass-through magnetometer. Shipboard measurements on half-cores were made at 5-cm intervals. A total of 747 discrete samples were taken at 50-cm intervals. Discrete sampleswere collected from the center of the half-cores to avoid deformation at the outer edges of thecore. Magnetic measurements on the cubes were performed in the magnetically shielded room at

PAGE 20

20the University of Florida using a 2G-Enterprises cryogenic magnetometer. The samples werestep-wise alternating-field (AF) demagnetized using a D-Tech D2000 AF demagnetizer.Magnetization component directions were determined using the method of Kirschvink (1980),applied to the 20to 60 mT peak field demagnetization interval.The astrochronology developed for Sites 1207, 1208, 1209, 1210, and 1211 was based oncycles seen in reflectance data (L*) measured shipboard on a purpose-built track. Reflectance ofvisible light from soft sediment cores was measured using a spectrophotometer at 2.5-cmintervals and provided a high-resolution record of color variations for visible wavelengths (400-700 nm). L* reflectance represents "lightness" of the sediment which is usually controlled bychanges in percent carbonate.The initial age model for each site was based on correlation of the polarity zone sequenceto the timescale of Cande and Kent (1995). Power spectra using the Blackman-Tukey methodwith a Bartlett window from the Analyseries software of Paillard et al. (1996) indicate thepresence of obliquity and eccentricity peaks. The reflectance data were then tuned to theastronomic solutions for obliquity from Laskar et al. (1993). This allowed astronomicallycalibrated ages to be assigned to the polarity reversal boundaries at Sites 1207, 1208, 1209, 1210and 1211. Site 1212 was not included in the astrochronology, as it contains a hiatus at 4to 5-Ma.Magnetostratigraphic InterpretationSite 1207 is the only site that has been drilled on the Northern High of Shatsky Rise. Thesequence of sediment recovered was mostly Neogene in age (0-163.8 mbsf) underlain byCampanian and older oozes and cherts. The sediment consists of nannofossil ooze with diatoms,radiolarians, and clay in varying amounts (Bralower, Premoli Silva, Malone et al., 2002). Thesamples taken for paleomagnetic analysis were AF demagnetized in 5-mT steps up to either 50,

PAGE 21

2160, or 70 mT, depending on the intensity of the natural remanent magnetization (NRM). Lessthan 10% of the NRM remains after demagnetization at these peak fields, indicating a low-coercivity remanence carrier, most likely magnetite. Orthogonal projections of demagnetizationdata (Figure 2-2) show well-defined components for most of the Neogene section after removalof the steep drilling related overprint at peak AF fields of 20 mT. Maximum angular deviation(MAD) values are low for most of the section (< 10), indicating well-defined characteristicmagnetization components; however, some intervals, particularly the interval between 50-60mbsf (Figure 2-3), have slightly higher MAD values and less well-defined components. Theinterpretation of the magnetic stratigraphy from shipboard and discrete sample data can beaccomplished by polarity zone pattern fit to the GPTS (Cande and Kent, 1992, 1995) (Table 2-2).This pattern fit is satisfactory to the base of Subchron C5An.1n (Figures 2-3, 2-4). Below thepolarity zone equivalent to Subchron C5An.1n, recovery was intermittent and biostratigraphyindicates a hiatus with Campanian age sediments below (Bralower, Premoli Silva, Malone et al.,2002). Sedimentation rates average 1-2 cm/k.y. throughout the section with some slightly higher(3-4 cm/k.y.) rates in the late Pliocene and late Miocene (Figure 2-5A). Component declinationhas been corrected for each core using Tensor orientation data measured shipboard. The meaninclination in normal polarity zones for the Site is 57.8, close to the expected inclination of 56for a geocentric axial dipole at this site; however, reversed polarity intervals have a meaninclination of -51.1, shallower than expected. This can be attributed to shallowing of reversedpolarity directions by the steep downward-directed drilling overprint, shown clearly in theorthogonal projections (Figure 2-2A).Sites 1209, 1210, 1211, and 1212 are located on the southern high of Shatsky Rise(Figure 2-1). Multiple holes were drilled at each site and composite sections were constructed

PAGE 22

22using shipboard MST data. Discrete sample cubes were only collected from Holes 1209A,1210A, 1211A, and 1212A. The shipboard data from the pass-through magnetometer areconsistent between the different holes at each site and confirms the interpretation of the magneticstratigraphy (see Shipboard Scientific Party, 2002).As for Site 1207, orthogonal projections from discrete sample data show two components:-a steep downward drilling related overprint and well-defined characteristic components (Figure2-2 B, C, D, E). In most cases the drilling related overprint was easily removed in peak AF fieldsof 10-20 mT. Little of the natural remanent magnetization remained at peak fields of 60 mT.MAD values are generally <5 throughout the sections. The expected inclination for theSouthern Rise is 51; again, all the sites show slightly steeper than expected inclinations innormal polarity zones and shallower than expected inclination in reversed polarity zones. Themagnetostratigraphic age models indicate mean sedimentation rates between 1and 3 cm/k.y. formost of the Neogene (Figure 2-5 B, C, D, E).The polarity interpretation at Sites 1209, 1210, and 1211 is unambiguous back to SubchronC3Bn (Table 2-2) (Figures 2-6, 2-7, 2-8). Below this level, interpretation becomes difficult dueto decreasing sedimentation rates leading to a hiatus recognized at all sites between the upperMiocene, and Oligocene and older sequences. At Site 1212, a hiatus accounts for the intervalbetween 4 and 5 Ma (Chron C3), and the polarity interpretation can be accomplished toSubchron C4n.2n (Figure 2-9). This interpretation of the sequence of polarity zones is confirmedby the shipboard biostratigraphy. The interpretation of the polarity stratigraphy was carried outusing data measured shipboard augmented with discrete sample cubes. When themagnetostratigraphic data were placed on the composite depth scale, the reversal boundarieswere found to be consistent between holes, indicating that there is very little error in the depths

PAGE 23

23of polarity zone boundaries or in composite depth calculations. The magnetic measurementsmade shipboard do include a small amount of error due to the response function of the shipboardmagnetometer. The response function of the wide-access magnetometer used to measure half-cores is ~10 cm, resulting in a cm-scale uncertainty in the placement of the reversal boundaries.Site 1208 is located on the Central High of Shatsky Rise and also provided an expandedlate Miocene to Holocene section. The magnetic stratigraphy from Site 1208 will be presented inChapter 3.AstrochronologyCycles were visually identifiable in L* reflectance data from all six of the sites in thisstudy. For Sites 1209, 1210, and 1211, we worked with spliced composite records rather thandata from a single hole. Reflectance data were initially placed on the magnetostratigraphic agemodel based on the polarity timescale of Cande and Kent (1995). Power spectra for untunedsections of reflectance data placed on this age model consistently show a concentration of powerat orbital frequencies, particularly around the 41 k.y. obliquity cycle (Figure 2-10).The reflectance records were then tuned to the astronomical solution for obliquity fromLaskar et al. (1993), as this was the most visually identifiable cycle in the reflectance data andthe power spectra for different time intervals in all the sites showed a concentration of power atthe obliquity frequency (Figure 2-10). In constructing the astrochronological age model, weassume that there was no phase lag between the orbital forcing and the response. Forconvenience, the reflectance data were broken up into 1 Myr intervals when compared to theastronomical solution and each site was tuned independently. Cycles were readily apparent in thereflectance data for all sites, and tuning of the record required a minimum of adjustment of peaksin the reflectance data to the astronomical solution (Figures 2-11, 2-12, 2-13, 2-14).Astronomically tuned ages were calculated for polarity reversals in the 1 to 8-Ma interval at Site

PAGE 24

241207 (Table 2-3). At Site 1209, tuning was performed in the 1 to 7-Ma interval and at Sites 1210and 1211 in the 1 to 5-Ma interval. Site 1208 has also provided an astrochronological age modelfor the 1 to 6-Ma interval (Figure 2-12) and is included in Table 3. The tuned age models arecompared to each other (Table 2-3) and are compared with other recently publishedastrochronologies for this time period (Table 2-4). The output of a band-pass filter centered on41 k.y. is shown below the astronomical solution for obliquity and the raw reflectance data inFigures 2-11, 2-12, 2-13, 2-14, and 2-15.To test the validity of the timescale we used cross-spectral analysis performed using theBlackman-Tukey method and Analyseries software (Paillard et al., 1996). Coherence betweenthe reflectance data and the astronomical solution for obliquity was significant at all the sites,although the coherence values depend on which time interval is being examined. At Site 1207coherence was ~ 0.8 for the 1.2 to 1.8-Ma and 6.2 to 6.8-Ma intervals (Figure 2-16A). Thecoherence values at Site 1208 were > 0.8 for the entire 1 to 6-Ma interval. Sites 1209, 1210, and1211 also showed coherence values between 0.8 and 1 (Figure 2-16C, 2-16D, 2-16E).DiscussionComparison of the tuned ages for polarity reversal boundaries at the five sites in the 1.5to 2-Ma interval showed that polarity chron ages are in good agreement. For other time intervalsthere are some significant differences (more than an obliquity cycle) between sites (Table 2-3).Intervals with enhanced 41 k.y. power in reflectance data are considered more reliable (italics inTable 2-3). Site 1208 showed the strongest cyclicity, with Site 1207 also showing a clear signalin some intervals particularly the 2.1 to 2.7-Ma and 4.5 to 5-Ma intervals.During ODP Leg 138 to the eastern equatorial Pacific, 11 sites were drilled and most ofthem showed a prominent cyclicity in GRA density. Shackleton et al. (1995) used these cycles inGRA bulk density records to produce an orbitally-tuned age model for the 0 to 12.5-Ma interval.

PAGE 25

25They worked entirely in the time domain comparing smoothed GRA bulk density records withthe target record of summer insolation at 65N. In their tuning they assumed that no phase lagexisted between insolation and GRA bulk density controlled by proportion of SiO2 and CaCO3(high density), high carbonate content being associated with high Northern Hemisphereinsolation. Age control points were added to the data to align prominent groups of densitymaxima. The records were broken into 0.8-m.y. intervals for convenient viewing. Each site wastuned independently over the chosen time interval. Shackleton et al. (1995) found that someintervals in these records were more easily tuned than others, similar to results from Leg 198.Shackleton et al. (1995) noted that it was difficult to tune the 0to 1-Ma interval, which was alsothe case at four of the Leg 198 Sites (1208, 1209, 1210, and 1211). The 1to 2-Ma interval forthe Leg 138 sites carries a clear 41 k.y. obliquity cycle. For Leg 198 sites, the 1.2to 1.6-Mainterval also carries a very clear obliquity cycle (Figures 2-11, 2-12, 2-13, 2-14, 2-15). In the 2.4-to 2.6-Ma interval, a very strong obliquity cycle was observed at Site 846 (Leg 138), and thissame interval also carries a strong 41 k.y. signal at Sites 1209, 1210, and 1211. Comparisonbetween the Site 1207 age model and ages from Shackleton et al. (1995) indicate consistency forthe 1-to 8-Ma time interval (Table 2-4).Hilgen et al. (1995) developed an astronomical timescale for the interval from 3to 9.7-Mausing lithologic cyclicity seen in sedimentary sections in the Mediterranean. These sectionscomprise open marine sediments that alternate between carbonate-rich and carbonate-poor marlsor homogeneous marls and sapropels. The individual sapropels were related to precessionminima, and the clusters of sapropels to the 400-k.y. eccentricity cycles. In tuning the section,the target curve used was the 65N summer insolation curve. To obtain an astronomical age forthe youngest polarity reversal in the sequence, Hilgen et al. (1995) took the Shackleton et al.

PAGE 26

26(1995) age for the onset of Subchron C3An.2n of 6.576 Ma. They then matched the lithologiccycles in the section to the astronomical solution using the correlation of sapropel clusters toeccentricity. The age of the calibration point (6.576 Ma) had to be adjusted to 100 k.y. older toestablish a consistent correlation between sapropel clusters and eccentricity maxima. The agesfrom Hilgen et al. (1995) differ significantly with those from Leg 198 in the 6-to 8-Ma interval(Table 4). At the top of Subchron C3Bn the difference is more than 200 k.y. In the interval from7.2to 8.1-Ma, the difference is ~ 100 k.y. which is the amount of adjustment of the 6.576-Matie point used by Hilgen et al. (1995) for the age of the youngest polarity reversal in their section.ODP Site 926 on the Ceara Rise also produced an orbitally tuned timescale from 5to 14-Ma (Shackleton and Crowhurst, 1997). This timescale cannot be directly compared with the Leg198 timescale because of a lack of polarity reversals at Site 926. Backman and Raffi (1997) usedthe cyclostratigraphic age model from Site 926 to calibrate ages of the calcareous nannofossildatums for the late Miocene. These ages were then compared with the biomagnetochronologyfrom Site 853 (ODP Leg 138). The center of the peak in abundance of transitional morphotypesof Triquetrorhabdulus rugosus at Site 853 occurred 120-130 k.y. after the corresponding peak atSite 926. The age estimates of Hilgen et al. (1995) were then applied to the Site 853 data and thepeak center was found to coincide at Sites 853 and 926. Therefore, Backman and Raffi (1997)considered that the Hilgen et al. (1995) ages are more reliable in this interval than the ages ofShackleton et al. (1995).Lourens et al. (2004) have recalibrated the Miocene astronomic timescales of Shackletonand Crowhurst (1997) and Hilgen et al. (1995) using the astronomic solution of Laskar et al.(2004). For the last 13 Ma the retuning resulted in almost negligible changes in the ages ofreversal boundaries (Lourens et al., 2004). For the 6to 8Ma interval the ATNTS2004 is in

PAGE 27

27close agreement with that of Hilgen et al (1995) and therefore differs significantly with theresults of this study.ConclusionsFive sites from Shatsky Rise have produced high-quality magnetic stratigraphies from thelate Miocene to Holocene. Cycles identified in reflectance data from Sites 1207, 1208, 1209,1210, and 1211 have allowed astronomic calibration of the polarity reversal sequence from ~8Ma to present. The assumption that there is no phase lag between sedimentary cyclicity and theastronomical parameters allowed the cycles to be tuned to the astronomical solution forobliquity. Cross-spectral analysis on the tuned age model indicated high coherence between theastronomic solution and the reflectance data and confirms the reliability of the tuning. The agemodel has been compared with other published astrochronologies and is found to be in goodagreement with Hilgen (1991a, 1991b) (and, therefore, Cande and Kent [1995]) in the 1-to 6-Mainterval. In the 6-to 8-Ma interval the age model differs significantly from that of Lourens et al.(in press) and Hilgen et al. (1995) from the Mediterranean. It is in better agreement with theODP Leg 138 timescale of Shackleton et al. (1995) from the Pacific Ocean.

PAGE 28

28Table 2-1. Latitude, longitude, water depth, the oldest Neogene magnetic polarity chronidentified, and the basal age of the Neogene section. Site Latitude Longitude Waterdepth Basal Chron Basal Age(Ma) 1207 37.4287' N 162.0530'E 3100m C5An2n 12.184 1209 32 39.1001'N 158.3560'E 2387m C3Bn 7.091 1210 32 13.4123'N 158.5618'E 2573m C3Bn 7.091 1211 32 0.1300'N 157.9999'E 2907m C3Bn 7.091 1212 32 26.9000'N 157.7016'E 2682m C4n.2n 8.072

PAGE 29

29Table 2-2. Magnetostratigraphic age model for Sites 1207, 1209, 1210, 1211 and 1212. Polaritychron labels are according to Cande and Kent (1992, 1995). Ages of chrons are fromCande and Kent (1995). Depths are in meters below sea floor (mbsf) for Site 1207and meters composite depth (mcd) for Site 1209, 1210, 1211 and 1212. Chron Ma (CK95) 1207 (mbsf) 1209 (mcd) 1210 (mcd) 1211 (mcd) 1212 (mcd) C1n 0.00 0.00 0.00 0.00 0.00 0.00 base 0.780 12.35 11.28 14.89 8.00 11.95 C1r.1n 0.990 16.26 13.32 18.07 9.550 14.12 base 1.070 16.77 14.22 19.71 10.27 14.98 C2n 1.770 24.38 25.28 32.03 16.74 23.62 base 1.950 28.40 28.21 34.70 18.38 25.78 C2r.1n 2.197 29.73 37.68 23.70 26.95 base 2.229 30.25 38.09 30.08 27.78 C2An.1n 2.581 43.13 37.69 46.51 30.90 32.61 base 3.040 51.77 49.43 56.88 32.34 39.00 C2An.2n 3.110 53.25 58.52 33.98 39.81 base 3.220 56.79 60.37 37.37 41.67 C2An.3n 3.330 58.77 52.34 61.81 41.17 43.00 base 3.580 66.91 58.03 67.35 42.51 48.68 C3n.1n 4.180 80.23 66.22 75.36 43.94 base 4.290 83.34 68.23 77.62 44.66 C3n.2n 4.480 87.19 80.90 46.92 base 4.620 90.46 82.34 49.28 C3n.3n 4.800 92.39 73.24 83.78 50.72 base 4.890 94.32 74.08 85.42 51.70 C3n.4n 4.980 96.84 75.59 86.45 52.69 base 5.230 99.95 78.76 91.17 53.80 C3An.1n 5.894 105.73 82.94 94.05 54.29 54.44 base 6.137 106.77 84.95 95.69 55.15 56.17 C3An.2n 6.269 109.29 86.12 97.02 58.31 base 6.567 114.18 90.13 99.12 59.14 C3Bn 6.935 116.56 93.81 100.35 59.88 base 7.091 120.41 96.15 101.46 61.03 C3Br.1n 7.135 61.60 base 7.170 63.58 C3Br.2n 7.341 123.08 64.57 base 7.375 123.53 65.14 C4n.1n 7.432 125.16 65.39 base 7.562 126.05 66.95 C4n.2n 7.650 126.34 67.37 base 8.072 129.01 70.00 C4r.1n 8.225 130.79 base 8.257 132.27 C4An 8.699 134.50 base 9.025 136.72 C4Ar.1n 9.23 137.76 base 9.308 138.35 C4Ar.2n 9.580 140.28 base 9.642 140.72 C5n.1n 9.740 141.32 base 9.880 142.36 C5n.2n 9.920 142.65 base 10.949 151.40 C5r.1n 11.052 153.62 base 11.099 154.07 C5r.2n 11.476 155.11 base 11.531 155.40 C5An.1n 11.935 157.81 base 12.078 160.77 C5An.2n 12.184 161.76 base 12.401

PAGE 30

30Table 2-3. Comparison of astrochronological age models for sites 1207, 1208, 1209, 1210 and1211. Italics indicate the most reliable ages in intervals where the cyclicity inreflectance is best defined. In italics and brackets are the differences between tunedages and those of Cande and Kent (1995). Chron Ka(CK95)(Hilgen1991a,b) Site 1207Ka(difference) Site 1208Ka(difference) Site 1209Ka(difference) Site 1210Ka(difference) Site 1211Ka(difference) C1n 0 0 C1r.1r 780 776.7 (-3.3) C1r.1n 990 992.8 (2.8) C1r.2r 1070 1089.4 (19.4) 1073.9 (3.9) 1069.4 (-0.6) C2n 1770 1786.4 (16.4) 1776.2 (6.2) 1770.0 (0) 1777.8 (7.8) 1777.8 (7.8) C2r.1r 1950 1954.2 (4.2) 1948.7 (-1.3) 1975.4 (25.4) 1972.2 (22.2) 1972.2 (22.2) C2r.1n 2140 2095.7 (-44.3) 2133.5 (-6.5) C2r.2r 2150 2112.0 (-38) 2170.4 (20.4) C2An.1n 2581 2620.5 (39.5) 2564.7 (-16.3) 2550.3 (-30.7) 2642.7 (61.7) 2536.1 (-44.9) C2An.1r 3040 3042.5 (2.5) 3045.2 (5.2) 3032.0 (-8) 3032.9 (-7.1) 3022.2 (-17.8) C2An.2n 3110 3118.0 (8) 3105.8 (-4.2) 3114.9 (4.9) 3110.9 (0.9) C2An.2r 3220 3236.5 (16.5) 3229.8 (9.8) 3248.5 (28.5) 3242.3 (22.3) C2An.3n 3330 3354.5 (24.5) 3340.9 (10.9) 3361.4 (31.4) 3340.5 (10.5) 3352.8 (22.8) C2Ar 3580 3593.3 (13.3) 3599.6 (19.6) 3648.8 (68.8) 3597.5 (17.5) 3644.4 (64.4) C3n.1n 4180 4154.0 (-26) 4190.9 (10.9) 4172.5 (-7.5) 4182.8 (2.8) 4169.4 (-10.6) C3n.1r 4290 4262.5 (-27.5) 4351.9 (61.9) 4305.9 (43.4) 4305.9 (15.9) 4305.6 (15.6) C3n.2n 4480 4489.8 (9.8) 4523.6 (43.6) 4501.0 (21) 4457.9 (-22.1) C3n.2r 4620 4637.0 (17) 4683.8 (63.8) 4665.3 (45.3) 4589.3 (-30.7) C3n.3n 4800 4760.5 (-39.5) 4806.9 (6.9) 4809.0 (9) 4798.8 (-1.2) C3n.3r 4890 4857.3 (-32.7) 4880.9 (-9.1) 4880.9 (-9.1) 4891.2 (1.2) C3n.4n 4980 4972.5 (-7.5) 4991.8 (11.8) 4981.5 (1.5) 4973.3 (-6.7) 4950.7 (-29.3) C3r 5230 5245.4 (15.4) 5201.2 (-28.8) 5240.2 (10.2) C3An.1n 5894 5886.0 (8) 5952.7 (58.7) 5915.8 (21.8) C3An.1r 6137 6143.0 (6) 6073.9 (36.9) C3An.2n 6269 6241.5 (-27.5) 6318.3 (49.3) C3Ar 6567 6526.2 (-40.8) 6548.3 (-18.7) C3Bn 6935 6878.0 (-57) 6971.3 (35.3) C3Br.1r 7091 7095.8 (4.8) 7027.7 (-63.3} C3Br.1n 7135 C3Br.2r 7170 C3Br.2n 7341 7348.2 (8.2) C3Br.3r 7375 7388.3 (13.3) C4n.1n 7432 7453.5 (3.5) C4n.1r 7562 7540.9 (-21.1) C4n.2n 7650 7634.1 (15.9) C4r.1r 8072 8038.0 (-34)

PAGE 31

31Table 2-4. Astrochronological ages for Leg 198 compared to ages Hilgen (1991a, 1991b), Hilgenet al. (1995) and Shackleton et al, (1995). In italics and brackets are the differencesbetween Leg 198 tuned ages and Hilgen et al. (1995) and Shackleton et al. (1995). Chron Ka (CK95)Hilgen(1991a,b) Leg 198 Shackletonet al. (1995)(differenceto 198) Hilgen et al.(1995a)(differenceto 198) C1n 0 base 780 776.7 C1r.1n 990 992.8 base 1070 1089.4 C2n 1770 1786.4 base 1950 1954.2 C2r.1n 2140 2133.5 base 2150 2170.4 C2An.1n 2581 2564.7 base 3040 3042.5 3046 (3.5) C2An.2n 3110 3118.0 3131 (13) base 3220 3236.5 3233 (-3.5) C2An.3n 3330 3354.5 3331 (-23.5) base 3580 3593.3 3594 (0.7) C3n.1n 4180 4190.9 4199 (8.1) base 4290 4351.9 4316 (-35.9) C3n.2n 4480 4523.6 4479 (-44.6) base 4620 4683.8 4623 (-60.8) C3n.3n 4800 4806.9 4781 (-25.9) base 4890 4880.9 4878 (-2.9) C3n.4n 4980 4972.5 4977 (4.5) base 5230 5201.2 5232 (30.8) C3An.1n 5894 5952.7 5875 (-77.7) base 6137 6143.0 6122 (-21) C3An.2n 6269 6241.5 6256 (14.5) base 6567 6526.2 6555 (28.8) 6677 (150.8) C3Bn 6935 6878.0 6919 (41) 7101 (223) base 7091 7095.8 7072 (-23.8) 7210 (114.2) C3Br.1n 7135 7256 base 7170 7301 C3Br.2n 7341 7348.2 7455 (106.8) base 7375 7388.3 7492 (103.7) C4n.1n 7432 7453.5 7406 (-47.5) 7532 (78.5) base 7562 7540.9 7533 (-7.9) 7644 (103.1) C4n.2n 7650 7634.1 7618 (-16.1) 7697 (62.9) base 8072 8038.0 8027 (-11) 8109 (71)

PAGE 32

32 Figure 2-1. Bathymetric map of Shatsky Rise showing the location of sites drilled during ODPLeg 198.

PAGE 33

33 Figure 2-2. Representative orthogonal projections of AF demagnetization data from (A) Site1207, (B) Site 1209, (C) Site 1210, (D) Site 1211 and (E) Site 1212. Low AFdemagnetization treatment and the high treatment are given, as is mbsf or mcd of thesample. Open circles represent the vector end-point projection on the vertical plane,while closed circles represent the vector endpoint projection on the horizontal plane.

PAGE 34

34 Figure 2-3. Site 1207 component inclination and declination from discrete samples (opensquares) for 0-80 meters. Inclination and rotated declination from the shipboard pass-through magnetometer after AF demagnetization at peak fields of 20 mT (gray line).Chrons are labeled according to Cande and Kent (1992). Black indicates normalpolarity, white reversed polarity. Also shown are the MAD values calculated fordiscrete sample data (after Kirschvink, 1980).

PAGE 35

35 Figure 2-4. Site 1207 component inclination and declination from discrete samples (opensquares) for 80-160 meters. Inclination and rotated declination from the shipboardpass-through magnetometer after AF demagnetization at peak fields of 20 mT (grayline). Chrons are labeled according to Cande and Kent (1992). Black indicates normalpolarity, white reversed polarity. Also shown are the MAD values calculated fordiscrete sample data (after Kirschvink, 1980).

PAGE 36

36 Figure 2-5. Interval sedimentation rates and age versus depth for the initial age model at a) Site1207, b) Site 1209, c) Site 1210, d) Site 1211 and e) Site 1212.

PAGE 37

37 Figure 2-6. Site 1209 component inclination and declination from discrete samples (opensquares). Inclination and rotated declination from the shipboard pass-throughmagnetometer after AF demagnetization at peak fields of 20 mT (gray line). Chronsare labeled according to Cande and Kent (1992). Black indicates normal polarity,white reverse polarity. Also shown are the MAD values calculated for discrete sampledata (after Kirschvink, 1980).

PAGE 38

38 Figure 2-7. Site 1210 component inclination and declination from discrete samples (opensquares). Inclination and rotated declination from the shipboard pass-throughmagnetometer after AF demagnetization at peak fields of 20 mT (gray line). Chronsare labeled according to Cande and Kent (1992). Black indicates normal polarity,white reverse polarity. Also shown are the MAD values calculated for discrete sampledata (after Kirschvink, 1980).

PAGE 39

39 Figure 2-8. Site 1211 component inclination and declination from discrete samples (opensquares). Inclination and rotated declination from the shipboard pass-throughmagnetometer after AF demagnetization at peak fields of 20 mT (gray line). Chronsare labeled according to Cande and Kent (1992). Black indicates normal polarity,white reverse polarity. Also shown are the MAD values calculated for discrete sampledata (after Kirschvink, 1980).

PAGE 40

40 Figure 2-9. Site 1212 component inclination and declination from discrete samples (opensquares). Inclination and rotated declination from the shipboard pass-throughmagnetometer after AF demagnetization at peak fields of 20 mT (gray line). Chronsare labeled according to Cande and Kent (1992). Black indicates normal polarity,white reverse polarity. Also shown are the MAD values calculated for discrete sampledata (after Kirschvink, 1980).

PAGE 41

41 Figure 2-10. Power spectra from a) Site 1207, b) Site 1208 c) Site 1209, d) Site 1210, and e) Site1211 for reflectance data placed on a Cande and Kent (1995) age model.

PAGE 42

42 Figure 2-11. The astronomical solution for obliquity (Laskar et al., 1993) compared with tunedL* reflectance data from Site 1207 for the 0-8 Ma interval. The reflectance datafiltered using a band-pass filter centered on 41kyrs is shown in the lower part of eachframe. Black indicates normal polarity and white reverse polarity. Heavy line on L*reflectance data indicates intervals where the cyclicity is best developed.

PAGE 43

43 Figure 2-12. The astronomical solution for obliquity (Laskar et al., 1993) compared with tunedL* reflectance data from Site 1208 for the 1-6 Ma interval. The reflectance datafiltered using a band-pass filter centered on 41kyrs is shown in the lower part of eachframe. Black indicates normal polarity and white reverse polarity. Heavy line on L*reflectance data indicates intervals where the cyclicity is best developed.

PAGE 44

44 Figure 2-13. The astronomical solution for obliquity (Laskar et al., 1993) compared with tunedL* reflectance data from Site 1209 for the 1-7 Ma interval. The reflectance datafiltered using a band-pass filter centered on 41kyrs is shown in the lower part of eachframe. Black indicates normal polarity and white reverse polarity. Heavy line on L*reflectance data indicates intervals where the cyclicity is best developed.

PAGE 45

45 Figure 2-14. The astronomical solution for obliquity (Laskar et al.,, 1993) compared with tunedL* reflectance data from Site 1210 for the 1-5 Ma interval. The reflectance datafiltered using a band-pass filter centered on 41kyrs is shown in the lower part of eachframe. Black indicates normal polarity and white reverse polarity. Heavy line on L*reflectance data indicates intervals where the cyclicity is best developed.

PAGE 46

46 Figure 2-15. The astronomical solution for obliquity (Laskar et al., 1993) compared with tunedL* reflectance data from Site 1211 for the 1-5 Ma interval. The reflectance datafiltered using a band-pass filter centered on 41 kyrs is shown in the lower part of eachframe. Black indicates normal polarity and white reverse polarity. Heavy line on L*reflectance data indicates intervals where the cyclicity is best developed.

PAGE 47

47 Figure 2-16. Cross-spectral analysis from a) Site 1207, b) Site 1208 c) Site 1209 d) Site 1210 ande) Site 1211. Power spectra for the tuned reflectance data (black) and for theastronomical solution for eccentricity and obliquity (red). Coherence values betweenthe astronomical solutions and reflectance data are shown below.

PAGE 48

48CHAPTER 3INTEGRATED NEOGENE MAGNETIC, CYCLE AND BIOSTRATIGRAPHY FROM ODPSITE 1208 (SHATSKY RISE, PACIFIC OCEAN)IntroductionODP Site 1208 was drilled in 2001 on Shatsky Rise, a large igneous province in the NWPacific Ocean. A single hole drilled at the site has produced a magnetic polarity stratigraphy forthe 0-12 Ma interval. Sedimentation rates decrease from 4-5 cm/kyr in the Brunhes andMatuyama chrons to less than 1 cm/kyr at the base of the studied section. A revised planktonicforaminifer biostratigraphic zonation has been developed for the NW Pacific Ocean using theseventeen most isochronous foraminifer datums. This scheme has been integrated withnannofossil events, and with the magnetic stratigraphy. Cycles in the reflectance (L*) can bematched to astronomic solution for obliquity allowing astronomic calibration of polarity chronboundaries, and planktonic foraminifer and calcareous nannofossil datums. Astronomic ages forpolarity chron boundaries are consistent with the ATNTS2004 timescale (Lourens et al., 2004) inthe 1-5.2 Ma interval, however, between 5.2 and 6.2 Ma, astronomic ages from Site 1208 differsignificantly (by ~ 300 kyr) from the ATNTS2004 timescale.As polarity reversals can be considered globally synchronous, the integration of polaritychron boundaries and biostratigraphies has become a powerful means of calibratingbiostratigraphic zonations, and determining synchroneity of biostratigraphic events (seeBerggren et al., 1995a,b). Early work on Neogene foraminifer biostratigraphy in the NorthPacific Ocean on Deep Sea Drilling Project (DSDP) Sites 173, 296 and 310 (Keller, 1979a,b,c)was augmented by foraminifer and nannofossil work on ODP Leg 138 in the eastern equatorialPacific (Raffi and Flores, 1995; Shackleton et al., 1995a). ODP Leg 138 biostratigraphies wereintegrated into well-defined magnetic stratigraphies (Schneider, 1995) and cyclostratigraphiesbased on gamma ray attenuation (GRA) bulk density data. Correlation of GRA bulk density

PAGE 49

49cycles to astronomical calculations for solar insolation provided robust age models for LateMiocene to Recent sediments (Shackleton et al., 1995b). The ODP Leg 138 age models wereamong the first astrochronologies developed for the Late Miocene to Quaternary, and hence thebiostratigraphies generated from ODP Leg 138 sites were rather precisely calibrated. Berggrenet al. (1995a,b) incorporated ages and bio-magnetostratigraphic data from ODP Leg 138 intotheir review of bio-magnetostratigraphic correlations for the Cenozoic and Quaternary.ODP Site 1208 offers the opportunity to refine biomagnetostratigraphic correlations for thelate Neogene. The attributes of ODP Site 1208 include: good preservation of foraminifers andcalcareous nannofossils, relatively high sedimentation rates compared to ODP Leg 138 sites, anda robust age model based on magnetic polarity stratigraphy and correlation of reflectance data toastronomical solutions.The study of ODP Site 1208 is a continuation of the work presented in Chapter 2, whichdeals largely with ODP Leg 198 sites other than Site 1208. An initial astrochronology for ODPSite 1208 (Chapter 2, Evans et al., 2005) was based on correlation of the shipboard reflectance(L*) data (Shipboard Scientific Party, 2002a) to the astronomical solution for obliquity of Laskaret al. (1993). Here, we update the Site 1208 astrochronology using the new astronomic solutionsof Laskar et al. (2004), present the Site 1208 magnetostratigraphy, foraminiferal and nannofossilbiostratigraphy, and link these stratigraphies to the new astrochronology. The recalibration ofthe Site 1208 age model makes little difference to the chronology presented in the previouschapter (and in Evans et al., 2005) because the astronomic solutions in the 0-12 Ma interval donot change significantly in Laskars two calculations (Laskar et al., 1993; 2004).Today, Shatsky Rise (Figure 3-1) lies in a subtropical water mass toward the north end of awarm-water mass known as the Kuroshio Extension Current (Shipboard Scientific Party, 2002b).

PAGE 50

50North of the Northern High of Shatsky Rise (Figure 3-1) lies a significant front, a transitionregion between subtropical and subarctic water masses. The transition zone waters are derivedfrom off the coast of northern Japan, where the cold, nutrient-rich Oyashio Current mixes withthe warm, nutrient-poor Kuroshio Extension Current. Middle Miocene calcareous planktonassemblages are rather uniform and diverse across Shatsky Rise and display warm, subtropicalaffinities. Since the Late Miocene, however, a faunal and floral gradient has been establishedacross Shatsky Rise (Shipboard Scientific Party, 2002b). Calcareous plankton assemblagesprogressively loose their warm-water taxa along a traverse from south to north across theShatsky Rise. At Sites 1207 and 1208 (Figure 3-1), there is a marked decrease in diversity inassemblages that assume temperate (occasionally cold-temperate) affinities, relative to sitesfurther south. The changes in calcareous plankton assemblages are paralleled by a progressivedecrease in calcareous preservation from north to south (Shipboard Scientific Party, 2002b).One of the most noticeable features of the upper Miocene through Pleistocene sectionsrecovered at Shatsky Rise is the decimeterto meter-scale cycles between darker and lighterlithologies. The darker-colored intervals, in general, contain larger amounts of well-preservedbiosiliceous material, and contain calcareous plankton assemblages that have cold-wateraffinities and have undergone relatively enhanced dissolution. Calcareous plankton preservationis enhanced in the light-colored layers that are poorer in diatoms and represent warmer-waterintervals when Site 1208 was located in a subtropical water mass, similar to the situation at Site1208 today and for the Southern High through most of the Neogene (Shipboard Scientific Party,2002b).

PAGE 51

51Site Location and LithologyODP Site 1208 is located in 3346m of water on the Central High of Shatsky Rise (Figure3-1). The Central High of the Rise had not been drilled prior to ODP Leg 198, and thesedimentary sequence at the site revealed ~260 m of Upper Miocene to Recent sediments with~60m of more condensed Lower and Middle Miocene below. A total of 314.17 m of Neogeneage sediment was recovered at the site with an average recovery of 95%. The Upper Miocene toRecent section is composed of nannofossil ooze and nannofossil clay with diatoms andradiolarians, and an average carbonate content of 53% (Shipboard Scientific Party, 2002b).Since 3 Ma, the average sedimentation rates were 4.2 cm/kyr. Prior to 3 Ma, sedimentation ratesdecrease progressively reaching 1 cm/kyr at ~8 Ma. The character of the seismic reflectionrecord at the site, along with the relatively high sedimentation rate that prevailed during thePliocene-Pleistocene, suggests that the stratified lens of sediment at the site constitutes a driftdeposit formed by current redistribution of sediment that settled on the Central High (ShipboardScientific Party, 2002b). The sediment drift deposits at Site 1208 are somewhat similar to thosedrilled along the Meiji Seamount during ODP Leg 145 in that both sections comprise fine-grained sediment devoid of sedimentary structures other than bioturbation (Rea et al., 1993).Magnetic StratigraphyMagnetic measurements on half cores from Site 1208, using the shipboard pass-throughmagnetometer, revealed an unambiguous magnetic stratigraphy, ranging in age from Recent toUpper Miocene (Figures 3-2, 3-3 and 3-4). The shipboard data are based on a singledemagnetization step (20 mT). This abbreviated treatment was necessary to preserve thesediment magnetization for later shore-based study, and to maintain core-flow through theshipboard core laboratory during the cruise. These shipboard data are supported using discretesample cubes (7cm3) collected from the working halves of cores, which were measured at the

PAGE 52

52University of Florida. The discrete sample cubes were AF demagnetized in 5mT increments upto peak fields of 80 mT. A steep drilling related overprint was removed by 20 mT peak field(Figure 3-5), and the primary magnetization was defined using the standard least squares method(Kirschvink, 1980), giving low maximum angular deviation (MAD) values indicating welldefined component magnetizations (Figures 3-2, 3-3 and 3-4).An initial age model and initial estimate of interval sedimentation rates were calculatedusing the magnetostratigraphy and the geomagnetic polarity timescale (GPTS) of Cande andKent (1995) (Figure 3-6a). Sedimentation rates decrease down section averaging 4-5 cm/kyr inthe Pleistocene, 3.5-4 cm/kyr in the Pliocene and 1-2 cm/kyr in the Miocene. The duration of theReunion subchron given in the Cande and Kent (1995) GPTS (10 kyr) causes a large increase inthe sedimentation rates in the polarity zone correlative to the Reunion subchron (Figure 3-2).Using a revised age and duration for the Runion subchron (Channell et al., 2003), sedimentationrates in the polarity zone correlative to the Reunion subchron are reduced to ~4 cm/kyr inkeeping with surrounding sedimentation rates.Numerous excursions can be identified in the shipboard magnetic stratigraphy particularlyin the Matuyama Chron, one of which (at 103 mbsf) is confirmed by a single discrete samplecorresponding to an age of 2.283 Ma. Channell et al., (2002) identified seven excursions in theMatuyama chron at ODP Site 983 in the North Atlantic that have been labeled: Santa Rosa (932ka), C1r.1n.1r (1048 ka), Punaruu (1115 ka), Bjorn (1255 ka), Gardar (1472-1480 ka), Gilsa(1567-1575 ka), and C2r.1r.1n (1977 ka).At ODP Site 1208, shallow inclinations are seen in shipboard data that appear to becorrelative to Santa Rosa (950 ka), Punaruu (1123 ka), Gardar (1450 ka), Gilsa (1522 ka), andC2r.1r.1n (1976 ka) (Figure 3-2). An interval of shallow inclination at 896 ka at Site 1208 may

PAGE 53

53correspond to the Kamikatsura excursion that originates from the work of Maenaka (1983).Three intervals of shallow inclination are also noted in the Brunhes chron with ages of 134, 193and 262 ka close to the published ages for the Blake excursion (120 ka), Iceland Basin excursion(189 ka) and 8! (260 ka) of Lund et al. (2001) (Figure 3-2). One potential excursion is noted inthe Gauss chron (Figure 3-3) and three potential excursions in the Gilbert chron (Figure 3-5).The ages for the Santa Rosa, Blake, Iceland Basin and 8! are calculated by assuming constantsedimentation rates within the Brunhes and subchron C1r.1r. Ages for other excursions are atSite 1208 are calculated from the astronomic age model described below. All the postulatedexcursions in the Site 1208 record should be regarded with some caution as they are based on asingle demagnetization step (20 mT peak field) from shipboard data.Cycle StratigraphyShipboard gamma ray attenuation bulk density data and L* reflectance data (ShipboardScientific Party, 2002), show a prominent cyclicity in the 1-6 Ma interval, that, based on theinitial age model has a period close to 41 kyr (see Chapter 2, Evans et al., 2005). Using theastronomical solutions of the Laskar et al. (2004), the L* reflectance data was tuned to obliquityby matching the L* output of a filter centered on 41 kyr to the orbital solution for obliquity(Figure 3-7). The resulting interval sedimentation rates for the 1-6 Ma interval are given inFigure 3-6b. The tuned ages for reversal boundaries, based on this match, are given in Table 3-2.The astronomical calibration of Site 1208 presented here is a recalibration of the astronomicaltimescale of Evans et al., (2005) using the updated astronomical solutions of Laskar et al. (2004).The recalibration to the new astronomic solutions resulted in little change to the astronomic agesfor Site 1208 relative to those given by Evans et al. (2005). Comparison of polarity reversal ageswith other timescales, and with results from IODP Site U1313 (Chapter 6), indicates close

PAGE 54

54agreement with differences < ~60 kyrs between 2.6 Ma and 5 Ma (Table 3-2). Beyond 5.2 Ma,the differences with respect to other timescales increase to over 100 kyr. In the 5.5-6 Ma interval,the polarity reversal ages from Site 1208 are closest to those of Shackleton et al. (1995b) fromODP Leg 138 (equatorial Pacific). The largest discrepancy beyond 5.2 Ma is with ATNTS2004timescale (Lourens et al., 2004) where the difference in ages is ~ 300 kyrs. The ATNTS2004timescale uses the work of Hilgen et al. (1995) from the Mediterranean in this interval.Calcareous NannofossilsCalcareous nannofossils were semi-quantitatively analyzed using smear slides andstandard light microscope techniques (Bown and Young, 1998). The following abundance andpreservation categories were used: Species abundance: abundant: >10 specimens per field ofview (FOV), common: 1 specimens per FOV, few: 1 specimen per 210 FOV, rare: 1specimen per 11 FOV. Total nannofossil abundance: abundant: >10%, common: 1%%,few: 0.1%%, rare: <0.1%, barren and questionable occurrence. Nannofossil preservation:good, moderate, poor (See range chart of Bown, 2005). All core catcher samples were examinedand ~60 other samples collected through the Late Miocene to Recent section. Biostratigraphy isdescribed with reference to the zonal scheme of Bukry (1973, 1975; zonal code numbers CN andCP added and modified by Okada and Bukry, 1980) for Cenozoic calcareous nannofossilbiostratigraphy.The middle MioceneHolocene section yielded a beautiful succession of rich and abundantnannofossil assemblages. Preservation improved up-section but was also dependent upon whichpart of the light/dark sedimentary cycle was sampled. The darker, diatom-rich intervals yieldedmore poorly preserved nannofossil assemblages (Shipboard Scientific Party, 2002). The Neogenenannofossil biostratigraphy indicates a relatively complete stratigraphy for the Pliocene-Pleistocene (Figure 3-8) and Miocene (Figure 3-9), with all nannofossil zones from CN5 through

PAGE 55

55CN15 identified by their primary zonal fossils (Figure 3-10). Calcareous nannofossil range chartsare shown in Bown (2005). Zones CN1CN5 could not be easily distinguished because of theabsence of the marker species Sphenolithus belemnos, Helicosphaera ampliaperta, andDiscoaster kugleri. In addition, a number of CN subzones could not be recognized due to theabsence of D. kugleri (Subzone CN5b), Discoaster loeblichii, Discoaster neorectus (SubzoneCN8b), and Amaurolithus amplificus (subdivisions within Zone CN9) and an anomalously lowlast occurrence (LO) of Triquetrorhabdulus rugosus (Subzone CN10b) (Bown, 2005).The astronomically calibrated ages of Pliocene to Quaternary calcareous nannofossildatums from Site 1208 (Table 3-3) are generally consistent with ages from Berggren et al.(1995b), that are based largely on work from the Mediterranean (Rio et al., 1990). The LO ofDiscoaster brouweri, however, differs significantly from Berggren et al. (1995b) in both age andcorrelative polarity chron (Table 3-3). The age of 1.95 Ma given by Berggren et al. (1995b),correlative to the onset of the Olduvai subchron, is based on correlation to Deep Sea DrillingProject (DSDP) Site 606 in the North Atlantic (Backman and Pestiaux, 1987).At Site 1208, the FO of Discoaster berggrenii in the Late Miocene is ~ 0.5 Myrs youngerthan the age reported in Berggren et al. (1995a). This age is based on correlation to polaritychron C4r.2r from ODP Leg 138. The age is more consistent with that seen at DSDP Site 608where the datum is correlated to polarity chron C4n (Ruddiman et al., 1987). The FO ofDiscoaster hamatus is a controversial datum (Berggren et al., 1995a) that has very inconsistentcorrelation to polarity chrons regardless of latitude. In ODP Leg 138 sites, it is correlative tosubchron C5n.2n, as at Site 1208. The FO of Catinaster coalitus is another controversial datumthat, at ODP Site 1208, is correlated to subchron C5n.2n similar to the correlation at ODP Leg138 sites. Berggren et al. (1995a) give an age of 10.8 Ma for the FO of Coccolithus

PAGE 56

56miopelagicus, 200 kyrs younger than the age from Site 1208 (Table 3-3), however the correlationof this datum to polarity subchron C5r.1r at Site 1208 is consistent with the correlation at DSDPSite 608.Planktonic Foraminifera158 samples were analyzed for planktonic foraminifers at ~1.5 m intervals, together withcore-catcher samples (from the base of each core) collected shipboard from the 320-m-thickupper Neogene section at ODP Site 1208. The samples were soaked in a slightly basic solution,shaken, washed over a 63 !m sieve and dried at 60C. Specimens of planktic and benthicforaminifers were picked from the >125 !m fraction. Specimens of all recognizable plankticspecies were identified following the classic taxonomies of Kennett and Srinivasan (1983), Bolliand Saunders (1985), Jenkins (1985), and Iaccarino (1985). Shipboard and shore-basedoccurrence tables were combined to determine a planktic foraminifer biostratigraphy.Occurrence estimates were based on the following percentages: Rare=1%, rare to few=3%,few=5%, few to common=8%, common=10%, common to abundant=15%, abundant =>20%,Planktonic foraminiferal abundance varies from abundant to common through thePleistocene and upper Pliocene but declines in the Miocene to few to rare relative to siliceousmicrofossils and clay. Temperate-water species dominate many of the Neogene planktonicforaminiferal assemblages at Site 1208 (Shipboard Scientific Party, 2002b).The magnetostratigraphically-interpolated ages for many foraminiferal datums on ShatskyRise differed significantly from those reported from the southwest Pacific, due to regionalmigration patterns. Application of zonal schemes proposed for the southwest Pacific (Jenkins,1985) and the mid-latitudes were complicated by unexpected changes in the sequence offoraminiferal datums observed at Shatsky Rise. A revised temperate foraminifer biostratigraphy

PAGE 57

57for the late Neogene uses seventeen of the most isochronous foraminiferal datums at ShatskyRise as zonal markers (shown in Figure 3-11).Discrepancies between published ages for planktonic foraminifer datums (Berggren et al.,1995a,b; Lourens et al., 2004) and those identified at Site 1208 are large in some cases. (Tables3-4 and 3-5). The majority of the magneto-biostratigraphic correlations used in Berggren et al.,(1995a,b) and Lourens et al. (2004) are from the Mediterranean (Hilgen, 1990), South Atlantic(Hodell and Kennett, 1987) or South Pacific (Srinivasan and Sinha, 1993). The comparison ofthe ages of the planktic foraminifer datums is affected by regional differences and varying zonalschemes (Tables 3-4 and 3-5).The LO of Gr. tosaensis at ODP Site 1208 is at 0.292 Ma, significantly younger than theage of 0.65 Ma given by Berggren (1995b). This datum is taken from the work of Berggren et al.(1985) and Srinivasan and Sinha (1993) from the southern Pacific and Indian Oceans. The LOdatum of Gr. punticulata has an age of 1.882 Ma at Site 1208, however, an age of 2.41 Ma wasobtained at DSDP Site 607 (North Atlantic) where it is correlative to polarity subchron C2An.2n.The FO of Gr. truncatlinoides occurs at the same stratigraphic level as the FO of Gr.toseanis at ODP Site 1208. Following Berggren et al. (1995b), the FO of Gr. toseanis has an ageof 3.35 Ma. At Site 1208, the astronomically calibrated age of the event is 2.015 Ma. The Site1208 age for this datum is more consistent with the astronomically calibrated age for the FO ofGr. truncatlinoides of 2.39 Ma from ODP Leg 138 in the eastern equatorial Pacific (Shackletonet al., 1995a). The LO of Gr. margaritae was assigned an age of 3.85 Ma in the ATNTS2004(Lourens et al., 2004) from ODP Sites 925 and 926 from Ceara Rise. The astronomic age for thedatum at Site 1208 is 3.761 Ma.

PAGE 58

58ConclusionsODP Site 1208 has produced a clear magnetic stratigraphy for the 0-12 Ma interval withsedimentation rates in the Brunhes and Matuyama chrons varying in the 4-5 cm/kyr range. Thesesedimentation rates are some of the highest sedimentation rates seen in this interval in pelagicsediments from the midand low latitude Pacific Ocean. This anomalously high sedimentationrate appears to be due to formation of a drift-type deposit on the Central High of Shatsky Rise.The relatively high sedimentation rates have allowed identification of polarity excursions in theMatuyama Chron that have not been previously identified in sediments from the Pacific Ocean.It is important to stress that these excursions are identified in shipboard pass-through magneticdata, and are not based on identification of magnetization components. For this reason, theratification of these excursional directions must await further (u-channel) studies of thesesediments.Reflectance (L*) cycles identified in the sediments have allowed astronomic calibration ofreversal boundaries and biostratigraphic datums, by correlation of L* reflectance data to theastronomic solution for obliquity (Laskar et al., 2004). Calcareous nannofossil biostratigraphy islargely consistent with the most recent review of bio-magnetostratigraphic correlations for thistime interval (Berggren et al., 1995a, b). Based on the correlation of planktonic foraminiferdatums to the magnetic stratigraphy at Site 1208, a new planktonic foraminifer zonation for thenorthwest Pacific Ocean has been developed that can be precisely correlated to polarity chronsand astronomically calibrated ages.

PAGE 59

59Table 3-1. Depths of reversal boundaries from ODP Site 1208. Chrons are labeled according toCande and Kent (1992). Ages for polarity chrons are from Cande and Kent (1995)and Channell et al., (2003). Chron Ma (CK95) mbsf Chron Ma (CK95) mbsf C1n 0 0 C3Br.1n 6.946 base 0.78 42.92 base 6.981 C1r.1n 0.99 52.57 C3Br.2n 7.153 base 1.07 55.85 base 7.187 240.33 C1r.2r.1n 1.201 61.18 C4n.1n 7.245 241.08 base 1.211 61.67 base 7.376 241.83 C2n 1.77 85.01 C4n.2n 7.464 242.95 base 1.95 92.81 base 7.892 250.78 C2r.1n* 2.115 99.86 C4r.1n 8.047 251.71 base* 2.153 101.01 base 8.079 252.46 C2An.1n 2.581 119.45 C4An 8.529 256 base 3.04 137.8 base 8.861 260.66 C2An.2n 3.11 140.64 C4Ar.1n 9.069 262.53 base 3.22 144.58 base 9.146 264.02 C2An.3n 3.33 147.76 C4Ar.2n 9.428 265.69 base 3.58 156.88 base 9.491 268.12 C3n.1n 4.18 172.4 C5n.1n 9.592 269.05 base 4.29 176.47 base 9.735 271.85 C3n.2n 4.48 182.51 C5n.2n 9.777 base 4.62 185.46 base 10.834 282.1 C3n.3n 4.8 189.28 C5r.1n 10.94 287.14 base 4.89 191.01 base 10.989 287.69 C3n.4n 4.98 194.09 C5r.2n 11.378 290.49 base 5.23 200.04 base 11.434 291.42 C3An.1n 5.894 216.47 C5An.1n 11.852 292.17 base 6.137 221.51 base 12 294.03 C3An.2n 6.269 222.25 C5An.2n 12.108 298.69 base 6.567 231.39 base 12.333 299.63 C3Bn 6.935 235.31 base 7.091 240.34 age from Channell et al. (2003)

PAGE 60

60Table 3-2. Astronomically calibrated ages for reversal boundaries from ODP Site 1208 comparedto ATNTS2004 (Lourens et al., 2004), Cande and Kent (1995), IODP Site U1313(Evans et al., in preparation, Chapter 6), Hilgen et al., (1995) and ODP Leg 138(Shackleton et al., 1995b). Differences between Site 1208 ages and published ages aregiven in parentheses. Chron 1208tuned age(Ma) CK95 (Ma) Hilgen et al.(1995) (Ma) ATNTS2004 (Ma) ODP Leg 138 IODP SiteU1313Chapter 6 C1n base 0.780 C1r.1n 0.990 base 1.062 1.070 (-0.008) 1.072 (0.01) C1r.2r.1n 1.158 1.201 (0.043) 1.173 (0.015) base 1.167 1.211 (0.044) 1.185 (0.018) C2n 1.763 1.770 (-0.007) 1.785 (0.022) 1.778 (0.015) base 1.944 1.950 (-0.006) 1.942 (-0.002) 1.945 (0.001) C2r.1n 2.204 2.140 (0.064) 2.129 (-0.075) 2.128 (-0.076) base 2.214 2.150 (0.064) 2.149 (-0.065) 2.148 (-0.066) C2An.1n 2.616 2.581 (0.035) 2.582 (-0.34) 2.581 (-0.035) 2.600 (-0.016) 2.616 (0) base 3.048 3.040 (0.008) 3.032 (0.016) 3.032 (-0.016) 3.046 (-0.002) 3.074 (0.026) C2An.2n 3.091 3.110 (-0.019) 3.116 (0.025) 3.116 (0.025) 3.131 (0.04) 3.153 (0.062) base 3.207 3.220 (-0.013) 3.207 (0) 3.207 (0) 3.233 (0.026) 3.268 (0.061) C2An.3n 3.350 3.330 (0.020) 3.330 (-0.02) 3.330 (-0.02) 3.331 (-0.019) 3.346 (-0.004) base 3.584 3.580 (0.004) 3.569 (-0.015) 3.596 (0.012) 3.594 (0.01) 3.549 (-0.035) C3n.1n 4.164 4.180 (-0.016) 4.188 (0.024) 4.187 (0.023) 4.199 (0.035) 4.144 (-0.02) base 4.307 4.290 (0.017) 4.300 (-0.010) 4.300 (-0.007) 4.316 (0.009) 4.277 (-0.03) C3n.2n 4.484 4.480 (0.004) 4.493 (+0.009) 4.493 (0.009) 4.479 (-0.005) 4.500 (0.016) base 4.601 4.620 (-0.019) 4.632 (0.031) 4.631 (0.03) 4.623 (0.022) 4.631 (0.03) C3n.3n 4.785 4.800 (-0.015) 4.799 (0.014) 4.799 (0.051) 4.781 (-0.004) 4.760 (-0.025) base 4.897 4.890 (0.007) 4.879 (-0.018) 4.896 (-0.001) 4.878 (-0.019) 4.889 (-0.008) C3n.4n 4.987 4.980 (0.007) 4.998 (0.011) 4.997 (0.01) 4.977 (-0.01) 5.009 (0.022) base 5.182 5.230 (-0.048) 5.236 (0.054) 5.235 (0.053) 5.232 (0.05) 5.273 (0.091) C3An.1n 5.735 5.894 (0.159) 5.952 (0.217) 6.033 (0.298) 5.875 (0.14) base 5.955 6.137 (0.182) 6.214 (0.259) 6.252 (0.297) 6.122 (0.167)

PAGE 61

61Table 3-3. Nannofossil datums for ODP Site 1208 (Bown, 2005). Ages for the datums areinterpolated from the magnetic stratigraphy (this work) and the correlative polaritychron is given. The datums are compared to ages given by Berggren et al. (1995a, b).Tuned ages for the datums are compared to ATNTS2004 (Lourens et al., 2004) andODP Leg 138 ages for nannofossil datums only (Raffi and Flores, 1995; Shackletonet al., 1995a). Datum Depth(mbsf) 1208Mag. strat.Age (Ma) ChronSite 1208 Berggren etal.1995a b(Ma) chron ODPLeg138 1208Tunedage (Ma) ATNTSage(Ma) FO Emiliania huxleyi 14.24 0.258 C1n 0.26 0.26 0.29 LO P. lacunosa 30.30 0.551 C1n 0.46 0.46 0.44 FO G. omega 43.11 0.784 C1r.1r FO G. caribbeanica 87.90 1.837 C2n 1.841 LO D. brouweri 100.16 2.143 C2r.1n 1.95 Olduvai 1.96 2.146 2.06 LO D. pentaradiatus 116.40 2.510 C2r.2r 2.46-2.56 M/Gboundary 2.52 2.499 2.39 LO D. surculus 119.08 2.572 C2r.2r 2.55-2.59 M/Gboundary 2.63 2.556 2.52 LO D. tamalis 128.70 2.812 C2An.1n 2.78 top Gauss 2.78 2.802 2.80 LO LargeReticulofenestra 163.90 3.851 C2Ar 3.833 FO D. tamalis 166.66 3.65 C2Ar 3.95 LO Sphenolithus 166.66 3.65 C2Ar 3.6 base Gauss 3.66 3.95 LO Amaurolithus 168.88 4.56 C2Ar 4.03 FO D. asymmetricus 168.88 4.56 C2Ar 4.2 top Cochiti 4.13 4.03 FO C. cristatus 187.90 4.735 C3n.2r 4.750 LO D. quinqueramus 207.00 5.551 C3r 5.6 C3r 5.55 5.472 5.59 FO Amaurolithus 235.52 6.941 C3Bn FO D. quinqueramus 250.80 8.075 C4r.1r FO D. berggrenii 250.80 8.075 C4r.1r 8.6 C4r.2r 8.45 FO D. hamatus 265.94 9.586 C4Ar.2n 9.4 C4Ar.2r FO C. calyculus 270.10 9.792 C5n.1n 10.79 FO D. hamatus 274.20 10.125 C5n.2n 10.7 10.38 10.55 FO C. coalitus 279.70 10.699 C5n.2n 10.9 10.89 LO C. miopelagicus 285.06 11.009 C5r.1r 10.8 11.02 LO C. premacintyrei 295.41 13.19 C5An.1r 12.65 11.21 LO C. floridanus 295.41 13.19 C5An.1r 13.19 13.33

PAGE 62

62Table 3-4. Plio-Pleistocene foraminfer datums, with depths, correlative polarity chron, tuned ageand compared to Berggren et al. (1995a, b) and ATNTS 2004 (Lourens et al., 2004)from ODP Legs 138 and 111. Event Depthmbsf 1208mag stratage ChronSite 1208 Berggren etal 95ab age(Ma) 1208tunedage ATNTSAge LO Gr. crassula 0.4 0.007 C1n LO Gr. tosaensis 16.1 0.292 C1n 0.65 C1n 0.61 LO Gs. bulloideus 38.2 0.694 C1n LO B. praedigitata 43.7 0.797 C1r.1r LO Gt. woodi 53.2 1.005 C1r.1n 1.004 2.3 LO Gs. bollii 60.5 1.182 C1r.2r 1.197 LO Gs. obliquus 66.7 1.331 C1r.2r 1.34 1.3 LO N. acostaensis 76.2 1.559 C1r.2r 1.562 1.58 LO N. humerosa 79.6 1.640 C1r.2r 1.653 LO Gt. decoraperta 83.6 1.736 C1r.2r 1.742 2.75 LO Gr. puncticulata, FO Gs. tenellus,FO Gs. elongatus 89.7 1.878 C2n 2.41 1.882 2.41 FO Gr. hirsuta 92.3 1.938 C2n 1.930 FO Gr. toseansis, LO Gr. cibaoensis, FO Gr.truncatulinoides 95.2 2.014 C2r.1r 3.35C2An.2n 2.015 1.93 LO Pu. primalis, LO Gr. limbata FO Ga. parkerae 98.5 2.103 C2r.1r 2.116 FO Pu. obliquiloculata, FO B. digitata 101.9 2.171 C2r.2r 2.164 LO Gq. venezuelana 108.1 2.316 C2r.2r 2.282 LO Ss. paenedehiscens, LO Gt. apertura 111.4 2.393 C2r.2r 2.382 LO N. dupac, FO Ge. siphonifera 121.5 2.632 C2An.1n 2.621 LO Gr. pseudomiocenica 123.0 2.670 C2An.1n 2.654 LO Gr. juanai, FO Gr. bermudezi 125.8 2.740 C2An.1n 2.714 LO Gr. plesiotumida, LO Gs. extremus,LO Gs. triloba 132.5 2.907 C2An.1n 2.902 FO Gr. limbata 135.3 2.978 C2An.1n 2.966 LO Gq. conglomerata, LO Gr. sphericomiozea 137.2 3.250 C2An.1n 3.024 FO Pu. primalis 142.2 3.154 C2An.2n 3.147 LO Gr. inflata 146.2 3.276 C2An.2r 3.301 FO Gt. rubescens 147.4 3.318 C2An.2r 3.338 FO Gr. puncticulata, LO Gr. conoidea 150.0 3.391 C2An.3n 4.5 Nunivak 3.427 FO Sa. dehiscens 151.5 3.433 C2An.3n 5.2 E.Gilbert 3.457 LO Ge. pseudobesa 158.2 3.631 C2Ar 3.637 FO Ge. calida, FO Gr. crassula, LO D. altispira, LO Ss.seminulina 159.6 3.740 C2Ar 3.682 LO Ss. kochi 161.0 3.739 C2Ar 3.708 4.53 LO Gr. margaritae 162.4 3.793 C2Ar 3.58 G/Gboundary 3.761 3.85 FO Gs. bulloideus 165.0 3.894 C2Ar 3.871 FO Gq. conglomerata, FO Gr. crassiformis 172.0 4.165 C2Ar 4.187 FO Ga. uvula, FO Gs. extremus,LO Gr. conomiozea 178.7 4.360 C3n.1r 4.413 FO Gg. umbilicata, FO Ge. aequilateralis 182.9 4.499 C3n.2n 4.54 FO Gr. sphericomiozea 186.5 4.669 C3n.2r 5.6 C3r 4.696 FO Gr. conomiozea, LO Gt. nepenthes,FO Gr. pseudomiocenica 190.8 4.879 C3n.3n 4.2 Cochiti 4.873 4.37

PAGE 63

63Table 3-5. Miocene foraminifer datums, with depths, correlative polarity chron, tuned age andcompared to Berggren et al. (1995a, b) and ATNTS 2004 (Lourens et al., 2004) fromODP Legs 138 and 111. Datum Depth(mbsf) 1208 magstrat age(Ma) ChronSite 1208 Berggrenet al.(1995b) 1208tunedage ATNTS2004 FO Gr. tumida, LO Gs. kennetti 203.1 5.354 C3r 5.6 C3r 5.327 5.57 FO Gs. bollii 209.4 5.608 C3r 5.591 FO Gs. kennetti 211.3 5.685 C3r 5.675 FO Gd. hexagona 214.3 5.806 C3r 5.816 FO N. dutertrei, FO Gs. conglobatus 217.3 5.934 C3An.1n 5.961 6.2 FO Ss. paenedehiscens, LO Gr.merotumida 224.5 6.342 C3An.2n FO Ge. pseudobesa, FO Gr. margaritae,FO Ss. kochi, FO Gr. plesiotumida, FO N. humerosa, FO Gr. scitula 227.3 6.434 C3An.2n 6.0 C3An LO Gr. miotumida c.f. 233.6 7.0 C3Ar FO Gr. cibaoensis, FO N. acostaensis,FO Gr. miotumida c.f. 240.2 7.4 C3Bn 7.8 C4n.2n LO Gq. baroemoenensis 245.9 7.9 C4n.2n FO Gr. juanai 251.1 8.1 C4r.1r FO B. praedigitata, FO Gs. obliquus,FO N. pachyderma (dextral), (sinistral) 255.6 8.7 C4r.1r FO Gs. ruber, FO Gt. apertura, LO Gq.dehiscens 263.1 9.4 C4Ar.1n FO Gr. merotumida 270.5 9.9 C5n.1n LO Gr. praemenardii 272.6 10.0 C5n.2n FO N. dupac 276.8 10.2 C5n.2n LO Gt. druryi 282.3 11.0 C5r.1r LO Ss. disjuncta 284.3 11.1 C5r.1r 11.49 FO Gt. decoraperta 285.0 11.1 C5r.1r FO Gr. miozea 291.9 11.8 C5r.3r LO Gr. mayeri 294.1 12.1 C5An.1r LO N. continuosa 295.2 12.1 C5An.1r LO Cs. parvulus, LO Gr. panda 296.8 12.2 C5An.1r 11.8C5r.3r FO Gt. nepenthes, FO Gr. mayeri, FO Gr.menardii 298.9 12.3 C5An.2n 11.8C5r.3r 11.63 FO Gt. druryi 299.7 12.3

PAGE 64

64 Figure 3-1. Bathymetric map showing the location of Shatsky Rise in the Pacific Ocean and alarger map of Shatsky Rise showing the Sites drilled on ODP Leg 198 including Site1208 on the Central High of the Rise (after Bown, 2005).

PAGE 65

65 Figure 3-2. Inclination, declination and MAD values plotted against meters below sea floor. Grayline indicates the AF demagnetization data from the shipboard pass-throughmagnetometer at the 20 mT demagnetization step. Open squares indicate data fromdiscrete samples. The polarity interpretation is shown in black (normal polarity) andwhite (reverse polarity) and chrons are labeled according to Cande and Kent (1992,1995). Excursions are labeled according to Channell et al. (2002) and Singer et al.(1999). Ages for excursions are calculated from the astronomic age model for Site1208.

PAGE 66

66 Figure 3-3. Inclination, declination and MAD values plotted against meters below sea floor. Grayline indicates data from the shipboard pass-through magnetometer at the 20 mTdemagnetization step. Open squares indicate data from discrete samples. The polarityinterpretation is shown in black (normal polarity) and white (reverse polarity) andchrons are labeled according to Cande and Kent (1992, 1995). Excursions are labeledaccording to Channell et al. (2002) and Singer et al. (1999). Ages for excursions arecalculated from the astronomic age model for Site 1208.

PAGE 67

67 Figure 3-4. Inclination, declination and MAD values plotted against meters below sea floor. Grayline indicates data from the shipboard pass-through magnetometer. Open squaresindicate data from discrete samples. The polarity interpretation is shown in black(normal polarity) and white (reverse polarity) and chrons are labeled according toCande and Kent (1992, 1995). Gray bar indicates indeterminate polarity.

PAGE 68

68 Figure 3-5. Orthogonal projections showing AF demagnetization data from discrete samples.Open circles represent the vector end point projections on the vertical plane andclosed circles represent vector end point projections on the horizontal plane.

PAGE 69

69 Figure 3-6. a) Interval sedimentation rates (black line) and age versus depth (red line) calculatedfrom the magnetostratigraphic data. b) interval sedimentation rates calculated for thetuned age model for the 1-6 Ma interval.

PAGE 70

70 Figure 3-7. Reflectance (L*) data (black line) tuned to the astronomic solution for obliquity fromLaskar et al. (2004). Lower plots shows the output of a gaussian filter centered on theobliquity frequency (0.024), applied to the reflectance (L*) data.

PAGE 71

71 Figure 3-8. Plio-Pleistocene planktonic foraminifer and calcareous nannofossil datums, corerecovery, and magnetostratigraphy, plotted against meters below the sea floor. Agesin bold are astronomically calibrated ages from this study. Ticks indicate position ofsamples taken for foraminifer analysis. Depths in mbsf of datums are given inparentheses before the datum.

PAGE 72

72 Figure 3-9. Miocene planktonic foraminifer and calcareous nannofossil datums core recovery,and magnetostratigraphy, plotted against meters below the sea floor (after Venti,2006). Ages in bold are astronomically calibrated ages from this study. Ticks indicateposition of samples taken for foraminifer analysis. Depths in mbsf of datums aregiven in parentheses before the datum.

PAGE 73

73 Figure 3-10. Calcareous nannofossil biostratigraphy including the zonations of Martini (1971)and Okada and Bukry (1980) modified from Bukry (1973, 1975). Ages in bold areastronomically calibrated ages from this study.

PAGE 74

74 Figure 3-11. A proposed biostratigraphy for the mid-latitude North Pacific uses 16 plankticforaminifer datums to divide the late Neogene into 15 biozones. The new stratigraphyis integrated into the Geomagnetic Polarity Timescale and compared to pre-existingplanktic foraminifer zonal schemes for temperate and tropical region, as well as totropical calcareous nannofossil zonations (after Venti, 2006). Abbreviations forzonations are as follows: B69: Blow (1969) modified by Kennett and Srinivasan(1981a, 1981b) BKSA95: Berggren et al. (1995b) J85: Jenkins (1985) SK81:Srinivasan and Kennett (1981a) M71: Martini (1971) B73,75: Bukry (1973, 1975),OB80 Okada and Bukry, (1980): Ages in bold are astronomically calibrated agesfrom this study.

PAGE 75

75CHAPTER 4PALEOINTENSITY-ASSISTED CHRONOSTRATIGRAPHY OF DETRITAL LAYERS ONTHE EIRIK DRIFT (NORTH ATLANTIC) SINCE MARINE ISOTOPE STAGE 11IntroductionThe Eirik Drift drapes the top of the underlying Eirik Ridge located off the southern tip ofGreenland (McCave and Tucholke, 1986). Magnetic anomalies have not been identified directlybeneath the Eirik Ridge, although the adjacent oceanic crust in both the Irminger Basin andLabrador Sea is associated with marine magnetic anomaly 24 of Paleocene-Eocene boundary age(Srivastava and Tapscott, 1986). The Eirik drift is 800 km long and has been constructed by theinteraction of the southwestward flowing Western Boundary Undercurrent (WBUC) andbasement topography (Chough and Hesse, 1985). The WBUC carries water masses originatingfrom the Norwegian and Greenland Seas that enter the North Atlantic over the Iceland-ScotlandRidge and Denmark Strait (McCave and Tucholke, 1986; Lucotte and Hillaire-Marcel, 1994).The WBUC moves over, and constructs the Eirik Drift and then follows bathymetric contoursaround the Labrador Basin (McCave and Tucholke, 1986).Drilling on the Eirik Drift includes Site 646 (ODP Leg 105), and piston and gravity corescollected during cruises by the CSS Hudson in 1990, the Marion Dufresne in 1999 and the R/VKnorr in 2002. Seismic records used to extrapolate the sequence recovered at Site 646 indicatethat the drift has been constructed since the middle to early Pliocene (Arthur et al., 1989).Although sedimentation on the drift sequence was more or less continuous during the LatePliocene and Pleistocene, sedimentation rates vary considerably with glacial/interglacialconditions and with location on the drift.Piston cores HU90-013-012 (water depth: 2830 m) and HU90-013-013 (water depth: 3380m) (Figure 4-1, Table 4-1), collected in 1990 during a cruise of the CSS Hudson, record the lastglacial cycle at differing water depths on the Eirik Drift (Hillaire-Marcel et al., 1994). Core

PAGE 76

76HU90-013-013 shows high sedimentation rates in the Holocene while Core HU90-013-012 hasvery low Holocene sedimentation rates due to winnowing by the WBUC (Stoner et al., 1995a,1996). Increases in magnetic concentration and grain size during the early Holocene and at theMIS 6/5e transition in HU90-013-013, were attributed to detrital influx associated with retreat ofthe Greenland Ice sheet (Stoner et al., 1995b). In core HU90-013-013, four discrete detritallayers were identified within MIS 2 and 3 based on their magnetic properties (coarse magneticgrain size) and relatively high percent carbonate values. Stoner et al. (1996) correlated three ofthese detrital layers with Heinrich events 1, 2 and 4. Stoner et al., (1998) revised the chronologyfor core HU90-013-013 by correlation to SPECMAP (Martinson et al., 1987) and refined theages and correlation of the detrital layers to North Atlantic detrital layers.We present data from three jumbo piston cores (JPC15, JPC18, JPC19) collected on theEirik Drift in the summer of 2002 during Cruise KN166-14 of the RV Knorr, and from CoreMD99-2227 collected during the 1999 Images campaign (Figure 4-1). JPC15 was taken on theupper slope of the ridge at a water depth of 2230 m. Core JPC19 was collected from the crest ofthe ridge at a water depth of 3184 m, and Core JPC18 from the southern flank of the ridge at awater depth of 3435 m. Core MD99-2227 was collected from the western toe of the drift at 3460m water depth. The recovered sediments are mostly dark gray bioturbated silty clays, with clayeysilt and sandy mud, and occasional gray nannofossil/foraminifer rich clayey silt layers (seeTuron, Hillaire-Marcel et al., 1999, for a lithologic description of MD99-2227).MethodsU-channel samples (2x2 cm square cross-section and 150 cm in length) were collectedfrom the center of the split face of piston core sections. These samples were measured on a 2G-Enterprises pass-through cryogenic magnetometer at the University of Florida. Natural remanentmagnetization (NRM) was demagnetized step-wise using alternating fields (AF) in 5 mT

PAGE 77

77increments for 0-60 mT peak fields, and in 10 mT increments for 60 mT-100 mT peak fields.Volume susceptibility was then measured using a susceptibility track specifically designed for u-channels (Thomas et al., 2003) that has a measurement resolution of a few centimeters.Anhysteretic remanent magnetization (ARM) was applied using an AF field of 100 mT and abias DC field of 50 !T. Isothermal remanent magnetization (IRM) was imparted using a 0.5 TDC field. Both artificial remanences were demagnetized with the same AF steps used todemagnetize NRM. Principal components were calculated from the NRM data using the methodof Kirschvink (1980) applied to the 20-80 mT interval. Relative paleointensity proxies weregenerated by normalizing the NRM data by both ARM or IRM, demagnetized at a common peakfield. A mean of nine normalized remanence values, in the 20-60 mT peak field range, was usedto generate the relative paleointensity proxies. ARM and susceptibility data were also used toascertain magnetic grain size changes that help define detrital layers. The parameter karm(anhysteretic susceptibility), obtained by normalizing ARM intensity by the strength of the dcfield used to acquire the ARM, was divided by volume susceptibility, to determine karm/k, aproxy for magnetite grain size.On completion of the magnetic measurements on the u-channel samples, X-radiographswere taken across detrital layers, identified by u-channel magnetic measurements and carbonateanalyses, to provide a picture of the internal structure of these layers and identify the presence orabsence of traction structures. Discrete toothpick-sized samples, collected at 1-cm intervalsacross detrital layers, were used for smear slide observation (Table 4-2) and for measurement ofmagnetic hysteresis parameters using a Princeton Measurements Corp. vibrating samplemagnetometer (VSM). Magnetic hysteresis parameters provide a means of estimating magnetitegrain size, and therefore of recognizing grading in detrital layers.

PAGE 78

78Cores were sub-sampled for oxygen isotope analysis at 5-cm spacing. Samples from CoreMD99-2227 were analyzed at GEOTOP (Montreal) while samples from the KN166-14 coreswere analyzed in the stable isotope laboratory at Rutgers University. For all the cores,foraminifer shells of the planktonic species Neogloboquadrina pachyderma (left coiling) werepicked in the 150-250 m fraction for the isotopic analyses. Planktonic foraminifer species wereused for the isotopic analyses due to the small amount of benthos present in the cores. For CoreMD99-2227, samples were collected at 5 cm intervals for carbonate analyses using an elementalanalyzer.Age models for the piston cores were constructed by matching relative geomagneticpaleointensity records and planktic "18O records to target curves, with the location of magneticexcursions (Laschamp and Iceland Basin) providing additional age constraints. The combinationof paleointensity records and oxygen isotope data provide enhanced temporal resolutioncompared to using either dataset independently.NRM and Normalized Remanence RecordThe natural remanent magnetization (NRM) data for all four cores are shown ascomponent inclination, corrected component declination, and maximum angular deviation(MAD) values (Figure 4-2). Cores were not oriented during collection, and therefore declinationdata were corrected by aligning the mean declination of each core to North. Twisting withincores during the coring process is indicated by anomalous declination changes in Core JPC18(114.5-189 cm) (Figure 4-2). Core MD99-2227 is affected by stretching in the upper 7 metersthat has significantly affected the magnetization directions (Figure 4-2).

PAGE 79

79Polarity ExcursionsBrief polarity excursions are a characteristic of the geomagnetic field, at least during thelast ~2 Myr, and excursions of known age provide useful stratigraphic markers. Componentmagnetizations from u-channels indicate directional excursions at 9.3 meters below seafloor(mbsf) in Core JPC15, at 13.4 mbsf in Core JPC18, and at 18.7 mbsf in Core JPC19 (Figures 4-2and 4-3). For Core JPC15, the observed excursion is correlated to the Laschamp excursion (~41ka). For Cores JPC18 and JPC19, the observed excursion is correlated to the Iceland Basinexcursion (~185 ka). Orthogonal projections of alternating field demagnetization data fromintervals recording the Iceland Basin excursion in Cores JPC18 and JPC19 (Figure 4-3) indicatethat the excursions are unambiguously recorded by u-channel samples and by discrete samplescollected alongside the u-channel trough.Relative PaleointensityIt is generally accepted that the generation of useful paleointensity proxies requires that thesediments contain magnetite as the only NRM carrier. Also the sediment should have a narrowrange of magnetite concentration, as indicated by magnetic concentration parameters varying byless than an order of magnitude, and have restricted magnetite grain-size in the few micron grain-size range, corresponding to pseudo-single domain grains (Tauxe, 1993). There is no evidencefrom demagnetization characteristics of NRM, or from hysteresis parameters, for high-coercivitymagnetic minerals such as hematite or pyrrhotite. Using plots of anhysteretic susceptibilityagainst susceptibility, and the calibration of King et al. (1983), we estimate that these sedimentsgenerally have magnetite grain sizes in the 1-10 m range (Figure 4-4). Records of ARM, IRMand susceptibility (Figure 4-5) show that the concentration parameters generally vary within anorder of magnitude, the limit deemed suitable for determination of relative paleointensity proxies

PAGE 80

80(Tauxe, 1993). The exception is within the coarser-grained intervals in the early part ofinterglacials, where the concentration parameters vary by more than an order of magnitude.NRM measured on u-channel samples was normalized using both ARM and IRM,demagnetized at the same peak fields as the NRM. To generate the paleointensity proxies, amean of nine demagnetization steps in the 20-60 mT interval were used to calculate meanNRM/ARM and mean NRM/IRM. Although the two proxies are generally consistent with eachother, mean NRM/ARM has the lower standard deviations and was therefore chosen as thepreferred paleointensity proxy.ChronologyTo construct age models for the four cores in this study, we correlate the planktonicoxygen isotope records to the benthic oxygen isotope stack (Lisiecki and Raymo, 2004). We thenadjust this correlation to optimize the fit of the relative paleointensity records to thepaleointensity record from ODP Site 983 (Channell et al., 1997; Channell, 1999). FollowingStoner et al. (2003), the paleointensity and oxygen isotope data from ODP Site 1089 were usedto improve the age model for ODP Site 983 particularly in the MIS 3-4 interval. For the EirikDrift cores, a combination of oxygen isotope data and relative paleointensity data can produce ahigher-resolution age model than would be possible using either data set independently.The magnetic excursion recorded at 18.7 mbsf in JPC19 (Figures 4-2 and 4-3) isinterpreted as the Iceland Basin excursion (Channell et al., 1997; Channell, 1999). It lies in aprominent paleointensity low at 185 ka in JPC19 (Figure 4-6), consistent with the expected ageof this excursion. According to the age model, Core JPC19, from the crest of the drift at a waterdepth of 3184 m, has an age at its base of 300 kyrs with a mean sedimentation rate of 10.5cm/kyr.

PAGE 81

81In Core JPC18, from southern flank of the Eirik ridge at a water depth of 3435 m,sediments coeval with interglacial periods are apparently missing, as shown by the lack ofHolocene oxygen isotope values (Figure 4-7). MIS 5e is also absent in the record, becauseoxygen isotope values in this interval are too high for full interglacial values. The polarityexcursion observed at 13.45 mbsf (Figures 4-2 and 4-3) is identified as the Iceland Basinexcursion and it occupies a distinct paleointensity low at 185 ka (Figure 4-7), an age consistentwith the observation of this excursion elsewhere. The overall mean sedimentation rate in CoreJPC18 is 9 cm/kyr.Core JPC15 was taken on the upper slope of Eirik ridge at a water depth of 2230 m. Thepolarity excursion observed at 9.3 mbsf in Core JPC15 (Figures 4-2 and 4-3) occurs within aprominent paleointensity low at ~ 40 ka (Figure 4-8) and is therefore interpreted as theLaschamp excursion. The base of JPC15 has an age of 160 ka and the mean sedimentation rate is15 cm/kyr (Figure 4-8).Core MD99-2227 shows significant stretching in the upper part of the core, however, thecorrelation to the calibrated ODP Site 983 paleointensity record is possible in the lower part(Figure 4-9). The paleointensity correlation is consistent with the correlation of the plankticoxygen isotope record to the benthic oxygen isotope stack of Lisiecki and Raymo (2005). Thesecorrelations give a basal age for Core MD99-2227 of 430 ka, and mean sedimentation rates of 10cm/kyr (Figure 4-9).Detrital Layer StratigraphyThe ratio of anhysteretic susceptibility to susceptibility (karm/k) has been shown to be auseful magnetite grain size proxy (e.g. King et al., 1983; Tauxe, 1993). Although the plots of karmversus k of each core (Figure 4-4) indicate magnetic grain sizes within a restricted (few micron)range, the karm/k data plotted versus age (Figure 4-10) indicate distinct broad intervals of low

PAGE 82

82values of karm/k that coincide with the early Holocene (when recorded), with MIS 5e, and withthe early parts of MIS 7, 9 and 11 (shaded in Figure 4-10). Low values of karm/k indicaterelatively coarse magnetite grain sizes in these intervals. Although Core JPC18 is missing part ofthe Holocene, and almost the entire MIS 5e, the intervals of low values of karm/k appear to bepartially recorded.Volume magnetic susceptibility data measured on u-channel samples from Cores JPC19and MD99-2227 show an increase in magnetic concentration in the early Holocene, MIS 5e, andin the early parts of MIS 7, 9 and 11 (Figure 4-10). These intervals of high magneticconcentration coincide with the intervals of low values of karm/k (Figure 4-10) that indicaterelatively coarse magnetite grain sizes.In Core JPC15, high sedimentation rates between 20-60 ka (500-1200 cm) allow theidentification of millennial-scale cycles in volume magnetic susceptibility (Figure 4-5). Theseappear to mimic the D/O cycles the Greenland Ice Core (GISP) oxygen isotope record, and arereminiscent of susceptibility cycles identified by Kissel et al. (1999) in cores along the path ofNorth Atlantic Deep Water (NADW), and attributed to changes in the strength of bottomcurrents. The depth of the WBUC, that varies in response to the relative outflows of watermasses from the Greenland and Norwegian Seas, could also be account for the variations.In addition to these broad decimeter-scale intervals defined by karm/k and k values, a totalof seventeen cm-scale layers with magnetic properties and percent carbonate values significantlydifferent from the surrounding sediments have been identified in MD99-2227 (Figure 4-11).These layers have been labeled according to marine isotope stage and their detrital carbonate(DC) content. For example, 6LDC indicates a low detrital carbonate (LDC) layer within MIS 6(Table 4-2).

PAGE 83

83Eight of the seventeen cm-scale layers are designated detrital carbonate layers (DC) on thebasis of their high detrital carbonate contents. Four of these layers (3DC, 7DCa, 8DC, 11DC) arerecognized by coarser grained magnetic material (compared to the background sediment), asindicated by low karm/k values (Figure 4-11). One of these DC layers (7DCa) shows a peak inmagnetic susceptibility while the other seven DC layers do not. Two DC layers (5DC, 9DC)show finer-grained magnetic material (compared to background sediment), and two DC layers(7DCb, 2DC) are not differentiated by magnetic grain size from the background sediment but allDC layers coincide with highs in percent carbonate and six show peaks in GRA bulk density(Figure 4-11). All DC layers are light in color, do not show a sharp base, and appear to showsome bioturbation. The X-radiographs of these layers confirm a high concentration of IRD, butno laminae or evidence for traction (Figure 4-12). Smear slides indicate a high percentage ofcoarse detrital carbonate material in these layers (Table 4-2).Nine of the seventeen cm-scale detrital layers are designated low detrital carbonate (LDC)layers (Figure 4-11). These do not feature an increase in percent carbonate, but show a peak inmagnetic susceptibility, a low in karm/k, and an increase in GRA bulk density. These LDC layersoccur within MIS 1, 2, 5, 6, 7, 9 and 11 and show sharp bases, bioturbated tops and are 4-18 cmthick (Figure 4-11, Table 4-2). The X-radiographs indicate a sharp base and laminae within thelayers (Figure 4-12), some of the laminae are inclined and indicative of traction, implying rapiddeposition from turbidity currents or contourites.Toothpick-sized samples collected at 1cm intervals through detrital layers were used todetermine magnetic hysteresis parameters that can be used as a means of assessing the grain sizeof magnetite (Day et al., 1977). All but one of the detrital layers exhibit hysteresis parametersthat fall within the pseudo-single domain (PSD) grain size range (Figure 4-13). The detrital

PAGE 84

84carbonate layer identified in MIS2 (2DC) shows coarse multi-domain magnetite that isanomalous compared to all other detrital layers (Figure 4-13). For five of the nine LDC layers,we see evidence for progressive change in hysteresis parameters through the detrital layerindicative of grading, fining upward from the base of the layer. Bioturbation of the detrital layerinto the overlying sediment could also cause the layer to appear graded. However, the presenceof distinct laminae within the LDC layers shows that no bioturbation of the layer has occurred.None of the DC layers show this grading in hysteresis parameters. The presence of grading inthe LDC layers indicates a turbiditic rather than a contourite origin for these layers. Smear slides indicate that LDC layers contain little clay and significant amounts of silt-sized opaque grains, green hornblende and quartz. Trace amounts of detrital carbonate arepresent in LDC layers and throughout the rest of the core, whereas the percentage of detritalcarbonate in the DC layers exceeds 10% (Table 4-2).DiscussionSedimentation rates on the Eirik Drift have been shown to be greatly affected by changesin the strength and bathymetry of the Western Boundary Undercurrent (WBUC) that is thoughtto be switched off during glacials and active during interglacials (Hillaire-Marcel et al., 1994;Hillaire-Marcel and Bilodeau, 2000). The core of this current is thought to occupy water depthsbetween 2500 and 3000 meters (Hillaire-Marcel et al., 1994), resulting in winnowing and almostcomplete removal of Holocene and MIS 5e sediment from these depths. Cores from outside theinfluence of the flow would be expected to have interglacial sedimentation rates comparable to,or higher than, glacial sedimentation rates.When combined with previous studies carried out on the drift, the new results indicate thatboth water depth and position on the drift influence interval sedimentation rates. Although thesite of Core JPC18 is located ~450 meters below the supposed core of the WBUC, sediment of

PAGE 85

85Holocene and MIS 5e age is missing at this site. This implies that the WBUC is active at deeperwater depths than previously supposed on the southern side of the Eirik ridge (Figure 4-1). Thismay be consistent with a deep branch of the WBUC, with a gyre in the outer Labrador Sea thatfeeds the Gloria Drift (Figure 4-1).Cores HU90-013-013 (water depth 3471 m), JPC19 (water depth 3184 m) and MD99-2227have relatively high Holocene sedimentation rates of 35 cm/kyr, ~13 cm/kyr, and 10 cm/kyrrespectively. Sedimentation rates in cores MD99-2227 and JPC19 appear to be low at the onsetof deglaciation and then increase. This may be due to increased winnowing by the WBUC at theonset of the deglaciation, offset by increased detrital input as the deglaciation proceeds.Core HU90-013-012 at 2830 meters water depth lies within the influence of the WBUCand has very low sedimentation rates in the Holocene (Stoner et al., 1995a, 1996). Higher up theslope, Core JPC15 at a water depth of 2230 meters has low sedimentation rates in the Holoceneand MIS 5e, although the site supposedly lies outside the main influence of the WBUC. Hillaire-Marcel et al. (1994) noted that, in core HU90-013-06 at even shallower water depths (1105 m)on the Eirik ridge, active bottom currents also resulted in very low Holocene sedimentation rates.Hillaire-Marcel et al. (1994) interpreted DC and LDC layers deposited during the lastglacial cycle at Orphan Knoll, on the western side of the Northwest Atlantic Mid-Ocean Channel(NAMOC), as being related to ice advances of the Laurentide Ice Sheet that triggered turbiditicflows down the NAMOC (Figure 4-1). Sediment suspended by these flows is thought to havedeposited cm-scale sandy mud beds rich in detrital carbonate (DC layers) at Orphan Knoll. Notall the detrital layers observed at Orphan Knoll are recognized on Eirik Drift, although two LDClayers and one DC layer in Core HU90-0130-013 (Figure 4-1) were considered coeval withOrphan Knoll detrital layers (Stoner et al., 1996).

PAGE 86

86The cm-scale detrital layers identified in core MD99-2227 extend the record of detritallayers beyond the last glacial cycle. Detrital layers on Eirik Drift occur during both glacial andinterglacial conditions. However, the layers occurring in the interglacials are close to theTerminations in the Holocene, MIS 5, 7 and 11. It is only in MIS 9 that the DC layer appears tooccur in the later part of the interglacial implying that the Laurentide Ice Sheet was presentthroughout MIS 9.Detrital layer 1LDC with an age of 13 ka in MD99-2227 (Table 4-3) is tentativelycorrelated to DC0 of Stoner et al. (1998). Layer 2LDC has an age of 18 ka and is correlated toDC1 (16 ka) from Orphan Knoll (Stoner et al., 1998) and with H1 of Bond et al., (1999) from thecentral Atlantic. The DC layer 2DC correlates with DC2 of Stoner et al. (1998) and with H2(Bond et al., 1999). The detrital layer labeled 3DC (39 ka) is correlated to DC4 from OrphanKnoll and to H4 (38 ka). As discussed above, the characteristics of LDC layers impliesdeposition by turbidity currents (derived from the Greenland Slope). If so, this turbiditic activityis sometimes coeval with Heinrich layers of the central Atlantic and with detrital events atOrphan Knoll.The ages of layers designated 2LDC, 2DC and 3DC in this study are consistent with agesfor Heinrich events H1, H2, and H4 (Table 4-3). No identifiable events that coeval with Heinrichevents H3, H5 or H6 are found. Hiscott et al. (2001) identified Heinrich-like detrital layers incore MD95-2025 from near Orphan Knoll back to MIS 9. Two detrital carbonate layers withinearly MIS 5 at Orphan Knoll (H8 and H9 of Hiscott et al., 2001) appear to be coeval with DCevents identified on Eirik Drift, implying that instabilities of the Laurentide Ice Sheet arerecorded at both sites. Detrital carbonate layers within MIS 7 and MIS 9 at Orphan Knoll (H10and H13 of Hiscott et al., 2001) are coeval with a LDC layers (7LDC and 10 LDC) identified on

PAGE 87

87Eirik Drift (Table 4-3), implying that the LIS instabilities that triggered the detrital carbonatelayers at Orphan Knoll were coeval with instabilities on the Greenland slope that triggered theLDC layers on Eirik Drift. Such conclusions are highly dependent on the resolution ofstratigraphic correlation. While stratigraphic correlation of detrital layers from the Orphan Knollto the central Atlantic for the last glacial cycle is rather well constrained (Bond et al., 1999;Stoner et al., 1996, 2000), the correlations beyond the last glacial cycle are considerably morespeculative (e.g. Hiscott et al., 2001; van Kreveld et al., 1996) due to lack of stratigraphicresolution that inhibits unequivocal correlation of detrital layers.ConclusionsPiston cores collected from Eirik Drift have produced records of relative paleointensity andof the Laschamp and Iceland Basin polarity excursions that augment oxygen isotope data forgenerating age models. Magnetic data from cores JPC19 and MD99-2227 show broad intervalsof increased magnetic grain size and concentration during MIS 5e and at the MIS 2/1 transition,consistent with observations from Core HU90-013-013 (Stoner et al., 1995b). Core MD99-2227also shows a similar increase in magnetic grain size and concentration at the onset of interglacialMIS 7, 9 and 11, implying that retreat of the Greenland Ice Sheet produced a characteristicdetrital signal at the onset of all interglacial stages over the last 400 kyr.Seventeen cm-scale detrital carbonate and low detrital carbonate layers are identified inMD99-2227 (Figure 4-11, Table 4-2). They occur in both glacial and interglacial stages. Thedetrital layers can be subdivided into two classes. Detrital carbonate (DC) layers are composedof carbonate-rich IRD. They usually, but not always, carry a magnetic signal indicating highmagnetic concentration and increased magnetic grain size relative to background sediment. Lowdetrital carbonate (LDC) layers have <10% detrital carbonate, usually show evidence (frommagnetic hysteresis ratios) for fining-upward grading, and X-radiograph evidence for traction.

PAGE 88

88These layers are also usually marked by high magnetic concentration and increased magneticgrain size relative to background sediment.Based on the differences between DC and LDC layers, we interpret the former as HudsonStrait derived detrital layers, and the latter as layers dominated by material from turbiditesderived from the Greenland slope. 1LDC, 2LDC, 2DC and 3DC are correlative with detritallayers observed at Orphan Knoll (Stoner et al., 1996) (Table 4-3). Three of them (1LDC, 2DCand 3DC) are coeval with central Atlantic Heinrich layers H1, H2 and H4 (Bond et al., 1999).Beyond the last glacial cycle, the correlation of detrital layers from Eirik Drift (this paper) toOrphan Knoll (Hiscott et al., 2001) and to the central Atlantic (van Kreveld et al., 1996) islimited by the imprecision of stratigraphic correlation (Table 4-3). Nonetheless, as illustratedhere, the use of paleointensity-assisted chronostratigraphy, the combination of relativepaleointensity with standard oxygen isotope stratigraphy, improves stratigraphic correlationsacross the northern North Atlantic Ocean (and beyond), and thereby facilitates the interpretationof detrital layers in terms of their correlation, aerial extent and provenance.

PAGE 89

89Table 4-1. Core, latitude, longitude, water depth and base age of the core. Core Latitude Longitude Waterdepth Base age(kyr) JPC15 -45.57 58.20 2230 150 JPC18 -47.13 57.19 3435 300 JPC19 -47.60 57.58 3184 250 MD99-2227 -48.22 58.12 3460 430

PAGE 90

90Table 4-2. DC and LDC layer properties in Core MD99-2227. event Thick-ness(cm) depth(cm) Age(ka) MIS Name kpeak karm/k %carb GRAPEdensity %detr.carb. X-Ray Sharpbase Grading 1 5 440.22 13.04 1/2 1LDC yes coarse low peak 10 traction yes yes 2 14 616.85 18.2 2 2LDC yes coarse low peak trace traction yes yes 3 15 663 21.4 2 2DC no high peak 15 no no 4 21 858.7 39.1 3 3DC no coarse high peak 40 no no 5 6 1872.3 111.68 5 5LDC yes coarse low peak trace yes no 6 16 2019.4 129.34 5 5DC no fine high peak 70 no no 7 14 2192.9 152.17 6 6LDC yes coarse low peak trace traction yes yes 8 16 2505.4 191.58 7 7DCa yes coarse high 20 no no 9 12 2700 214.98 7 7DCb no high peak 25 no no 10 6 2872.3 233.42 7 7LDC yes coarse low peak trace traction yes no 11 11 3083.7 266.57 8 8DC no coarse high peak 20 no no 12 17 3229.2 289.89 9 9DC no fine high peak 30 no no 13 5 3536.4 335.6 9/10 9LDC yes coarse low peak trace traction yes no 14 18 4008.2 391.03 11 11LDCa yes coarse low peak trace traction yes no 15 4 4084.2 403.48 11 11LDCb yes coarse low peak 5 traction yes yes 16 7 4133.7 409.57 11 11LDCc yes coarse low peak 10 traction yes yes 17 7 4240 421.43 11 11DC no coarse high peak 50 IRD rich no no

PAGE 91

91Table 4-3. Detrital Layers from other studies considered to be correlative to detrital layersidentified on Eirik drift. Event Name depth(cm) MD99-2227Age (ka) Stoner et al.(1998)(age ka) H-layers Bondet al. (1999)(age ka) Hiscott et al(2001)(age ka) Van Kreveldet al (1996) (age ka) 1 1LDC 440.22 13.04 DC0 (12) H1(11-12) h1 (15) 2 2LDC 616.85 18.2 LDC1 (18) H1 (16.8) 3 2DC 663 21.4 LDC3 (21) H2 (24) H2 (18-22) h2 (21) 4 3DC 858.7 39.1 DC4 (36) H4 (38) H4 (39-42) h4 (40-43) 5 5LDC 1872.3 111.68 H8(92-108) 6 5DC 2019.4 129.34 H9(121-126) h7 (128-131) 7 6LDC 2192.9 152.17 8 7DCa 2505.4 191.58 h12 (189) 9 7DCb 2700 214.98 10 7LDC 2872.3 233.42 H10(231-240) 11 8DC 3083.7 266.57 12 9DC 3229.2 289.89 13 9LDC 3536.4 335.6 H13(335-340) 14 11LDCa 4008.2 391.03 15 11LDCb 4084.2 403.48 16 11LDCc 4133.7 409.57 17 11DC 4240 421.43

PAGE 92

92 Figure 4-1. Location map showing the Labrador Sea from Hillaire-Marcel and Bilodeau (2000)and the location of piston cores JPC15, JPC18, JPC19, and MD99-2227. Blackarrows indicate the path of the Western Boundary Undercurrent. NAMOC: NorthwestAtlantic Mid-Ocean Channel.

PAGE 93

93 Figure 4-2. Component inclination, corrected component declination and maximum angulardeviation (MAD) values for cores JPC15, JPC19, JPC18 and MD99-2227.

PAGE 94

94 Figure 4-3. a). Component inclination, declination and maximum angular deviation (MAD)values recording Laschamp and Iceland Basin polarity excursions from piston coresJPC15, JPC18 and JPC19. Key: U-channel data (closed circles), deconvolved u-channel data (open squares-dashed line) using the method of Guyodo et al. (2003), 8-cm3 discrete sample cubes (open squares) and 1-cm3 cubes (diamonds).3b).Orthogonal projections from the Iceland Basin excursion from cores JPC19 andJPC18, from u-channel data, deconvolved u-channel data, and discrete samples. Opencircles represent the vector end point projection on the vertical plane. Closed circlesrepresent the vector end point projection on the horizontal plane.

PAGE 95

95 Figure 4-4. Anhysteretic susceptibility (karm) plotted against volume susceptibility (k) for JPC18,JPC19, JPC15 and MD99-2227. Diamonds indicate background sediment, redsquares indicate coarse decimeter-scale interglacial intervals, and blue circles indicatecm-scale detrital layers. Black lines indicate magnetic grain-size boundaries placedusing the calibration of King et al. (1983).

PAGE 96

96 Figure 4-5. NRM, ARM, IRM and volume susceptibility for MD99-2227, JPC15, JPC18 andJPC19. Orange-IRM, green-ARM, red-NRM, blue-volume susceptibility.

PAGE 97

97 Figure 4-6. JPC19: Relative paleointensity record correlated to that from ODP Site 983(Channell et al., 1997; Channell, 1999). Lower plot: planktic "18O data from JPC19correlated to the benthic "18O stack of Lisiecki and Raymo (2005). Intervalsedimentation rates are shown in orange.

PAGE 98

98 Figure 4-7. JPC18: Relative paleointensity data correlated to ODP Site 983 (Channell et al.,1997; Channell, 1999). Lower plot shows planktic "18O data correlated to the benthic"18O stack of Lisiecki and Raymo (2005). Interval sedimentation rates are shown inorange.

PAGE 99

99 Figure 4-8. JPC15: Relative paleointensity data correlated to ODP Site 983 (Channell et al.,1997; Channell, 1999). Lower plot shows planktic "18O data correlated to the benthic"18O stack of Lisiecki and Raymo, (2005). Interval sedimentation rates are shown inorange.

PAGE 100

100 Figure 4-9. MD99-2227: Relative paleointensity data correlated to ODP Site 983 (Channell et al.,1997; Channell, 1999). Lower plot shows planktic "18O data correlated to the benthic"18O stack of Lisiecki and Raymo (2005). Black bar indicates stretched interval dueto coring. Interval sedimentation rates are shown in orange.

PAGE 101

101 Figure 4-10. karm/k and magnetic susceptibility versus age for cores, JPC19, JPC18 and MD99-2227 compared to Core HU90-013-013 (Stoner et al., 1995a). The benthic oxygenisotope stack of Lisiecki and Raymo (2005) is shown at the bottom of the figure.Shaded areas indicate magnetic coarse grain-size intervals (from karm/k) and magneticconcentration intervals (from susceptibility) in early and peak interglacial intervals.

PAGE 102

102 Figure 4-11. Core MD99-2227: karm/k, magnetic susceptibility, bulk (GRAPE) density, percentcarbonate, and planktic oxygen isotope data. Blue shading indicates detrital carbonate(DC) layers and green shading indicates low detrital carbonate (LDC) layers.

PAGE 103

103 Figure 4-12. Photographs and X-radiographs of three detrital layers identified in MD99-2227.Upper: MD99-2227 section 15 (6LDC), middle: MD99-2227 section 28 (11LDCc),and lower: MD99-2227 section 29 (11DC). In X-radiographs, light color indicateshigher density, dark color indicates lower density.

PAGE 104

104 Figure 4-13. Hysteresis ratios Mr/Ms plotted versus Hcr/Hc (after Day et al., 1977): individualdetrital layers from MD99-2227 are shown by colored symbols, with direction ofupward fining indicated by arrow. Black circles indicate background sediment.

PAGE 105

105CHAPTER 5RELATIVE PALEOINTENSITY STACK FOR THE LAST 85 KYR ON A REVISED GISPCHRONOLOGY, AND ENVIRONMENTAL MAGNETISM OF THE GARDAR DRIFTIntroductionFour cores were collected in the summer of 2002 along a NE-SW transect along the GardarDrift, at water depths from 1880 m to 3082 m. Chronologies for the cores were developed usingrelative paleointensity proxies and a benthic oxygen isotope record from the southernmost pistoncore (JPC13). The magnetic grain size proxy karm/k mimics benthic "18O, particularly in thenorthern and southern sites, indicating a link between magnetic grain size and bottom wateractivity. The benthic "18O record from Core JPC13 can be correlated to a similar record fromcore MD95-2042 from the Portuguese Margin, which has been correlated to both the GreenlandIce Core (GISP) and Vostok ice core records. As a result of this correlation, the relativepaleointensity records can be placed on the Shackleton-revised GISP chronology. A stack of 11relative paleointensity records from the North Atlantic region for the 0-85 ka interval has beendeveloped and placed on the revised GISP chronology. This stack (EHC06) shows somesignificant differences particularly in the 0-30 ka interval when compared to the GlobalPaleointensity Stack (GLOPIS). At 60 ka, when compared to both GLOPIS and the NorthAtlantic Paleointensity Stack (NAPIS), there is a difference in the age models of ~2400 years.The EHC06 stack is in better agreement with the independently dated South AtlanticPaleointensity Stack (SAPIS).Sedimentary relative paleointensity (RPI) records can provide important constraints onmechanisms in the geodynamo, can shed light on proposed geomagnetic-climate linkages, andmay provide a means of high-resolution global stratigraphic correlation. High-resolutionchronological control is critical for the study of millennial-scale climate change and RPI recordsprovide a potential means of global high-resolution correlation. When oxygen isotope

PAGE 106

106stratigraphy is combined with RPI records, the combination provides more robust, higherresolution age control than either data set alone. Stacks of RPI records allow the recognition ofregionally characteristic RPI features although the stacking process undoubtedly filters out thehigher frequency components. The North Atlantic Paleointensity Stack (NAPIS) wasconstructed from six relative paleointensity records from the North Atlantic (Laj et al., 2000).Chronological control for the NAPIS stack was based on correlation of the oxygen isotopestratigraphy from Core PS2644-5 to the GISP2 "18O record, allowing the stack to be placed on aGISP2 timescale. The NAPIS stack was augmented by addition of records from the NorthAtlantic, South Atlantic, Indian Ocean and Mediterranean to generate the Global PaleointensityStack (GLOPIS) of Laj et al. (2004). This stack was synchronized with NAPIS, and thereforeplaced on the same GISP2 chronology.Changes over glacial to interglacial cycles in deep and intermediate water circulation mayhave caused changes in the speed of bottom currents that may be detected in magnetic mineralconcentrations and grain size. Magnetic susceptibility and other magnetic concentrationparameters from around the North Atlantic Basin, along the path of NADW, have been shown torecord changes during MIS 3 that can be correlated to Dansgaard-Oeschger (D/O) events in theGreenland ice cores (Kissel et al., 1999). Lower North Atlantic Deep Water (NADW) productionis believed to have increased during interglacials and decreased during glacial periods (Broeckerand Denton, 1989). At the present time, a major contributor to NADW is Iceland-ScotlandOverflow Water (ISOW) that flows across the Gardar Drift from NE to SW at depths of ~1800-3000 meters (Bianchi and McCave, 2000). At deeper water depths (>3000 meters) in thesouthern section of the drift, Lower Deep Water (LDW) of southern hemisphere origin has beenidentified at the sea floor (Bianchi and McCave, 2000) by its high Si content (McCartney, 1992).

PAGE 107

107Upper NADW is found at water depths shallower than 2000 m in the North Atlantic and isdefined by a silicate minimum and salinity maximum (Kawase and Sarmiento, 1986).In this study we present records of relative paleointensity and benthic oxygen isotopesfrom a transect across the Gardar Drift, roughly along the path of NADW that forms a looparound the drift (Figure 5-1). A new relative paleointensity stack has been developed using threenew relative paleointensity records from this work and eight published records (Table 5-1).Magnetic property data show significant changes across the drift that can be related to changingwater masses and bottom current speed.Site LocationsThe Gardar Drift rests on a basement high on the east side of the mid-ocean ridge andstretches for about 1100 km from its northeastern end south of Iceland (<1500 m water depth) tothe southwestern end, just north of the Charlie Gibbs fracture zone (>3000 m water depth)(Bianchi and McCave, 2000) (Figure 5-1). The Gardar Drift is being formed by deposition fromdeep currents transporting detritus from Iceland and the nearby European landmass, therebycreating a smooth, thick, eastward-dipping sediment cover (Bianchi and McCave, 2000). Themain flow of Iceland-Scotland Overflow (ISOW) water travels south of Iceland between waterdepths of 1300 and 2200 meters along the eastern flank of the Reykjanes Ridge and over GardarDrift (McCave and Tucholke, 1986). The Gardar Drift may have been initiated in the late EarlyMiocene between 20 and 17 Ma by a prolonged interval of production of northern componentwater (Miller and Tucholke, 1983). Previous drilling on the Gardar Drift was carried out in 1995during Ocean Drilling Program (ODP) Leg 162, in 1983 during Deep Sea Drilling Project(DSDP) Leg 94, and by a cruise of the RV Hudson in 1991 that collected Core HU91-045-080,close to the coring site of one of the cores discussed here (JPC13).

PAGE 108

108The three cores discussed here were collected on the Gardar Drift in the summer of 2002during a cruise KN166-14 of R/V Knorr that provided site survey data for IODP Expedition 303.The cores were taken in a northeast to southwest transect across the drift (Figure 5-1). Jumbopiston Core JPC13 and accompanying gravity Core GGC12 were taken in a depression at thesouthernmost tip of the Gardar Drift in 3082 m water depth (Table 1). Core JPC2 was taken atthe northern end of the drift close to ODP Sites 984 and 983 in a water depth of 1880 m. CoreJPC5 was taken near the center of the drift in 2841 m water depth (Figure 5-1, Table 5-1). In2004, IODP Expedition 303 revisited the location of Core JPC13 and recovered a sedimentarysection down to 244 meters below seafloor (mbsf) that reached the Olduvai Subchronozone witha mean sedimentation rate of 15 cm/kyr (Channell et al., 2006).Cores JPC2 and JPC5 are typical of sediments deposited on the Gardar Drift during the lastglacial cycle and are dominated by fine silts and silty clays with subsidiary amounts ofnannofossil ooze. Core JPC13, from the southern part of the drift (Figure 5-1), is atypical andcomposed of nannofossil ooze and silty clay interspersed with numerous cmto dm-scaleintervals of diatom-rich sediments. Bodn and Backman (1996) identified diatom rich layers inCore EW93-03-17 (57.0N, 37.0W) from the west side of the Reykjanes Ridge, about 550 kmNW of Core JPC13. They described the laminated diatom ooze as being monospecific and madeup of Thalassiothrix longissima. The same species of diatom was identified at IODP Site U1304(Expedition 303 Scientists, 2006) at the same location as Core JPC13.MethodsU-channel samples were collected from the center of the split face of the jumbo pistoncores and gravity cores. The u-channel samples were measured on a 2-G Enterprises narrow-access long-core magnetometer in a magnetically shielded room at the University of Florida.Volume magnetic susceptibility was measured using a track designed for u-channel samples

PAGE 109

109(Thomas et al., 2002). The natural remanent magnetization (NRM) of u-channel samples wasstepwise demagnetized using peak alternating fields from 10 mT to 50 mT in increments of 5mT, and from 50 mT to 100 mT in increments of 10 mT. Magnetization components werecalculated using the method of Kirschvink (1980) for the 20-80 mT demagnetization interval.Anhysteretic remanent magnetization (ARM) was acquired using a 100 mT alternating field anda 50 !T DC bias field. Isothermal remanent magnetization (IRM) was acquired using a 0.5 Tfield. Both artificial remanences (ARM and IRM) were AF demagnetized at the same peak fieldsas the NRM. Normalized remanence was calculated by the dividing the NRM by either ARM orIRM at each peak demagnetization field and calculating a mean of 9 steps in 20-60 mT interval.The normalized remanence data provide relative paleointensity proxy records that can becorrelated to other cores from the Gardar Drift and the North Atlantic Ocean.Comparison of the volume magnetic susceptibility records from Core GGC12 and CoreJPC13 indicate that the upper part of Core JPC13 is stretched relative to Core GGC12. Toaccount for the stretching, the magnetic susceptibility record from the upper part of Core JPC13was correlated to the susceptibility record from Core GGC12 (Figure 5-2). This correlationgenerated a corrected depth scale for the upper part (0-955 cm) of Core JPC13. For the partsection of Core JPC13 (955-2357 cm), 404 cm was subtracted from the original depth scale,which is the amount by which the upper part of the core had to be shortened to account forobserved stretching.Samples for stable isotope analysis were collected by dissecting the Core JPC13 and CoreGGC12 u-channels into 5-cm intervals after completion of the magnetic measurements. Thesamples were washed and sieved to retain the (>63 m) sand fraction. Isotope analyses werecarried out on the benthic foraminifers Cibicidoides wuellerstorfi and, in the Holocene where C.

PAGE 110

110wuellerstorfi is scarce, Hoeglundina elegans. Benthic foraminiferal tests were cleaned in anultrasonic bath to remove fine-grained particles and soaked in 15% H2O2 to remove organicmatter. Carbon dioxide gas was produced using a Thermo Finnigan Kiel III carbonatepreparation device by reacting foraminiferal calcite with 3 drops of H3PO4 at 90C. Oxygenisotope ratios were measured on-line using a Thermo Finnigan MAT252 mass spectrometer. Allisotope results are reported in standard delta notation relative to Vienna Pee Dee Belemnite(VPDB). Analytical precision was estimated by repeated measurements of NBS-19 and was 0.06 for "18O. Both C. wuellerstorfi and H. elegans were corrected to isotopic equilibriumusing the corrections of +0.64 and -0.4 respectively (Shackleton et al., 1984). A compositesection using data from both cores was produced by splicing the isotope record from CoreGGC12 to Core JPC13 at a depth of 274 cm.Gamma ray attenuation bulk density (GRA bulk density) for Core JPC13 was measured ona GEOTEK multi-sensor core logger using u-channel samples. The calibration standard for GRAbulk density measurements was specially constructed for u-channel samples.Directional Magnetic DataOrthogonal projections of NRM data from all four cores show well-definedmagnetization components on vector end-point projections (Figure 5-3), with less than 5% of theNRM remaining after demagnetization at peak fields of 100 mT. Maximum angular deviation(MAD) values for component magnetization directions resolved in the 20-80 mTdemagnetization interval are generally below 5 (Figure 5-4). Diatom mats in Core JPC13(shaded intervals in Figure 5-4) occasionally show a high coercivity component that is notdemagnetized at peak fields of 100 mT (sample JPC13, section 4 in Figure 5-3). We attribute thisto an ARM acquired during the NRM demagnetization procedure possibly due to the presence ofultra-fine magnetite susceptible to the acquisition of ARM. Two intervals between 300-315 cm

PAGE 111

111and 1600-1920 cm in Core JPC13 show shallower than expected inclination values. Theseintervals coincide with thick layers of diatom mats, and magnetization directions may beinfluenced by spurious ARM acquisition during the NRM demagnetization, as mentioned above.Normalized RemanenceIn the absence of secondary remanence acquisition, NRM intensity of sediments dependson the intensity of the geomagnetic field at time of deposition, magnetic mineralogy, grain sizeand concentration of the magnetic remanence carriers. To produce a proxy for geomagnetic fieldintensity, the effects of down-core changes in magnetic concentration must be removed.Assuming that the magnetization is carried by single domain or pseudo single domain magnetite,we can use artificial remanences such as ARM and/or IRM to normalize for the changes in theconcentration of magnetic carriers. Tauxe (1993) stipulated that for sediments to be consideredsuitable for paleointensity studies, they must be relatively homogeneous. Magnetic susceptibilityin these cores does not change by more than an order of magnitude. Anhysteretic susceptibility(karm) plotted versus susceptibility (k) indicates uniform magnetite grain sizes in the 1-5 mrange for all three cores fining southward with Core JPC13 having the finest grain sizes (Figure5-5).Two relative paleointensity proxies (NRM/ARM and NRM/IRM) were calculated for eachof the cores from a mean of nine demagnetization steps in the 20-60 mT demagnetizationinterval. The two proxies for each of the four piston cores were normalized to 1 and compared(Figure 5-6). NRM/IRM shows departures from the NRM/ARM proxy in some intervals such asthe 1300-1450 cm interval of Core JPC5, although the shape of the curve is similar for bothpaleointensity proxies. We attribute these departures to finer magnetic grain sizes in theseintervals, as indicated by karm/k values.

PAGE 112

112Stable Isotope Data and Age ModelsThe age model for Core JPC13 was derived from the oxygen isotope stratigraphy bycorrelation to Core MD95-2042 (Shackleton et al., 2004). Core MD95-2042 was collected on thePortuguese margin at 37'N, 10'W in a water depth of 3146 m during the 1995 IMAGEScruise (Bassinot et al., 1996). A useful feature of this core is that the planktic "18O record isremarkably similar to the GISP ice-core "18O record (Figure 5-7). Shackleton et al. (2004)employed this correlation to obtain a revised GISP chronology by utilizing AMS14C ages offoraminifera in Core MD95-2042 calibrated using paired 14C and 230Th measurements on pristinecorals for the younger part of the record, and speleothems at the older end of MIS 3. Hereafter,we refer to this chronology as the Shackleton-revised GISP chronology.As the benthic oxygen isotope record from Core JPC13 can be satisfactorily correlated tothe benthic oxygen isotope record from Core MD95-2042 (Figure 5-7), the Shackleton -revisedGISP chronology can be applied to Core JPC13. This chronology can then be extended to CoresJPC2 and JPC5 using relative paleointensity correlations (Figure 5-8). Interval sedimentationrates (Figure 5-9), consistent with the Shackleton-revised GISP chronology, were calculated forCore JPC13 by correlation to Core MD95-2042 (through benthic "18O), and for Cores JPC2 andJPC5 by correlation to Core JPC13 (through relative paleointensity, Figure 5-8).Bulk Magnetic and Physical ParametersAnhysteretic susceptibility was calculated every cm down-core by normalizing the ARMby the DC bias field used to apply the ARM. ARM was then divided by the volume susceptibilityto calculate the magnetite grain size proxy karm/k for Core JPC13 (Figure 5-10). When karm isplotted versus k, using the calibration of King et al. (1983), the mean grain-sizes from all threecores can be seen to fall within a restricted range, usually less than 5 !m (Figure 5-5).

PAGE 113

113Diatom-rich sediments in Core JPC13 (shaded in Figure 5-10) are associated with MIS 5,with thin diatom-rich layers distributed within MIS 1-3. Diatom-rich intervals correspond toreduced values of susceptibility, lower and more variable density values, and relatively fine-grained magnetite as indicated by trends in karm/k (Figure 5-10). Volume magnetic susceptibilityshows decreasing values southward along the drift that can be attributed to biogenic dilution(Figure 5-11d).In Core JPC2, karm/k data can be correlated to the benthic oxygen isotope record from ODPSite 983 (Figure 5-11a), indicating that magnetite grain size at this site is varying with isotopicstage and changing abruptly at Termination I and II. Coarser magnetite grain sizes characterizethe interglacial intervals and finer grain sizes characterize the glacial intervals. Core JPC13, onthe other hand, shows a different pattern (Figure 5-11c) in which karm/k is positively correlated tothe benthic oxygen isotope record from the same core during MIS 3. Finer magnetite grain sizesoccur in the interglacials and interstadials, and coarser grain sizes in the glacials and stadials(Figure 5-11c). In Core JPC5, a similar positive correlation is only observed the later part ofMIS 5 (Figure 5-11b).Relative Paleointensity StackA stack of eleven North Atlantic relative paleointensity records has been developed usingthe records from Cores JPC2 and JPC5, as well as Core JPC13 to place the stack on the revisedGISP age model (Shackleton et al., 2004). The seven published records used to augment thestack (Figure 5-12 and Table 5-2) were chosen based on the presence of an isotopic age modelfor each record, and a correlation coefficient for each record to JPC13 of > 0.5. Although CoresJPC2 and JPC5 do not have independent isotopic age models, their relative paleointensityrecords correlate to that from Core JPC13 with a correlation coefficient exceeding 0.6 (Figure 5-8). The records were then normalized to 1 by dividing by the mean of the paleointensity proxy.

PAGE 114

114Each record was then optimally correlated to the relative paleointensity proxy from Core JPC13to place them on the Shackleton-revised GISP chronology. The resulting correlation was thenchecked for any violations of the individual isotopic age models. The records were then re-sampled at an even spacing of 500 years. An arithmetic mean of the eleven records wascalculated along with the standard deviation. A jack-knife test was used to assess errors on thestack in which each record was, in turn, excluded from trial stacks. The departure of the trialstacks from the mean value of the stack yields the estimate of standard deviation for the finalstack (Figure 5-13). Only seven of the sites used in the stack have Holocene relativepaleointensity records.The EHC06 stack, on the Shackleton-revised GISP age model generated through CoreMD95-2042, was then compared with the 30-100 ka 36Cl record from the GRIP ice core(Baumgartner et al., 1998) and with the 0-60 ka 10Be-derived paleointensity estimate (Muscheleret al., 2005), both placed on the GISP2 chronology (Figure 5-14). The stack shows minor offsetswith respect to the 10Be-derived paleointensity estimate in the 0-50 ka interval, and larger offsets(to older ages) with respect to the 36Cl record beyond 50 ka. These offsets are broadly consistentwith, and in the same sense as, the offsets between the revised GISP chronology and the standardGISP2 chronology (Shackleton et al., 2004). When the EHC06 stack is compared to the NAPISstack (Laj et al., 2000), similar offsets are observed (Figure 5-15). The NAPIS stack was placedon the GISP2 chronology using the marine to ice-core oxygen isotope correlation proposed byVoelker et al. (1998) from Core PS2644. The age of the paleointensity low associated with theLaschamp excursion differs by ~300 years in the two stacks, and the difference in age ofpaleointensity features increases to 1870 years at the ~ 60 ka paleointensity low (Figure 5-15).These age differences are very consistent with offsets between the Shackleton-revised GISP

PAGE 115

115chronology and the GISP2 chronology of Meese et al. (1997) (see Table 2 of Shackleton et al.,2004). Major differences between EHC06 and NAPIS occur in the 10-20 ka interval where theNAPIS stack shows a peak while the EHC06 stack shows a low, and between 42-50 ka where apeak in the NAPIS stack is not seen in the EHC06 record.The global GLOPIS stack (Laj et al., 2004) covers the 0-75 ka interval (Figure 5-15).EHC06 shows a good match to GLOPIS apart from the 0-12 ka interval where GLOPIS utilizesthe archeomagnetic data of Yang et al. (2000). GLOPIS shows an upward trend from 20 ka topresent that is not seen in EHC06, however, both records show a double peak in the Holocene(Figure 5-15). Comparison to the South Atlantic SAPIS stack (Stoner et al., 2002) indicates agood match in the 30-70 ka. However, in the 20-30 ka interval, the new stack shows a peakwhere SAPIS shows a significant paleointensity low (Figure 5-15). SAPIS is on a chronologyindependent of GISP and fits better with EHC06 than with either NAPIS or GLOPIS between 45ka and 70 ka. The SAPIS chronology was based on the age model for ODP Site 1089 (Hodell etal., 2001) that was constructed by correlating planktic and benthic isotope records to those fromCore RC11-83 which has 14 calibrated radiocarbon ages in the 11-41 ka interval (Charles et al.,1996).Environmental MagnetismSediment sorting takes place through differing rates of sediment transport so that anoriginally unsorted mixture is converted downstream into narrower grain size distributions(McCave et al., 1995). Sediments from the transect along the Gardar Drift show magnetic grainsizes in a restricted range indicative of sorting and deposition by bottom currents, with atendency for fining to the south accompanied by reduction in magnetic concentration (Figures 5-5 and 5-11d). Current sorting of sediment is thought to take place largely in the 10-63 m rangeor sortable silt grain-size fraction (McCave et al., 1995). Below this grain-size range, sediments

PAGE 116

116are dominantly cohesive as clay minerals (with their charge imbalances) enter the sedimentscompositional spectrum, and van der Waals forces play an important role in particle adhesion(McCave et al., 1995). Although the magnetic grain sizes estimated from karm /k values arelargely in the 1-5 m range, the karm/k ratio may be influenced by the coarsest magnetic grains inthe grain size population and these grains may lie in the sortable silt fraction.In Core JPC2, from the northern end of the Gardar Drift, coarser magnetic grain sizesoccur during interglacials whereas the finer magnetite is present in the glacials (Figure 5-11a).The magnetic grain size at this site shows a progressive coarsening down-core through MIS 3,fining in MIS 4 and coarsening in MIS 5. The location of Core JPC2, at 1880 m water depth, ispresently under the influence of ISOW (Bianchi and McCave, 2000). We interpret the karm/k datato indicate stronger ISOW bottom currents during interglacials that winnow finer sedimentleaving coarser material. In the glacials, slower ISOW current strength allows deposition of finerparticles. ISOW is an important precursor of NADW, therefore, the grain size data from CoreJPC2 implies increased NADW production during interglacials relative to glacial stages.Bianchi and McCave (2000) divided the Iceland Basin into two areas, north and south of58 30 N, based on studies of surface sediments. North of this boundary, the Gardar Drift isdraped by sediments with high (terrigenous) silt/clay ratio, containing abundant current-sortedsilt with a strong coarse-grained modal peak. They ascribed this coarser grain-size to the strengthof the near-bottom current flow, and the proximity of Iceland, consistent with the coarser grainsizes identified in the Holocene sediments in Core JPC2.In sediments from the Reykjanes Ridge, Snowball and Moros (2003) identified a saw-toothed pattern in magnetic grain size in MIS 3 during Dansgaard-Oeschger (D/O) cycles withmagnetic grain-sizes coarser during interstadials compared to stadials. The changes in magnetite

PAGE 117

117grain size were interpreted as a proxy for the speed of near-bottom currents, with a gradualintensification in current velocity followed by a sharp decrease. While these short-term changesare not identified in sediments from Core JPC2, the changes in current speed (faster currents inwarm periods and slower current in cold periods) are consistent with the changes seen in CoreJPC2. The short-term saw-toothed pattern in magnetic grain size identified by Snowball andMoros (2003) is superimposed upon a long-term trend of reduced current speed (finer grainsizes) leading up to the Last Glacial Maximum (LGM), similar to the trend seen in Core JPC2.In Core JPC5, magnetic grain size does not show a correlation to the benthic oxygenisotope record of either ODP Site 983 or Core JPC13 (Figure 5-11b), but the magnetic fraction isfiner-grained than Core JPC2 and coarser grained than Core JPC13 (Figure 5-11). This isindicative of transport by deep currents from a single source manifest by the decrease inconcentration and grain size as the currents move southward from the principal detrital source. Asimilar trend was noted by Ballini et al. (2006) along the path of NADW from just north of theFaeroe Shetland channel, to near the Reykjanes Ridge and into the Irminger Basin. The karm/kmagnetic grain size in Core JPC5, which is at a water depth of 2841 m, does not show acorrelation to the benthic isotope records, other than in the later part of MIS 5, and this mayindicate that the site is seeing a mixture of both LDW and ISOW.Core JPC13 was collected at deeper water depths (3082 m) than either Cores JPC2 or JPC5and shows the finest magnetic grain sizes of any of the sites (Figure 5-5 and 5-11). In MIS 3, themagnetic grain size proxy can be correlated to the benthic isotope record from the same core(Figure 5-11c). The oxygen isotope record from Core JPC13 can, in turn, be correlated to the "Drecord of the Vostok ice core (Jouzel et al., 1987), which reflects temperature variations overAntarctica (Figure 5-7). This implies that during the last glacial period the benthic oxygen

PAGE 118

118isotopes are recording water temperatures that reflect the Antarctic temperature record. TheAntarctic climate events A1, A2, A3 and A4 (Figure 5-7) are clearly recorded in the benthicisotope record from Core JPC13, suggesting that a water mass of southern hemisphere origin wasbathing the site during MIS 3 precluding the hypothesis that changes sediment source couldaccount for the grain size changes identified in the karm/k record. Magnetic grain size as indicatedby karm/k is coarser during Antarctic Interstadials A1-A4 indicating faster bottom currents (andpossibly increased flux of lower deep water (LDW) of Southern Hemisphere origin into theNorth Atlantic) during Antarctic warm events. During interglacials, the correlation of the benthic"18O record to the magnetic grain size record breaks down and the record shows some similarityto the Greenland Ice core record indicating changes in the bottom-water mass over the site(Figure 5-7).ConclusionsThe new EHC06 relative paleointensity stack of eleven records from the North Atlanticregion has been placed on the revised GISP chronology of Shackleton et al. (2004). This stackshows a good correlation with other paleointensity stacks (such as GLOPIS) indicating that thenew stack is providing a consistent representation of relative geomagnetic field intensity. Itshould be noted that three records utilized by EHC06 are also used in NAPIS and seven of therecords were used in GLOPIS, although the age model for EHC06 is not the same as for theother stacks. The previous North Atlantic stacks (NAPIS and GLOPIS) are placed on GISP agemodels by correlation of magnetic concentration parameters to Core PS-2644 (Kissel et al.,1999), and the "18O correlation from this core to GISP (Voelker et al., 1998). The age offsetsbetween EHC06 and NAPIS/GLOPIS are consistent with the age offsets between the GISPchronology and the Shackleton-revised GISP chronology (see Table 2 in Shackleton et al., 2004),

PAGE 119

119indicating consistency in age models. In the 45-70 ka interval, the Shackleton-revised GISPchronology applied to the EHC06 stack is more consistent with the age model for the SAPISstack, which is independent of GISP (Stoner et al., 2002), than with the NAPIS/GLOPIS agemodels that are tied to GISP chronologies. For the Holocene, EHC06 shows a double peak in theHolocene (0-12 ka), which compares well with the archeomagnetic stack of Yang et al. (2000)and with available lake sediment records for this interval (e.g. Snowball and Sandgren, 2004),although the scaling in this interval differs from that adopted in GLOPIS (Figure 5-15).Magnetite grain size, based on the parameter karm/k, shows a correlation to the benthicoxygen isotope records for Cores JPC13 and JPC2. Core JPC2 lies at 1880 m water depth and isbathed in ISOW. The karm/k values at Core JPC2 show coarser magnetic grain sizes in theinterglacials (faster bottom currents) and finer magnetic grain sizes in the glacials (slower bottomcurrents) with a trend from coarse to fine through MIS 3 indicating decreasing bottom watercurrent strength. Core JPC5 (2841 m water depth) does not exhibit a clear correlation betweenthe oxygen isotope record and the magnetic grain size implying the site is seeing some mixtureof ISOW and LDW. Core JPC13 is the deepest of the sites on the drift, and shows the finestmagnetic grain sizes. In MIS 3, the magnetic grain size record can be correlated with the benthicisotope record that in turn can be correlated to the Vostok ice core air-temperature record. Thebenthic isotope record in Core JPC13 appears to be recording bottom water temperatures that arelinked to Vostok air temperature, implying LDW of southern ocean origin. The link to karm/kindicates that magnetic grain size (and hence bottom current strength) also changes in concertwith bottom-water temperatures.

PAGE 120

120Table 5-1. Summary of the cores used in this study and the eleven cores used in the relativepaleointensity stack, including latitude, longitude, water depth, mean sedimentationrates and references. Core Latitude Longitude Sedimentation rates(cm/kyr) Water Depth(meters) Reference JPC2 61.04 N 22.92 W 8 1880 This work JPC5 56.35N 27.86 W 17 2841 This work JPC13 53.05 N 33.53 W 15 3082 This work HU90-013-012 58.92 N 47.12 W 8 2830 Stoner et al., 1995 HU90-013-013 58.20 N 48.37 W 12 3380 Stoner et al., 1995 MD95-2024 50.20 N 45.68 W 24 3448 Stoner et al., 2003 MD95-2009 62.73 N 03.98 W 21.5 1027 Laj et al., 2000 ODP Site 984 61.40 N 24.10 W 18.7 1995 Channell 1999 ODP Site 983 60.40 N 23.63 W 10.8 1660 Channell et al., 1997 ODP Site 919 62.67 N 37.47 W 15 2088 Channell, 2006 PS2644-5 67.87 N 21.77 W 12.4 777 Laj et al., 2000

PAGE 121

121 Figure 5-1. Location map for cores analyzed in this study (JPC2, JPC5 and JPC13) and thelocation of cores used in the paleointensity stack (modified after Raymo et al., 2004).Paths of major deep-water flows are indicated by arrows. Key: NADW-NorthAtlantic Deep Water, ISOW, Iceland Scotland Overflow Water, DSOW-DenmarkStrait Overflow Water, LSW-Labrador Sea Water, LDW-Lower Deep Water. P13-Core HU90-013-013, P12-Core HU90-013-012.

PAGE 122

122 Figure 5-2. a) Correlation of the magnetic susceptibility records from Core GGC12 and CoreJPC13 with tie-lines to show the correlation between the two cores. b) Magneticsusceptibility records from Core GGC12 and Core JPC13 after adjustment to accountfor the stretching in Core JPC13.

PAGE 123

123 Figure 5-3. Orthogonal projections of alternating field demagnetization data from Cores JPC2,JPC5 and JPC13. Open circles indicate the vector end-point projections on thevertical plane, closed circles indicate the vector end-point projection on the horizontalplane.

PAGE 124

124 Figure 5-4. Component inclination, declination and maximum angular deviation (MAD) valuesfor Cores JPC2, JPC5, and JPC13. Shading for Core JPC13 indicates the position ofdiatom-rich intervals. Inclination is shown by a black line, declination by blue dots,and MAD values by a red line.

PAGE 125

125 Figure 5-5. Plot of anhysteretic susceptibility (karm) versus volume susceptibility (k), opensquares represent Core JPC13, closed diamonds Core JPC5 and crosses Core JPC2.

PAGE 126

126 Figure 5-6. Paleointensity proxies: Mean NRM/ARM (red line) and mean NRM/IRM (blackline) versus depth compared for Cores JPC2, JPC5 and JPC13.

PAGE 127

127 Figure 5-7. Core JPC13 benthic oxygen isotope record (blue line), open circles indicate datafrom Core GGC12, correlated to the MD95-2042 benthic oxygen isotope record (redline) of Shackleton et al. (2004), and the Vostok ice core "D record (Jouzel et al.,1987), (on the revised GISP age model of Shackleton et al., 2004) (black line). Alsoshown is, the MD95-2042 planktic isotope record (green line) of Shackleton et al.(2004) and the GISP "18O ice core record (Grootes et al., 1997) (on the GISP2 agemodel of Meese et al., 1997).

PAGE 128

128 Figure 5-8. Relative paleointensity records from Cores JPC2, JPC5 correlated to Core JPC13 onthe revised GISP chronology (Shackleton et al., 2004) for the last 85 ka. The CoreJPC13 record is shown in blue, the Core JPC5 record in red and the Core JPC2 recordin green.

PAGE 129

129 Figure 5-9. Interval sedimentation rates for Cores JPC2, JPC5 and JPC13. The sedimentationrates for Core JPC13 were calculated using the correlation between the benthicisotope records, and the sedimentation rates for Cores JPC5 and JPC2 were calculatedbased on the correlation of the relative paleointensity records.

PAGE 130

130 Figure 5-10. Core JPC13: GRA bulk density (orange), anhysteretic susceptibility divided byvolume magnetic susceptibility (karm/k) (red), magnetic susceptibility (k) (purple), andbenthic oxygen isotope data from Core JPC13 (black closed symbols) and GGC12(open squares). Shaded intervals indicate diatom-rich sediment.

PAGE 131

131 Figure 5-11. a) Anhysteretic susceptibility divided by volume magnetic susceptibility (karm/k)from Core JPC2 (red line) plotted against the benthic "18O record (black line) fromODP site 983 (Channell et al., 1997). b) Anhysteretic susceptibility divided byvolume magnetic susceptibility (karm/k) from Core JPC5 (red line), with inverted scalerelative to (a), plotted against the benthic "18O record (black line) from ODP Site 983(Channell et al., 1997). c) Anhysteretic susceptibility divided by volume magneticsusceptibility (karm/k), with inverted scale relative to (a), plotted against the benthic"18O record for Core JPC13. d) Magnetic susceptibility for Core JPC2 (top), CoreJPC5 (middle) and Core JPC13 (bottom) showing the decrease in susceptibility fromNE to SW along the drift.

PAGE 132

132 Figure 5-12. Eleven relative paleointensity records from the North Atlantic Ocean used in thenew EHC06 stack (see Table 2 for references). The records have been correlated toCore JPC13 to place them on the revised GISP chronology (Shackleton et al., 2004)and then re-sampled at a 500-year interval.

PAGE 133

133 Figure 5-13. a) The new relative paleointensity stack (heavy black line) with the results of thejack-knife sampling. b) The eleven records used in the relative paleointensity stackplotted together on the revised GISP chronology (Shackleton et al., 2004).

PAGE 134

134 Figure 5-14. Comparison of the EHC06 paleointensity stack to 36Cl flux (red line) (Baumgartneret al., 1997) and the paleointensity estimate based on the 10Be flux (blue line)(Muscheler et al., 2005), from the GRIP and GISP Greenland ice cores, respectively.

PAGE 135

135 Figure 5-15. Comparison of the EHC06 paleointensity stack (black line) on the revised GISP age(Shackleton et al., 2004) to other paleointensity stacks: a) NAPIS (Laj et al., 2000), b)GLOPIS (Laj et al., 2004) and c) SAPIS (Stoner et al., 2002) (red lines).

PAGE 136

136CHAPTER 6RELATIVE GEOMAGNETIC PALEOINTENSITY IN THE GAUSS AND GILBERTCHRONS FROM IODP SITE U1313 (NORTH ATLANTIC)IntroductionIntegrated Ocean Drilling Program (IODP) Site U1313 (41.068 N, 32.44' W)constitutes a reoccupation of Deep Sea Drilling Project (DSDP) Site 607 located at the base ofthe upper western flank of the Mid-Atlantic Ridge, ~240 miles northwest of the Azores Islands,in a water depth of 3426 m (Figure 6-1). DSDP Site 607 utilized the hydraulic piston corer(VLHPC) and the Extended Core Barrel (XCB) to penetrate to a total depth of 311.3 metersbelow seafloor (mbsf) (Ruddiman, Kidd, Thomas, et al., 1987). DSDP Sites 607 and 609 (bothdrilled during DSDP Leg 94) constitute benchmark sites for the long-term (Myr) surface anddeep ocean environmental records from the sub-polar North Atlantic (Ruddiman et al., 1986;Raymo et al., 1989). Drilling of DSDP Leg 94 sites preceded the shipboard capability forconstruction of composite sections and continuous measurement of magnetic parameters withpass-through magnetometers. The magnetostratigraphy from Site 607 was based on discretesamples (Clement and Robinson, 1987), and indicated a clear sequence of reversals to the base ofthe Matuyama Chronozone, with poor definition of polarity zones in the Gauss and GilbertChrons.Four holes (Holes U1313A-D) were cored with the Advanced Piston Corer (APC) usingnon-magnetic core barrels to maximum depths of 308.6, 302.4, 293.4, and 152.0 mbsf,respectively, with an average recovery of 103.5% (Expedition 306 Scientists, 2006). TheHolocene to uppermost Miocene sedimentary succession at Site U1313 comprises nannofossilooze with varying amounts of foraminifers and clayto gravel-sized terrigenous components.Two major lithologic units were identified. Unit I consists of Holocene to Upper Pliocenealternating nannofossil ooze, silty clay nannofossil ooze, and nannofossil ooze with clay. Unit II

PAGE 137

137consists of Upper Pliocene to uppermost Miocene nannofossil ooze, characterized by high(~95%) and uniform carbonate concentrations. The shipboard magnetic stratigraphy at SiteU1313 was constructed on the basis of continuous measurements of natural remanentmagnetization (NRM) after AF demagnetization at peak fields of 20 mT. NRM intensities after20 mT peak field demagnetization are in the 10 to 10 A/m range above 150 mbsf, and fall tothe 10 to 10 A/m range in the lower part of the section (150-275 mbsf).U-channel samples collected post-cruise have allowed refinement of the shipboardmagnetic stratigraphy at Site U1313. The polarity stratigraphy can now be resolved for theGauss and Gilbert Chrons, and for the latest Miocene, down to ~285 meters composite depth(mcd). The nannofossil oozes have a weak low-coercivity magnetization carried by magnetite.Although volume magnetic susceptibility is weak and partially negative, it is reproducible asdemonstrated by replicate measurements on u-channel samples. Natural gamma radiation andreflectance data, collected shipboard on whole core sections, and magnetic susceptibility from u-channel samples can all be correlated to the benthic oxygen isotope stack of Lisiecki and Raymo(2004). The age model for Site U1313 is based on the correlation of the u-channel magneticsusceptibility record to a benthic oxygen isotope stack and, in part, to the astronomic solution forinsolation. The resulting reversal ages are consistent (within one obliquity cycle) withestablished reversal ages in current polarity timescales. Three relative paleointensity proxies(slopes of NRM/ARM, NRM/IRM, and NRM/ARM-acquisition) are broadly consistent witheach other, can be correlated to Pacific and Indian ocean records of the same age, and show littleevidence for the saw-tooth pattern of paleointensity decrease, observed in other records of thesame age.

PAGE 138

138MethodsAt Site U1313, u-channel samples (2x2 cm square cross-section and 150 cm in length)were collected from the center of the split face of core sections in the 120-285 meters compositedepth (mcd) interval of the shipboard-derived composite section that corresponds to the 2.4-6.2Ma interval. The natural remanent magnetization (NRM) of u-channel samples was measuredeach 1-cm down-section on a 2G-Enterprises pass-through cryogenic magnetometer at theUniversity of Florida. From 120 mcd to 170 mcd, the NRM of u-channel samples was AFdemagnetized in the 20-60 mT interval using 5 mT increments. Below 170 mcd, the weakmagnetization intensities led to the NRM being demagnetized in 2.5 mT increments in the 20-40mT peak field range. Volume susceptibility was then measured each 1-cm down-section using asusceptibility track specifically designed for u-channel samples that has a measurementresolution of a few centimeters (Thomas et al., 2003). The volume susceptibility values are verylow, but repeatable, varying in the -10-2 to 10-2 (SI) range. The mean of three replicatemeasurements, and the resulting standard deviation, constituted the susceptibility record.Anhysteretic remanent magnetization (ARM) was applied using an AF field of 100 mT anda bias DC field of 50 !T. Isothermal remanent magnetization (IRM) was imparted using a 0.5 TDC field. Both artificial remanences were demagnetized with the same AF steps used todemagnetize NRM. ARM acquisition was also measured for the 120-170 mcd interval usingincreasing AF peak fields in 5 mT increments in the 20-60 mT peak field range, and a bias DCfield of 50 !T.In the 120-170 mcd interval, principal components were calculated from the NRM datausing the method of Kirschvink (1980) applied to the 20-60 mT demagnetization interval. In thelower part of the section, below 170 mcd, a single demagnetization step (30 mT) was generallyused to determine the characteristic magnetization due to the weak magnetization intensities that

PAGE 139

139precluded the definition of magnetization components. NRM data were normalized by ARM,ARM-acquisition, and IRM to generate relative paleointensity (RPI) proxies. For the 120-170mcd interval, slopes of NRM/ARM and NRM/IRM were calculated for each measurementposition (each 1-cm) in the 20-60 mT peak field range. The linear correlation coefficient (R)provided a measure of the uncertainty in the value of the slope. Similarly, for normalization byARM-acquisition, the slope of the NRM-lost vs ARM-gained plot was calculated together with alinear correlation coefficient (R). The paleointensity proxy based on ARM-acquisition isanalogous to the pseudo-Thellier method of Tauxe et al. (1995). From 170 to ~245 mcd,NRM/ARM and NRM/IRM values were calculated using the mean of 5 values of the ratio,calculated for 5 demagnetization steps in the 20-30 mT interval (2.5 mT increments).ResultsThe magnetic polarity stratigraphy at IODP Site U1313, that is reported here, covers theinterval between the Miocene-Pliocene boundary and the onset of the Matuyama Chron. Thepolarity subchrons: Kaena, Mammoth, Cochiti, Nunivak, Sidujfall, and Thevra are all clearlydefined (Figures 6-2 and 6-3, Table 6-1), however, below ~250 mcd, polarity zones are poorlydefined. The age of the base of the u-channeled section is Late Miocene (~6.2 Ma).NRM intensities (prior to demagnetization of u-channel samples) are weak, lying in the 10-3-10-4 A/m range in the 120 mcd-170 mcd interval, and in the 10-4-10-5 A/m range below 170mcd (Figures 6-2 and 6-3). Magnetization components can be adequately defined down to 170mcd, as indicated by orthogonal projections of demagnetization data (Figure 6-4), and maximumangular deviation (MAD) values that accompany component magnetizations (Figure 6-2). In thelower part of the section, magnetization components could not be adequately resolved fromorthogonal projections because there is no systematic decrease in magnetization intensity to the

PAGE 140

140origin of the projection. The definition of the characteristic magnetization in this interval wasbased on a single demagnetization step (30 mT peak field). An initial age model was based onthe location of polarity reversals, the polarity timescale of Cande and Kent (1995), and theassumption of constant sedimentation rates within polarity chrons, yielding interval meansedimentation rates of 3-6 cm/kyr (Figure 6-5).Normalized remanence values (RPI proxies) were calculated using three normalizers:demagnetized ARM, demagnetized IRM, and ARM-acquisition. At each 1-cm interval, in the120-170 mcd interval, the normalized remanence values were calculated by determining theslopes of NRM versus the normalizer, in the 20-40 mT AF peak field range (Figure 6-6). Thelinear correlation coefficient (R) of the slope yields a measure of the uncertainty in the definitionof the slope. Although the three RPI proxies show similar variability, the NRM/ARM andNRM/ARM-acquisition records show anomalously high values, particularly in the 120-130 mcdinterval and at about 153 mcd (Figure 6-6). The NRM/IRM record does not show theseanomalous values, and was therefore used as the Site U1313 RPI proxy for comparison withother RPI records.Anhysteretic susceptibility divided by susceptibility (karm/k) is sensitive to magnetite grainsize, with high values indicating finer grains. The magnetite grain size proxy was calculated forthe 120-160 mcd interval where susceptibility values are positive. The karm/k values indicatethat, in intervals where the NRM/ARM and NRM/ARM-acquisition RPI proxies areanomalously high, the magnetite grain size is apparently relatively coarse (shaded intervals inFigure 6-7). These intervals are coincident with the darker (glacial) intervals in the section,implying the presence of coarser magnetic material, possibly ice rafting debris (IRD), during theglacial intervals. This interpretation is supported by other magnetic properties: ARM is very

PAGE 141

141weak in the affected intervals, and the IRM intensities are also relatively low. Magneticsusceptibility values in these intervals are, however, consistent with the surrounding sediment,consistent with the magnetite in these intervals being anomalously coarse grained.In the 170-220 mcd interval, NRM components could not be adequately defined because ofthe very weak NRM intensities (Figures 6-2 and 6-3). For this interval, we plot mean values forNRM/ARM and NRM/IRM for five demagnetization steps (2.5 mT increments) in the 20-30 mTpeak field range (Figure 6-8). The standard deviation about the mean gives a measure of thevariability in normalized remanence values in this demagnetization range.The RPI record (NRM/IRM) from IODP Site U1313 can be correlated to the East Pacificpaleointensity stack (EPAPIS) of Yamazaki and Oda (2005) back to 3 Ma, to the Valet andMeynadier (1993) record from the Pacific Ocean (ODP Leg 138) back to 4 Ma, and to the IndianOcean record (Core MD90-0940) of Meynadier et al. (1994) beyond 5 Ma (Figure 6-9). The agemodel for the ODP Leg 138 record is from the astronomical tuning of the gamma ray attenuation(GRA) density record (Shackleton et al., 1995). The age model for the EPAPIS stack is based onthe correlation of ARM intensity to the oxygen isotope record from ODP Site 1143 from theSouth China Sea (Tian et al., 2002). The age model for Core MD90-0940, from the IndianOcean, was determined by correlation to ODP Sites 709 and 758 using nannofossil events andreversal boundaries.Volume magnetic susceptibility records from the u-channel samples vary between lowpositive and low negative values (Figure 6-10). Triplicate measurements at 1-cm spacing foreach u-channel sample yielded standard deviations that indicate that the low susceptibility valuesare reproducible. The reflectance (L*) and natural gamma radiation records, acquired shipboardeach 5 cm (Expedition 306 Scientists, 2006), co-vary with the susceptibility record (Figure 6-11),

PAGE 142

142indicating that all three parameters are mainly controlled by biogenic (carbonate) dilution ofterrigenous (detrital) input. The susceptibility record can be correlated to the benthic oxygenisotope stack of Lisiecki and Raymo (2004), and this correlation (Figure 6-12) is used to acquirethe age model for the 2.4-5.3 Ma interval at Site U1313. The resulting age model yields a moredetailed picture of sedimentation rates, relative to that acquired from reversal ages and theassumption of constant sedimentation rates in polarity chrons (Figure 6-5). The polarity reversalages in the 2.4-5.3 Ma interval, based on correlations of susceptibility to the benthic isotopestack, are within one obliquity cycle of reversal ages in modern polarity timescales (Table 6-2),with the exception of the reversal boundaries of the Kaena and Mammoth subchrons (values insquare brackets in Table 2) that are up to 155 kyrs younger than given in the Lourens et al.,(2004) timescale. In this interval (2.8-3.4 Ma), a Gaussian-shaped filter centered at 41 kyr,applied to the magnetic susceptibility, can be compared to the astronomic solution for summerinsolation at 65N from Laskar et al. (2004) (Figure 6-13). The filtered susceptibility data showsa modulation that is consistent with the modulation of calculated summer insolation at 65 N.Minor adjustments of the filtered susceptibility record to fit the astronomic solution results inages for the Kaena and Mammoth subchrons that are consistent with published astronomically-based timescales (Table 2).DiscussionThe calibrated polarity reversal ages from Site U1313, based on a correlation of thesusceptibility record to the isotope stack of Lisiecki and Raymo (2004) and to the insolationsolution of Laskar et al. (2004), differ by less than an obliquity cycle (41 kyr) from reversal agesin modern polarity timescales (Table 2). The polarity timescales of Cande and Kent (1995) andLourens et al. (2004) utilized the Gauss-Gilbert timescale of Hilgen (1991a,b) that was based oncyclostratigraphy in the Trubi Limestones in Sicily. Hilgen (1991a,b) correlated the CaCO3

PAGE 143

143cycles in the Trubi Limestones to the precession and eccentricity orbital solutions of Berger(1978). Very minor adjustments to these ages, to make them compatible with the new astronomicsolutions of Laskar et al. (2004), were incorporated in the Lourens et al. (2004) timescale.The relative paleointensity record from IODP Site U1313 can be correlated with thePacific records of Valet and Meynadier (1993) and Yamazaki and Oda (2005), and to the IndianOcean record of Meynadier et al. (1994) (Figure 6-9). The so-called saw-tooth pattern inrelative paleointensity, whereby relative paleointensity declines within polarity chrons andabruptly recovers post-reversal, was first described from ODP Leg 138 sediments (Valet andMeynadier, 1993). This saw-tooth pattern received a great deal of attention as it may provideclues to mechanism of triggering of polarity reversals. The abrupt recovery in paleointensitypost-reversal followed by a slow decrease in intensity leading to the next reversal wasparticularly clear in subchron C2An.1n (late Gauss) in the ODP Leg 138 sediments (Figure 6-9).This saw-tooth pattern has been identified in other equatorial Pacific records, as well as IndianOcean records (Meynadier et al., 1994; Valet et al., 1994; Thibal et al., 1995) but is not seen inall RPI records from this interval (Tauxe and Shackleton, 1994; Kok and Tauxe, 1999). Kok andTauxe (1996a, 1996b) used a cumulative viscous remanence (VRM) model to explain the saw-tooth pattern of paleointensity in the Gauss Chronozone. Individual polarity zones are notuniformly affected by VRM acquisition, and VRM acquired over millions of years can beresistant to the AF demagnetization techniques. Mazaud (1996) developed an alternative modelwhereby some but not all magnetic particles in the sediment acquire NRM at the time ofdeposition, while the remaining grains acquire magnetization after deposition in the subsequentpolarity zone. This leads to a decrease in intensity of the magnetization leading up to a reversal

PAGE 144

144due to the competing effect of the magnetization acquired at the time of deposition and thatacquired progressively in opposite polarity.The RPI record from IODP Site U1313 is one of only a handful of high-sedimentation-rate RPI records for the interval from 2.5 Ma to 5.3 Ma, with sedimentation rates in the 3-6cm/kyr range. From 2.5 Ma to 4 Ma, the RPI record can be correlated to records from the PacificOcean (Valet and Meynadier, 1993; Yamazaki and Oda, 2005). Between 4 Ma and 5.3 Ma, theSite U1313 RPI record is less robust than for the younger part of the record, however, thestandard deviations associated with the mean normalized remanence values are low, indicatingconsistent normalized remanence values for different demagnetization steps. The older part ofthe RPI record can be adequately correlated to the only other RPI record available for this timeinterval, the record of Meynadier et al. (1994) from the Indian Ocean (Figure 6-9). The SiteU1313 RPI record does not clearly show the saw-tooth pattern of RPI variations firstrecognized in the Gauss Chronozone of ODP Leg 138 sediments from the Pacific Ocean (Valetand Meynadier, 1993). This observation tends to indicate that the saw-tooth pattern of RPImay be an artifact of delayed remanence acquisition as suggested by Kok and Tauxe (1996a,b)and Mazaud (1996), rather than a feature of the geomagnetic field at this time.The magnetic stratigraphy from Site U1313 covers the interval from 2.5 Ma to 6.2 Ma, andall known subchrons of the Gauss and Gilbert are recorded. Interestingly, there are no polarityexcursions in the Site U1313 record, although the relatively high sedimentation rates at SiteU1313 might be expected to reveal them, and no excursions have been unequivocally detectedelsewhere in this time interval. Cycles in the magnetic susceptibility record at Site U1313 haveallowed age-calibration by correlation to the benthic oxygen isotope stack of Lisiecki and Raymo(2004). The tuned ages of the reversal boundaries are consistent with current timescales (e.g.

PAGE 145

145Lourens et al., 2004) that all obtained their astronomically calibrated reversal ages for Gauss andGilbert subchrons from the work of Hilgen (1991a,b) in the Trubi Limestones of southern Italy,which is therefore ratified by this present work.

PAGE 146

146Table 6-1. Depth of polarity chrons from IODP Site U1313 in meters composite depth (mcd).Estimated uncertainties in the depth of the reversal boundaries are given inparentheses. ChronCK95 Depth(mcd) Age (ka)CK95 top C2An.1n (base Matuyama) 123.17(.13) 2.5810 base C2An.1n (top Kaena) 142.83(+0.21) 3.0400 top C2An.2n (base Kaena) 146.33(.10) 3.1100 base C2An.2n (top Mammoth) 150.67 (.10) 3.2200 top C2An.3n (base Mammoth) 154.50 (.10) 3.3300 base C2An.3n(base Gauss) 168.61 (.15) 3.5800 top C3n.1n (top Cochiti) 195.28(.10) 4.1800 base C3n.1n (base Cochiti) 200.83(.10) 4.2900 top C3n.2n (top Nunivak) 209.50 (.15) 4.4800 base C3n.2n (base Nunivak) 216.33 (.15) 4.6200 top C3n.3n (top Sidufjall) 224.17 (.10 4.8000 base C3n.3n (base Sidufjall) 228.83 (.15) 4.8900 top C3n.4n (top Thevra) 233.83 (.10) 4.9800 base C3n.4n (base Thevra) 245.67 (.05) 5.2300 top C3An.1n 267.22 (.20) 5.8940 base C3An.1n 277.17(.05) 6.1370 top C3An.2n 280.67(.05) 6.2690

PAGE 147

147Table 6-2. Polarity reversal ages determined at Site U1313, compared with the polarity reversalages in the polarity timescales of Cande and Kent (1995), Lourens et al., (2004) andShackleton et al. (1995). Difference between the Site U1313 reversal ages andtimescale ages are given in parentheses. Ages in square brackets indicate the ages ofreversals in the Gauss Chron based on the fit of susceptibility to the benthic oxygenisotope stack (Figure 6-12), prior to final tuning to the astronomical solution forinsolation (Figure 6-13). Chron U1313Age (Ma) CK95Age (Ma) ATNTSAge (Ma) Leg 138Age (Ma) Base Matuyama 2.616 2.581 (-0.035) 2.581 (-0.035) Top Kaena 3.074 [2.998] 3.040 (-0.034) 3.032 (-0.042) 3.046 (-0.028) Base Kaena 3.153 [3.052] 3.110 (-0.043) 3.116 (-0.037) 3.131 (-0.022) Top Mammoth 3.268 [3.107] 3.220 (-0.048) 3.207 (-0.061) 3.233 (-0.035) Base Mammoth 3.346 [3.176] 3.330 (-0.016) 3.330 (-0.016) 3.331 (-0.015) Base Gauss 3.549 3.580 (+0.031) 3.596 (+0.047) 3.594 (+0.045) Top Cochiti 4.144 4.180 (+0.036) 4.187 (+0.043) 4.199 (+0.055) Base Cochiti 4.277 4.290 (+0.013) 4.300 (+0.023) 4.316 (+0.089) Top Nunivak 4.500 4.480 (-0.02) 4.493 (-0.007) 4.479 (-0.021) Base Nunivak 4.631 4.620 (-0.011) 4.631 (0) 4.623 (-0.008) Top Sidufjall 4.760 4.800 (+0.04) 4.799 (-0.39) 4.781 (+0.021) Base Sidufjall 4.889 4.890 (+0.001) 4.896 (+0.007) 4.878 (-0.011) Top Thevra 5.009 4.980 (-0.029) 4.997 (-0.012) 4.977 (-0.032) Base Thevra 5.273 5.230 (0.043) 5.235 (0.038) 5.232 (-0.041)

PAGE 148

148 Figure 6-1. Location map for IODP Site U1313.

PAGE 149

149 Figure 6-2. Magnetic polarity stratigraphy from IODP Site U1313 in the 120-200 mcd interval.Inclination data from u-channel samples are shown by the blue line. Declination dataare shown by the dotted black line. MAD values are shown by the red line. Left plotshows NRM intensity data on a log scale at 0, 20 and 30 mT AF demagnetizationsteps. Black bars indicate normal polarity, white bars indicate reverse polarity.Chrons are labeled according the Cande and Kent (1992).

PAGE 150

150 Figure 6-3. Magnetic polarity stratigraphy from IODP Site U1313 in the 200-280 mcd intervalfrom u-channel samples using a single AF demagnetization peak field (30 mT).Inclination data (blue line), declination data (dotted black line). Left plot shows NRMintensity data on a log scale prior to AF demagnetization of u-channel samples. Blackbars indicate normal polarity, white bars indicate reverse polarity. Chrons are labeledaccording the Cande and Kent (1992).

PAGE 151

151 Figure 6-4. Vector end-point projections of AF demagnetization data for particular horizons fromu-channel samples. Open circles indicate the vector end-point projection on thevertical plane while closed circles indicate the vector end-point projection on thehorizontal plane.

PAGE 152

152 Figure 6-5. Interval sedimentation rates (black line), and age-depth plot (red line) calculatedusing the magnetic polarity stratigraphy only and interval sedimentation rates fromthe fit of susceptibility to the target benthic oxygen isotope curve (blue line).

PAGE 153

153 Figure 6-6. Gauss Chronozone at Site U1313: Three relative paleointensity proxies: slope ofNRM/ARM (red line), slope of NRM/ARM-acquisition (blue line) and slope ofNRM/IRM (green line) for the 120-170 mcd interval. All calculated in the 20-60 mTpeak field demagnetization interval. Black bar indicates normal polarity, white barreverse polarity. Chrons are labeled according the Cande and Kent (1992). Lower plotshows R-values for the slopes of NRM/ARM, NRM lost versus ARM gained andNRM/IRM.

PAGE 154

154 Figure 6-7. The magnetic grain size proxy, anhysteretic susceptibility divided by susceptibility(karm/k), for the 2.5-3.3 Ma interval. Shading shows the darker (glacial) intervalscharacterized by coarser magnetite grain size.

PAGE 155

155 Figure 6-8. Later part of the Gilbert Chronozone at Site U1313: Two relative paleointensityproxies NRM/ARM (red line) and NRM/IRM (green line) with standard deviationshown by a black bar, for the 170-220 mcd interval, calculated over five AFdemagnetization steps in the 20-30 mT peak field interval. Black indicates normalpolarity, white reverse polarity. Chrons are labeled according the Cande and Kent(1992).

PAGE 156

156 Figure 6-9. Relative paleointensity records from IODP Site U1313 (black line), the Pacificrecord of Valet and Meynadier (1993) (blue line), the Pacific EPAPIS stack(Yamazaki and Oda, 2005) (green line) and the Indian Ocean record (red line)(Meynadier et al., 1994) (between 2.5-4 Ma. Black bar indicates normal polarity,white bar reverse polarity. Chrons are labeled according the Cande and Kent (1992).Lower plot shows the IODP Site U1313 relative paleointensity record after a 9-pointsmooth (blue line) overlaid by the Indian Ocean paleointensity record (Meynadier etal., 1994) (red line) in the 3.5-5.2 Ma interval.

PAGE 157

157 Figure 6-10. Volume magnetic susceptibility from u-channel samples. Black line is a mean ofthree measurements, red bars indicate the standard deviation from the mean.

PAGE 158

158 Figure 6-11. Volume magnetic susceptibility from u-channel samples (blue line) and L*reflectance data measured shipboard (black line). Red line is smoothed naturalgamma radiation (NGR) data.

PAGE 159

159 Figure 6-12. Mean volume magnetic susceptibility (black line) tuned to the benthic oxygenisotope stack of Lisiecki and Raymo (2004) (red line) for the 2.5-5.3 Ma interval.

PAGE 160

160 Figure 6-13. Output of a gaussian filter centered on a period of 41 kyr (bandpass= 0.024 kyr-1)applied to the u-channel volume susceptibility record (black line) and the astronomicsolution for summer insolation at 65N from Laskar et al. (2004).

PAGE 161

161CHAPTER 7ODP SITE 1092 REVISED COMPOSITE DEPTH SECTION HAS IMPLICATIONS FORUPPER MIOCENE "CRYPTOCHRONS"IntroductionODP Site 1092 is located in the sub-Antarctic South Atlantic (46.7S, 7.8E, waterdepth= 1974m). A magnetic polarity stratigraphy was presented for two time intervals (1.95 to~3.6 Ma and ~5.9 to ~13.5 Ma) by Evans and Channell (2003). As is routine aboard the R/VJoides Resolution, composite stratigraphic depths (mcd) for site 1092 were constructed frommulti-sensor track (MST) data. Magnetic susceptibility, gamma ray attenuation porosity(GRAPE) and light reflectance data, were used to correlate among holes at the site and to derivean optimal record (splice) of the sedimentary section (Shipboard Scientific Party, 1999). Thecomposite depths for ODP site 1092 have now been revised using X-ray fluorescence (XRF)scans of half-cores. These new data have allowed improved correlation among the holes at thesite. The revised meters composite depth (rmcd) scheme has resulted in significant changes inhole-to-hole correlation, particularly within the interval correlative to subchron C5n.2n.Using the shipboard composite section, Evans and Channell (2003) identified four reversepolarity subzones within the polarity zone correlative to C5n.2n. The four polarity subzones wereconsidered to be correlative to "cryptochrons'' in the polarity timescale of Cande and Kent (1992)and were labeled as C5n.2n-1 to C5n.2n-4. The results imply that "cryptochrons" originallyidentified within marine magnetic anomaly 5 by Blakely (1974) signify polarity reversals ratherthan solely geomagnetic intensity minima. On the revised composite depth scale, the reversepolarity subchrons labeled as C5n.2n-2 and C5n.2n-3 by Evans and Channell (2003) become asingle subchron recorded in two different holes. The result supports the revised composite depthscale and indicates three, not four, subchrons within C5n.2n.

PAGE 162

162Revised Composite Depths (rmcd)Shipboard MST data (magnetic susceptibility, GRAPE) and light reflectance data from site1092 are often too uniform, particularly in the Upper Miocene section (120-185 mcd), for precisehole-to-hole correlation. As part of a project to assess carbonate sedimentation in the southernoceans, Westerhold and Bickert (in preparation) have measured most of the archive halves ofcores from site 1092 using the XRF core scanner at the Universitt Bremen (Rhl and Abrams,2000). Fe and Ca intensity data, measured every 2 cm, are often more variable than shipboardMST data and generally provide an efficient means of hole-to-hole correlation.Some core sections within the shipboard composite splice were not scanned for XRFbecause the working and archive halves had been too heavily sampled (partly for u-channels togenerate the magnetic data). For core sections without XRF data, shipboard magneticsusceptibility could be adequately matched to core sections with Fe intensity data from XRFscans. Some cores from outside the splice had to be stretched or squeezed to conform with theoverall depth scale of the shipboard composite section. Drilling related expansion andcontraction in these poorly consolidated sediments contributes to the lack of precise correlationof depth scales between holes (Shipboard Scientific Party, 1999). The depth scale of theshipboard composite section (with no stretching or squeezing of cores within the splice) wasadhered to in the construction of the revised composite section (rmcd).Above core 1092A-12H, the shipboard composite section (mcd) is consistent with hole-to-hole correlations based on both MST and XRF data. In the shipboard splice, core 1092C-12Hoverlies 1092A-12H. core 1092C-12H can be well correlated to 1092B-12H using XRF andmagnetic susceptibility, and this correlation is consistent with shipboard composite depths.Correlation from 1092C-12H to the underlying core in the splice (1092A-12H) is poor for bothMST and XRF data. However, 1092B-12H can be well correlated to 1092C-13H, but only when

PAGE 163

163the latter is moved 90 cm up relative to 1092B-12H. This shift is the uppermost modification ofthe composite section depths. Below this, 1092A-13H can be well correlated to its neighboringcores in the splice (1092C-13H and 1092C-14H), however, the correlation to 1092C-14Hrequires that this core should be moved up 2.58 m into 1092A-13H. The rationale for thisadjustment, based on Fe intensity (XRF) data, is illustrated in Figure 7-1.Implications for Magnetic StratigraphyAugmentation of the MST data by XRF data leads to offsets between the shipboardcomposite depths (mcd) and the revised composite depths (rmcd) that reach a maximum of 3.54m in 1092C-14H (Table 5-1). The resulting modification of the composite section provides newcomposite depths for polarity zone boundaries at site 1092 (Table 7-2 and Figure 7-2), with newage estimates for subchrons not included in the standard geomagnetic polarity timescale (GPTS).The magnetostratigraphic interpretation (the correlation of polarity zones to polarity chrons) isthe same as in Evans and Channell (2003) apart from the interval within C5n.2n. When utilizingthe shipboard composite section, this normal polarity chronozone appeared to contain four thinpolarity subzones that Evans and Channell (2003) associated with cryptochrons in marineoceanic magnetic anomaly data. The revision of the composite section indicates that there isduplication of polarity subzones that is an artifact of miscalculations in shipboard compositedepths. Polarity subchrons, that were originally labeled C5n.2n-2 (recorded in core 1092A-13H)and C5n.2n-3 (recorded in core 1092C-14H), become a single subchron (relabeled as C5n.2n-2).The realignment of cores 1092A-13H and 1092C-14H within the composite section (Figure 7-1b) results in coincidence of the records of C5n.2n-2 and C5n.2n-3 (Figure 7-3a,b). This notonly ratifies the adjustment of the composite section but also reduces the number of subchronswithin C5n.2n from four to three (Figure 7-3), consistent with the number of cryptochrons inthe GPTS of Cande and Kent (1992). Normalized remanence (mean NRM/IRM), used as a

PAGE 164

164proxy for geomagnetic paleointensity in Evans and Channell (2003), can also be well correlatedbetween cores 1092A-13H and 1092C-14H after revision of the composite depths (Figure 7-3c,d). The revised composite depths also alter the estimated duration of C5n.2n-2 and C5n.2n-3.C5n.2n-2 now has an estimated duration of 5 kyr, while the duration of C5n.2n-3 increases to 11kyr, assuming a uniform sedimentation rate within C5n.2n.In the Orera section (Spain), Abdul Aziz et al. (2003) found three normal polarity subzoneswithin C5r (two within C5r.2r and one within C5r.3r) that are not represented in the GPTS ofCande and Kent (1992, 1995). This augmented C5r could be correlated with the polarity zonesat site 1092 by moving the onset of C5r.3r to 179.41 rmcd (Krijgsman, pers comm., 2003). Thepolarity zones at site 1092 correlative to these three features have thicknesses of 1.75m (C5r.2r-1n), 0.23m (C5r.2r-2n) and 0.38m (C5r.3r-1n). This interpretation appears consistent with ahiatus at 180.48 rmcd advocated by Censarek and Gersonde (2002) from the diatombiostratigraphy. The hiatus was placed at 180.48 rmcd on the basis of the coincidence of the firstoccurrences of Denticulopsis praedimorpha and Nitzschia denticuloides, and the last occurrenceof Actinocyclus ingens var. nodus, although the ages of these diatom events are poorlyconstrained.In a recent study of chron C5 at ODP site 887, Bowles et al. (2003) found no evidence forreverse polarity subzones within C5n.2n and concluded that cryptochrons of this agerecognized in marine magnetic anomaly data represent fluctuations in geomagnetic fieldintensity. The mean sedimentation rate in C5n.2n at site 887 is 1cm/kyr, or ~ 30% of that at site1092, and it is therefore less likely that polarity intervals of the duration seen at ODP site 1092would have been recorded at site 887.

PAGE 165

165The short duration of the reverse polarity intervals within C5n.2n at site 1092 may indicatethat they are excursions rather than polarity subchrons. Various criteria have been suggested todistinguish excursions from polarity subchrons (Cande and Kent, 1992; Gubbins, 1999;Roberts and Lewin-Harris, 2000). Roberts and Lewin-Harris (2000) suggested that for a polarityexcursion to qualify as a polarity subchron it should be bounded by two field reversals, and thatdecreases in paleointensity should be apparent at both bounding reversals. Of the three subchronswithin C5n.2n at site 1092, only C5n.2n-3 exhibits a clear recovery in paleointensity between thereversals.The fundamental conclusion of Evans and Channell (2003) that short duration (5-11 kyr)polarity subchrons exist within C5n.2n, that are probably correlative to cryptochronsinterpreted from oceanic magnetic anomaly data, has not changed. However, the number ofpolarity subchrons within C5n.2n has been reduced from four to three by revision of compositedepths at site 1092.

PAGE 166

166Table 7-1. Adjusted depths of core tops from ODP site 1092. Core mbsf Ship mcd* Offsetmcd tombsf(m) Revisedmcd (rmcd) Offsetmbsf tormcd(m) Offset mcdto rmcd (m) 177-1092A12H 103 115.41 12.41 114.48 11.48 -0.93 13H 112.5 126.72 14.22 125.80 13.30 -0.92 14H 122 138.12 16.12 136.51 14.51 -1.61 15H 131.5 149.74 18.24 147.26 15.76 -2.48 16H 141 160.18 19.18 159.01 18.01 -1.17 17H 150.5 168.72 18.22 170.14 19.64 1.42 18H 160 179.85 19.85 181.57 21.57 1.72 19H 169.5 191.84 22.34 191.07 21.57 -0.77 20H 179 201.34 22.34 200.57 21.57 -0.77 177-1092B13H 111.4 122.22 10.82 121.57 10.17 -0.65 14H 121.4 134.32 12.92 132.68 11.28 -1.64 15H 130.9 146.62 15.72 144.14 13.24 -2.48 16H 140.4 155.72 15.32 154.74 14.34 -0.98 17H 149.9 166.76 16.86 165.23 15.33 -1.53 18H 159.4 175.31 15.91 176.96 17.56 1.65 177-1092C13H 108.5 119.93 11.43 119.01 10.51 -0.92 14H 118 132.82 14.82 129.28 11.28 -3.54 15H 127.5 142.70 15.20 140.22 12.72 -2.48 16H 137 155.70 18.70 151.72 14.72 -3.98 17H 146.5 166.38 19.88 162.34 15.84 -4.04 18H 156 175.79 19.79 173.91 17.91 -1.88

PAGE 167

167Table 7-2. Position of the polarity zone boundaries at site 1092 in shipboard mcd and rmcd. Agesof polarity chrons are from the geomagnetic polarity timescale (GPTS) of Cande andKent (1992, 1995). Ages of polarity subchrons not featured in the GPTS are markedby an asterisk and estimated assuming constant sedimentation rates within polaritychrons. Depth(mcd) rmcd Chron Age (Ma)CK95 121.48 120.55 Base C5n.1n 9.880 122.40 121.47 Top C5n.2n 9.920 127.92 127.00 Top C5n.2n.1 10.098* 128.01 127.15 Base C5n.2n.1 10.103* 132.88 131.95 Top C5n.2n.2 10.258* 133.02 132.16 Base C5n.2n.2 10.263* 151.31 148.60 Top C5n.2n.3 10.803* 151.43 148.95 Base C5n.2n.3 10.814* 156.60 154.12 Base C5n.2n 10.949 158.80 157.82 Top C5r.1n 11.052 160.00 159.02 Base C5r.1n 11.099 163.10 161.91 Top C5r.2n 11.476 164.54 163.37 Base C5r.2n 11.531 170.60 169.07 Top C5r.3r.1n 11.866* 170.79 169.26 Base C5r.3r.1n 11.877* 171.75 173.24 Top C5An.1n 11.935 173.00 174.42 Base C5An.1n 12.078 174.20 175.67 Top C5An.2n 12.184 174.49 175.94 Base C5An.2n 12.401 177.70 179.35 Top C5Ar.1n 12.678 178.50 180.15 Base C5Ar.1n 12.708 179.00 180.65 Top C5Ar.2n 12.775 179.42 181.13 Base C5Ar.2n 12.819 180.29 181.94 Top C5AAn 12.991 181.55 183.20 Base C5AAn 13.139 181.95 183.60 Top C5AAr.1n 13.208* 182.02 183.67 Base C5AAr.1n 13.220* 182.50 184.15 Top C5ABn 13.302

PAGE 168

168 Figure 7-1. (a) Fe intensity (XRF) data plotted as a five-point moving average on the shipboardcomposite depth (mcd) scheme, with the position of the three subchrons identified incore sections 1092C-13H-6, 1092A-13H-4 and 1092C-14H-2 (C5n.2n-1 to 3) byEvans and Channell (2003). The thick line indicates data from hole 1092A, the thinline from hole 1092B and the dashed line from hole 1092C. (b) Fe intensity (XRF)data plotted as a five-point moving average on the revised composite depth (rmcd)scheme. Line notation as for Fig. 1a. On the revised composite depth (rmcd) scheme,C5n.2n-2 and C5n.2n-3 merge into a single subchron (C5n.2n-2).

PAGE 169

169 Figure 7-2. Inclination of the characteristic magnetization component plotted against revisedcomposite depth (rmcd) for site 1092. Polarity chrons are labeled according to Candeand Kent (1992). Arrows indicate subchrons within C5n.2n, C5r.3r and C5AAr.1n.Polarity interpretation: black indicates normal polarity, white reverse polarity.

PAGE 170

170 Figure 7-3. Site 1092: (a) Inclination of the characteristic magnetization component plottedagainst shipboard composite depth (mcd) showing C5n.2n-2 and C5n.2n-3 accordingto Evans and Channell (2003). (b) Inclination of the characteristic magnetizationcomponent plotted against revised composite depth (rmcd) showing that subchronsC5n.2n-2 and C5n.2n-3 of Evans and Channell (2003) become a single subchron(now labeled C5n.2n-2). (c) Mean of the ratio of natural remanent magnetization(NRM) to isothermal remanent magnetization (IRM), calculated for ninedemagnetization steps in the 20-60 mT demagnetization range, plotted againstshipboard composite depth (mcd), (d) Mean of the ratio of natural remanentmagnetization (NRM) to isothermal remanent magnetization (IRM), calculated fornine demagnetization steps in the 20-60 mT demagnetization range, plotted againstrevised composite depth (rmcd).

PAGE 171

171CHAPTER 8ASTRONOMICAL AGES FOR MIOCENE POLARITY CHRONS C4AR-C5R (9.3-11.2 MA),AND FOR THREE EXCURSION CHRONS WITHIN C5N.2NIntroductionSite 1092 was drilled in January 1998 on Meteor Rise, close to DSDP Site 704, duringODP Leg 177 in the South Atlantic. The site produced a clear magnetic stratigraphy from 4-13Ma including the interval between C4Ar.1n and C5r.1n when sedimentation rates were ~3cm/kyr (Figure 8-1). Four short reverse polarity intervals (excursion chrons) were identifiedwithin subchron C5n.2n (Evans and Channell, 2003). This number was reduced to three due torecognition of an error in the Site 1092 composite splice, revealed by correlation of X-rayfluorescence (XRF) core scanning data, that resulted in duplication of one of the excursion zones(Evans et al., 2004).The three "cryptochrons" in C5n.2n listed by Cande and Kent (1992, 1995), hereafterreferred to as CK92/95, originate from the work of Blakely (1974) who identified three short-wavelength magnetic anomalies (tiny wiggles in the terminology of CK92/95) withinAnomaly 5 from a stack of marine magnetic anomaly (MMA) records from the NE PacificOcean. The term cryptochron expresses the uncertainty in origin of these tiny wiggles thatmay be attributed to polarity excursions/chrons or fluctuations in geomagnetic paleointensity.The resolution of Blakelys (1974) record did not allow precise estimation of the spacing of theshort wavelength anomalies. They were placed at ~300 kyr intervals within C5n.2n, and Blakely(1974) attributed these short wavelength anomalies to full polarity reversals of the geomagneticfield. These polarity subchrons within C5n.2n were included in some subsequent timescalesincluding those of Ness et al. (1980) and Harland et al. (1982, 1990), but were relegated tocryptochrons in CK92/95.

PAGE 172

172In the last decade, CK92/95 has been the standard polarity timescale used in the vastmajority of studies that involve the integration of magnetic, bioand chemostratigraphies. Thetimescale was constructed by deriving a composite geomagnetic polarity sequence from marinemagnetic anomaly spacings. In the 0-5 Ma interval, CK95 used astrochronologically-derivednumerical ages for polarity chrons available at the time (Shackleton et al., 1990; Hilgen et al.,1991). Beyond 5 Ma, using the assumption of smoothly varying spreading rates, a splinefunction was used to fit 8 radiometric age-calibration points, in the 14.8-84.0 Ma interval, to theLate Cretaceous-Cenozoic polarity record.Since the publication of CK92/95, the astrochronological calibration of the polaritytimescale has been extended beyond the last 5 Myrs. The majority of these developments havebeen incorporated into the recently published ATNTS2004 timescale of Lourens et al. (2004).For the Late Miocene, these authors used a blend of previously published astronomicaltimescales (Abdul Aziz et al., 2003; Hilgen et al., 1995; 2003). adjusted to the latest astronomicalsolutions (Laskar et al., 1993). This adjustment resulted in minor modification of the ages of thereversal boundaries from those given in the primary publications.For the polarity chrons in the C4Ar.1r -C4Ar.3r interval, Lourens et al., (2004) utilizedrecords from the Mediterranean (Hilgen et al., 1995), and from Monti dei Corvi (northern Italy)(Hilgen et al., 2003). At Monti dei Corvi, Hilgen et al. (2003) tuned a cyclic alternation of marls,marly limestones and organic-rich beds to the 65N summer insolation time series (Laskar et al.,1993). This allowed astronomic calibration of the polarity chrons in the interval from C4An tothe young end of C5n.2n. In the C5n.2n-C5Ar interval, Lourens et al. (2004) incorporated thework of Abdul Aziz et al. (2003) from the lacustrine Orera section in Spain. This sectionproduced a reliable magnetic stratigraphy from the onset of C5n.2n to C5Ar.2n. The astronomic

PAGE 173

173calibration of the reversal boundaries was accomplished using the cyclic alternation ofmudstones and dolomitic carbonates identified in the sequence.In this study, we use new oxygen isotope records from ODP Site 1092 (Paulsen et al., inpress) to astronomically calibrate polarity chrons C4Ar-C5r (9.3-11.2 Ma). Spectral analysisreveals a dominant obliquity (41-kyr) cycle in the oxygen isotope record and we use this tocalibrate the Site 1092 record to the astronomical solution (Laskar et al., 2004). This studydiffers from previous astronomical timescales for this interval (Abdul Aziz et al., 2003; Hilgen etal., 2003) in that it uses oxygen isotope records rather than lithologic cycles as the means ofastronomical calibration.Methods and ResultsAt ODP Site 1092, oxygen isotope data for the Middle to Late Miocene (7-15 Ma) weregenerated from three species of foraminifers (Figure 8-2) (see (Paulsen et al., in press). Benthicoxygen isotope data were generated from the benthic foraminifer Cibicidoides kullenbergi.Planktic oxygen isotope data were generated from two species: Globigerina bulloides andGloborotalia scitula. A power spectrum using the Blackman-Tuckey method with a Bartlettwindow, was generated in the depth domain from the stacked oxygen isotope record, using theAnalyseries program of Paillard et al. (1996) (Figure 8-3a). This showed power at twofrequencies: 0.78 m-1 and 0.25 m-1. A gaussian filter centered at 0.78 ( 0.234) m-1 was thenapplied to the stacked oxygen isotope records to extract this dominant cycle. The record was thenplaced on an initial age model based on the magnetic stratigraphy (Evans and Channell, 2003)and the ATNTS2004 timescale (Lourens et al., 2004). The dominant cycle was identified as the41-kyr obliquity cycle (Figure 8-3b) and individual (obliquity) cycles were numbered fromyoungest to oldest (1-45) (Figure 8-2). The second peak at a frequency of 0.25 m-1 was identifiedas close to the 100 kyr eccentricity period.

PAGE 174

174The oxygen isotope stack was tuned to an astronomical target curve, which was derivedfrom the sum of normalized values (minus the mean and divided by the standard deviation) ofeccentricity (E), obliquity (T) and negative precession (P) (E+T-P) (Laskar et al., 2004). Tuningof the isotope record was only possible in the 9.3-11.2 Ma interval due to lower sedimentationrates and condensed horizons outside this interval.For Neogene sections, it is often assumed that the 41 kyr component of "18O is globallycorrelative, and not likely to be variable in phase relative to orbital forcing (Clemens, 1999).Much of the power in the climate spectrum since the early Oligocene appears to be concentratedin the obliquity band (Zachos et al., 2001). At Site 1092, the final age model was obtained bytuning the initial age model (from ETP tuning) until the coherence calculated using cross-spectral analysis was maximized between the filtered "18O record (filter centered at 41 kyr) andthe orbital obliquity signal. Coherence between the oxygen isotope stack and ETP is close to oneat the obliquity frequency (Figure 8-3c). The 1.2 Myr modulation of the obliquity cycle is clearlyvisible in the filtered isotope record (Figure 8-4) facilitating an unambiguous match to the orbitalobliquity target. In this way, we produced an orbitally tuned age model for the 9.3-11.2 Mainterval at Site 1092.The resulting astronomically tuned ages for C5n.2n are 44 kyrs younger at the onset, and19 kyrs younger at the termination, than ages in ATNTS2004 (Lourens et al., 2004). The newages are also significantly different from the CK92/95 ages, with the onset of C5n.2n being 47kyrs older and the termination 62 kyrs older (Table 8-1). Although the difference is close to oneobliquity cycle, an offset by one obliquity cycle would give an inappropriate match between the"18O records and the ETP curve (Figure 8-4). For example, if we shift the oxygen isotope recordsone obliquity cycle younger then the light "18O values of G. bulloides and C. kullenbergi in the

PAGE 175

175interval 10.78 to 10.72 Ma (Figure 8-4) would be located in the ETP minimum at ~10.7 Mawhich can be considered unrealistic. Interval sedimentation rates at Site 1092, calculated for theC4Ar-C5r interval using the age-depth tie points from the tuning of the oxygen isotope records,vary from 1.7 cm/kyr to 3.7 cm/kyr for the entire interval and vary from 2.5 cm/kyr to 3.7 cm/kyrfor C5n.2n (Figure 8-5).Comparison with Other TimescalesThe greatest potential source of error in the age model is uncertainty in the orbital solution,which may be as high as 20 kyr at 10 Ma (Lourens et al., 2004; Laskar et al., 2004), whereas ourtuning errors should be no more than a few thousand years. A component of the uncertainty inplacement of the reversal boundaries can be estimated using the mean sedimentation rate (3cm/kyr) and the response function width (at half-height) of ~4.5 cm for the 2G Enterprises u-channel magnetometer, giving a nominal error of ~2 kyrs for each reversal boundary. This errorwas mitigated by deconvolution (Guyodo et al., 2002) of the u-channel record across theexcursional intervals (Evans and Channell, 2003), resulting in a modified error estimate of ~1kyr for the C5n.2n polarity excursions. These estimates do not include error in placement ofpolarity zone boundaries associated with delayed remanence acquisition, referred to as post-Depositional Remanent Magnetization (pDRM). Following Channell and Guyodo, (2004), thesediment lock-in beneath the bioturbated surface layer in pelagic sediments is abrupt, and cantherefore the lock-in depth can be estimated from the mean sedimentation rate and the thicknessof the surface bioturbated mixed layer (<10 cm in most pelagic environments (Trauth et al.,1997; Smith and Rabouille, 2002)). In the case of Site 1092, assuming a 10 cm bioturbatedsurface layer, the delay in remanence acquisition would be about 3 kyr.A data gap occurs in the oxygen isotope records at 155.8-157.3 revised meters compositedepth (rmcd). The gaussian filter identifies two obliquity cycles in this data gap (Figure 8-2). If

PAGE 176

176we assume that three cycles occurred in this gap, the sedimentation rates would be anomalouslylow (2.3 cm/kyr), while a single cycle causes an increase in sedimentation rates (5.1 cm/kyr).Two obliquity cycles yields sedimentation rates of 3.2 cm/kyr consistent with those adjacent tothis interval. The revised composite section is well constrained in this interval (Evans et al.,2004), and there are no indications in physical properties of a likely change in sedimentationrate.Comparison of the new astronomically-tuned ages for subchrons C4Ar.1n to C5r.1n (9.3-11.2 Ma) with ATNTS2004 (Lourens et al., 2004) reveal differences of 5-48 kyrs (Table 8-1). Alarge part of the age discrepancy is probably due to the low resolution of the paleomagneticrecord in the Monti de Corvi section (Hilgen et al., 2003) that provides the basis for theATNTS2004 timescale in this interval. In this section, the polarity reversals are poorly definedand the pattern fit of polarity zones to polarity chrons is ambiguous, due to weak and unstablemagnetic remanence. Hilgen et al. (2003) gave errors of 25-77 kyrs for the astronomical ages forthe reversal boundaries at Monti dei Corvi, due largely to poor definition of polarity zones (seeTable 3 of Hilgen et al., 2003). For C4Ar.1n-C4Ar.2n, the differences between the astronomicalages obtained at Site 1092 and those obtained at Monti dei Corvi are within these error estimates,and the differences reach 71 kyrs for subchron C5n.1n where the error estimates at Monti deiCorvi are largest.Site 1092 and CK92/95 ages differ by ~100 kyrs in the interval between C4Ar.1n andC4Ar.2n. Between the top of chron C5n.1n and the base of C5r.1n, the differences are 44-67 kyrs(Table 6-1). This narrow range indicates that the durations of subchrons in this interval are veryconsistent between the two timescales. CK92/95 relies on two calibration points for the middle tolate Miocene interval. The first is placed at the older end of subchron C3n.4n with an age of 5.23

PAGE 177

177Ma from the astrochronological work of Hilgen (1991). The second age calibration point at 14.8Ma at the young end of subchron C5Bn, was derived from radioisotopic age constraints on thecorrelative N9/N10 foraminifer zone boundary (see Cande and Kent, 1992).Shackleton et al. (1995) constructed a timescale for the Late Neogene based on gamma rayattenuation (GRA) bulk density data from sediment cores obtained during ODP Leg 138. For the0-6 Ma interval, cycles identified in the GRA bulk density data were tuned to the orbitalinsolation record of Berger and Loutre (1991). The Late Miocene (6-14.8 Ma) timescale wasrecalibrated using two tie-points at 5.875 Ma (termination of C3An) and 9.64 Ma (termination ofC5n) and fitting a cubic-spline to estimate spreading rates in the manner adopted by CK92. Theage control point at the termination of C5n (9.64 Ma) was generated by taking the radiometricage of 9.66 +/-0.05 Ma from Baksi (1992) and adjusting it to the closest age that allowed theGRA bulk density to be matched directly to the insolation record. The ages obtained byShackleton et al. (1995) are 153-225 kyrs younger than those obtained for Site 1092 (Table 8-1).There are several possible factors that could contribute to these differences. (1) The sedimentrecord from the ODP Leg 138 sites may not be complete in the older part, possibly attributable touse of the XCB coring system. (2) The quality of the GRA bulk density data deteriorates, and thematch to the insolation record becomes ambiguous, in the older part of the record. (3)Sedimentation rates are low (~1-2 cm/kyr) in the Late Miocene at Leg 138 sites (Shackleton etal., 1995).The Monti Gibliscemi section in Sicily (Italy) is a deep marine cyclically beddedhemipelagic succession of Miocene age (Hilgen et al., 2000). Due to weak magnetic intensitiesand overprinting, a magnetic stratigraphy was not obtained from the section. Hilgen et al. (2000)]indirectly estimated astronomical ages for polarity chron boundaries by transferring the

PAGE 178

178astronomical ages of calcareous nannofossil events at Monti Gibliscemi to ODP Leg 138 sites inthe equatorial Pacific that have reliable magnetic stratigraphies (Schneider, 1995). Linearinterpolation of sedimentation rates between nannofossil datums yielded ages for polarity chronboundaries (Hilgen et al., 2000). In the interval from C5n.1n to the base of C5n.2n, the ages fromMonti Gibliscemi are consistently older than ages from Site 1092 with the mean difference being~40 kyrs (Table 8-1). For subchron C5r.1n, the ages are younger than those obtained in thisstudy by 37 and 38 kyrs at the young and old end of the subchron, respectively.Excursion ChronsPrevious estimates of the duration of the polarity excursion chrons within C5n.2n fromODP Site 1092 have relied on the assumption of constant sedimentation rates within the chron(Evans and Channell, 2003). Based on a mean sedimentation rate within C5n.2n of ~3 cm/kyr,the excursion chrons were estimated to have a duration of 6-11 kyrs. The new astronomicalcalibration yields durations for these excursion chrons of 3-4 kyrs (Table 8-1).DSDP Site 608 has recently yielded a revised magnetic stratigraphy for the Middle to LateMiocene (Krijgsman and Kent, 2004). Discrete samples collected every 2.5 cm at Site 608indicate three excursions within C5n.2n, albeit represented by single samples, with estimateddurations of 1-6 kyrs. Three reverse polarity intervals at ODP Site 884 on the Detroit Seamountin the NW Pacific Ocean were placed within C5n.2n (Roberts and Lewin-Harris, 2000), andwere calculated by the authors to have durations of 6, 26 and 28 kyrs. Ambiguities in theinterpretation of the magnetic stratigraphy at Site 884, and the apparent duration of these reversepolarity intervals, makes it unlikely that they correlate to the excursional directions identified atSite 1092 (see Evans and Channell, 2003).Roperch et al. (1999) studied a 4.5 km thick middle Miocene continental red bed section inthe Bolivian Altiplano. Magnetostratigraphic results indicate that the sequence was deposited

PAGE 179

179during the 14-9 Ma interval, and has a mean sedimentation rate of 97 cm/kyr in the 11.5-9.2 Mainterval. Roperch et al. (1999) identified one reverse polarity interval represented by fivesamples (at 3714-3719 m above base of section) within the normal polarity interval correlative toC5n.2n. Using an estimate for the mean sedimentation rate within C5n.2n (97 cm/kyr), thisreverse interval has a duration of ~5 kyrs. The Ulloma tuff lies ~100 meters below the reversepolarity zone and has yielded an age of 10.35 +/-0.06 Ma from 40Ar/39Ar dating of sanidinecrystals (Marshall et al., 1992). Assuming a constant sedimentation rate from the top of thepolarity zone correlative to C5n.2n to the Ulloma tuff the reverse polarity zone has an age of10.21 Ma and a duration of ~8 kyrs.Bowles et al. (2003) studied the sedimentary section at ODP Site 887 from the NorthPacific that covers the C5n interval. The core was sampled using discrete samples at 2.5 cmspacing. The mean sedimentation rate within C5n.2n (1 cm/kyr) implies a sampling resolution of2500 yrs, however no reverse polarity intervals were detected within C5n.2n. In view of thesedimentation rates at Site 887, it is possible that polarity intervals of the duration seen at ODPSite 1092 would not have been recorded using this sampling regime.The Bowers et al. (2001) deep-tow marine magnetic anomaly (MMA) record from thesouthern East Pacific Rise (EPR) (Figure 8-7) is one of the most detailed MMA records for thistime interval with an the average half-spreading rates of 42 mm/yr. In Figure 8-7, we correlatethe Site 1092 paleointensity record from Evans and Channell (2003) to the deep-tow MMArecord. The three brief excursion chrons observed in C5n.2n at Site 1092 can then be placed intothe deep-tow MMA record (arrows from below in Figure 8-7). The preferred correlationbetween the relative paleointensity record from Site 1092 and the deep tow magnetic anomalyrecord yields a correlation of Site 1092 excursion chrons to the deep-tow record that differs from

PAGE 180

180the Bowers et al. (2001) correlation (arrows from top in Figure 8-7) of CK92/95 tiny wiggles(cryptochrons) from the North Pacific stack to the EPR deep-tow record.Oxygen isotope records from ODP Site 1092 have allowed astronomic calibration of theages of eight polarity chron boundaries (C4Ar.1n-C5r.1n), and of three excursion chrons withinC5n.2n (Evans and Channell, 2003; Evans et al., 2004). This is the first time astronomicallycalibrated ages have been assigned to the excursion chrons within C5n.2n, and they indicatedurations of 3-4 kyr. This duration estimate is consistent with the model of Gubbins (1999) thatpredicts that excursions should have durations less than the magnetic diffusion time (3 kyrs) forthe inner core (Hollerbach and Jones, 1995). The duration of these excursions is less than theduration for reversal transitions such as the Matuyama-Brunhes boundary (5-10 kyrs, e.g.Channell and Kleiven, 2000) implying that the outer core must maintain the opposite ortransitional polarity state for greater than ~3 kyrs to allow the outer core field to diffuse throughthe inner core and hence stabilize the outer core field (Gubbins, 1999). The duration forexcursions, such as those within C5n.2n, which appear as abrupt swings to reverse polarity andreturn to normal polarity, was apparently insufficient for establishment of a prolonged reversepolarity interval.

PAGE 181

181Table 8-1. Astronomical ages from recent timescales compared with those inferred at ODP Site1092. Numbers in parentheses indicate the difference between Site 1092 estimates(this paper) and other timescales. CK95Cande and Kent (1995), ATNTS2004-Lourens et al., (2004), A2003Abdul Aziz et al. (2003), S1995Shackleton et al.(1995), H1995Hilgen et al. (1995), H2000Hilgen et al. (2000), H2003Hilgen etal. (2003). Subchron Depth (rmcd) 1092 age (ka)(errors) CK95 age (ka) ATNTS2004(ka) A2003 (ka) S1995 (ka) H1995(ka) H2000 (ka) H2003 (ka) Top C4Ar.1n 105.13 (.03) 9351 () 9230 (-121) 9312 (-39) 9142 (-209) 9364 (+13) Base C4Ar.1n 106.96 (.19) 9443 () 9308 (-135) 9409 (-34) 9218 (-225) 9428 (-15) Top C4Ar.2n 112.60 (.05) 9671 () 9580 (-91) 9656 (-15) 9482 (-189) 9629 (-42) 9652 (-19) 9687 (+16) Base C4Ar.2n 115.60 (.04) 9765 () 9642 (-123) 9717 (-48) 9543 (-222) 9740 (-25) 9762 (-3) 9729 (-36) Top C5n.1n 116.80 (.05) 9807 () 9740 (-67) 9779 (-28) 9639 (-168) 9841 (+34) 9770 (-37) Base C5n.1n 120.79 (.05) 9942 () 9880 (-62) 9934 (-8) 9775 (-167) 10000 (+58) 9871 (-71) Top C5n.2n 121.61 (.07) 9968 () 9920 (-62) 9987 (+19) 9815 (-153) 10037 (+66) 10004 (+36) Top C5n.2n.1 127.00 10154 () 10091 (-63) Base C5n.2n.1 127.09 10157 () 10093 (-64) Top C5n.2n.2 131.96 10309 () 10248 (-61) Base C5n.2n.2 132.10 10313 () 10252 (-61) Top C5n.2n.3 148.83 10826 () 10782 (-44) Base C5n.2n.3 148.95 10829 () 10785 (-44) Base C5n.2n 154.12 (.09) 10996 () 10949 (-47) 11040 (+44) 11043 (+47) 10839 (-157) 10998 (+2) Top C5r.1n 157.71 (.12) 11108 () 11052 (-57) 11118 (+10) 11122 (+14) 10943 (-165) 11071 (-37) Base C5r.1n 159.03 (.03) 11149 () 11099 (-50) 11154 (+5) 11158 (+9) 10991 (-158) 11111 (-38)

PAGE 182

182 Figure 8-1. Magnetic component inclination for the C4Ar.1n-C5r.1n interval from ODP Site1092 (Evans and Channell, 2003) compared to the geomagnetic polarity timescale ofCande and Kent (1992; 1995). rmcd= revised meters composite depth.

PAGE 183

183 Figure 8-2. Oxygen isotope records from the C4An-C5r.1n interval at ODP Site 1092. The topframe shows the output of a gaussian filter centered at a frequency of 0.78 m-1 appliedto the stacked "18O record. The stacked "18O record with numbered obliquity cyclesis shown superimposed on the same record with a 5-point smoothing. The three "18Orecords from different planktic and benthic foraminiferal species were used togenerate the stack.

PAGE 184

184 Figure 8-3. a) Power spectrum generated from the oxygen isotope stack in the depth domain(solid line). b) Dashed line is the power spectrum generated from the ETP target(Laskar et al., 2004) and the solid line is the power spectrum generated from thestacked oxygen isotope records after tuning. c) Coherence between the "18O stack andthe ETP target curve, line indicates 95% confidence limit for coherence peaks.

PAGE 185

185 Figure 8-4. Upper plot shows the correlation of the filtered (filter centered at 0.0244 0.0073kyr1) oxygen isotope stack to the astronomical solution for obliquity (Laskar et al.,2004). Lower plot shows the correlation of the three oxygen isotope records from Site1092 to the ETP solution (Laskar et al., 2004). Crosses mark the tie points betweenthe oxygen isotope stack and the ETP curve. Shaded areas indicate critical intervals inthe correlation between the records that facilitate an unambiguous match between theoxygen isotope record and the ETP astronomic solution.

PAGE 186

186 Figure 8-5. Interval sedimentation rates for the C4Ar.1n-C5r.1n interval calculated using the newastrochronology. Asterisks indicate the position of polarity excursions within C5n.2n.

PAGE 187

187 Figure 8-6. Comparison of the age estimates of polarity chrons at ODP Site 1092 (this paper) tothe timescale of Cande and Kent (1992; 1995), to the ATNTS2004 timescale(Lourens et al., 2004), and to the timescales of Hilgen et al. (1995; 2000), and AbdulAziz et al. (2003).

PAGE 188

188 Figure 8-7. The Site 1092 relative paleointensity record for C5n.2n (base), the deep-towmagnetic anomaly record from the East Pacific Rise at 19S (middle) and the revisedNorth Pacific Stack (Bowers et al., 2001). Numbering on the revised North Pacificstack is after [35]. Arrows from above indicate the proposed correlation (Bowers etal., 2001) of CK92 cryptochrons to the revised N. Pacific Stack and the EPR 19Sdeep-tow record. Arrows from below indicate our preferred correlation of the polarityexcursion chrons to the deep-tow record.

PAGE 189

189CHAPTER 9CONCLUSIONS AND FUTURE WORKThe work presented in this dissertation illustrates the spectrum of timescales upon whichthe sedimentary record of the geomagnetic field can be used as a tool for stratigraphiccorrelation. The amount of information that can be gained from a particular sedimentary recorddepends on a number of factors: the type of sediment, the magnetic remanence carrier, the lengthof the record and the geographic location the core was collected. As such, this work hasdemonstrated the enormous possibilities that sedimentary records of the geomagnetic field havein terms of improving our understanding of changes in the geomagnetic field over time, theiruses in stratigraphic correlation on varying timescales, and the importance of environmentalmagnetism to paleoclimatology. As more sedimentary cores are collected from the worldsoceans, our understanding of the paleomagnetic field can only increase.In Brunhes age sediments from the North Atlantic a combination of relative paleointensityand oxygen isotope records have been used to develop paleointensity-assistedchronostratigraphies. Detrital layers identified on the Eirik drift have been placed in apaleointensity-assisted chronostratigraphic framework, allowing improved correlation to otherrecords of detrital layers from the North Atlantic. A new relative paleointensity stack for the 0-85 ka interval has been developed using three new paleointensity records and eight existingrecords. This stack has been placed on the Shackleton-revised GISP chronology by correlation ofa benthic oxygen isotope record to Core MD95-2042 from the Portuguese Margin.Miocene to Pleistocene age sediments from the Pacific and Atlantic Oceans have producedreliable magnetic stratigraphies back to 12 Ma. Integration of the magnetic stratigraphy withcycle stratigraphy has allowed astronomic calibration of the interval from 9.3-11.2 Ma at ODPSite 1092, in the 1-6 Ma interval at ODP Site 1208 and between 2.5 Ma and 6 Ma at IODP Site

PAGE 190

190U1313. Integration with biostratigraphic data has resulted in a new Late Miocene to Recentplanktonic foraminifer biostratigraphic zonation for the northwest Pacific. In Pliocene agesediments from IODP Site U1313 (a re-occupation of DSDP Site 607) a record of relativepaleointensity between 2.5 Ma and 6 Ma is one of only a handful of records for this interval.Even though the geomagnetic polarity timescale has recently been revised with the entireNeogene section of the timescale being astronomically calibrated, much work is still required onolder (Paleogene) parts of the timescale. I have been working on a collaborative project withThomas Westerhold, Ursula Rohl and others to improve the Paleogene GPTS. This has resultedin the compilation of magnetostratigraphic results from two ODP Legs (198 and 208) along withXRF scanning data and other physical properties from the cores, resulting in the firstastronomically calibrated timescale for the Paleocene (Westerhold et al., in preparation). Byintegrating published ODP data and land-based records with ODP Leg 198 and 208 sites, aPaleogene cyclostratigraphy has been accomplished (Westerhold et al., in preparation). The nextphase of this work will be on Eocene age sediments from ODP Leg 198.

PAGE 191

191LIST OF REFERENCESAbdul Aziz, H., Krijgsman, W., Hilgen, F.J., Wilson, D.S., Calvo, J.P., 2003. An astronomicalpolarity timescale for the late middle Miocene based on cyclic continental sequences. J.Geophys. Res.108, 2159, doi:10.1029/2002JB001818.Arthur, M.A., Srivastava, S.P., Kaminski, M., Jarrad, R., Osler J, 1989. Seismic stratigraphy andhistory of deep circulation and sediment drift development in Baffin Bay and the LabradorSea. In: Srivastava, S.P., Arthur, M.A., Clement, B., (Eds.), Proceedings of the ODP, Sci.Results 105. Ocean Drilling Program. College Station, TX, 957-975.Backman, J., Raffi, I., 1997. Calibration of Miocene nannofossil events to orbitally tunedcyclostratigraphies from Ceara Rise. In: Shackleton, N.J., Curry, W.B., Richter, C.,Bralower, T.J. (Eds.). Proceedings of the ODP, Sci. Results 154. Ocean Drilling Program.College Station, TX, 83-99.Backman, J., Pestiaux, P., 1987. Pliocene Discoaster abundance variations, Deep Sea DrillingProject Site 606: Biochronology and paleoenvironmental implications. In: Ruddiman,W.F., Kidd R.B., Thomas, E., Shipboard Scientists. Init. Repts. DSDP 94. Washington,D.C., US Government Printing Office, 903-910.Baksi, A., 1992. A40Ar/39Ar age for the termination of Chron 5; a new calibration point for theMiocene section of the GPTS. Trans. Am. Geophys. Union (EOS) 73, 630.Ballini, M., Kissel, C., Colin, C., Richter, T., 2006. Deep-water mass source and dynamicassociated with rapid climatic variations during the last glacial stage in the North Atlantic:A multiproxy investigation of the detrital fraction of deep-sea sediment. Geochem.Geophys. Geosys. 7, doi:10.1029/2005GC001070.Bassinot, F., Labeyrie, L., Shipboard Scientific Party, 1996. IMAGES MD101, a bord duMarion-Dufresne du 29 mai au 11 juillet 1995, 217 pp., Inst. Francais pour la Rech. et laTechnol. Polaires, Plouzane, France.Baumgartner, S., Beer, J., Masarik, J., Wagner, G. Meynadier, L., Synal, H.-A., 1998.Geomagnetic modulation of the 36Cl flux in the GRIP ice core, Greenland. Science 279,1330-1332.Berger, A., 1988. Milankovitch theory and climate. Rev. Geophys. 26, 624-657.Berger, A. 1978. Long-term variations of daily insolation and Quaternary climatic change. J.Atmos. Sci. 35, 2362-2367.Berger, A., Loutre, M.F., 1991. Insolation values for the climate of the last 10 million years.Quat. Sci. Rev. 10, 297-317.

PAGE 192

192Berggren, W.A., Kent, D.V., Swisher, C.C., Aubry, M-P., 1995a. A revised Cenozoicgeochronology and chronostratigraphy. In: Berggren, W.A., Kent, D.V., Aubry, M-P.,Hardenbol, J., (Eds.). Geochronology time scales and global stratigraphic correlation.SEPM Special publication 54, pp. 129-206.Berggren, WA., Hilgen F.J., Langereis, C.G., Kent, D.V., Obradovich, J.D., Raffi, I., Raymo,M.E., Shackleton, N.J., 1995b. Late Neogene chronology: New perspectives in high-resolution stratigraphy. GSA Bulletin 107, 1272-1287.Berggren, W.A., Kent, D.V., van Couvering J.A., 1985. Neogene geochronology andchronostratigraphy. In: Snelling, N.J. (Ed.). The chronology of the geological record.Geological Society of London Memoir 10, 211-250.Bianchi, G.G., McCave, I.N., 2000. Hydrography and sedimentation under the deep westernboundary current on Bjorn and Gardar Drifts, Iceland Basin. Marine Geology 165, 137-169.Blakely, R.J., 1974. Geomagnetic reversals and crustal spreading rates during the Miocene. J.Geophys. Res. 79, 2979-2985.Bodn, P., Backman, J., 1996. A laminated sediment sequence from northern North AtlanticOcean and its climatic record. Geology 24, 507.Bolli, H.M., Saunders, J.B., 1985a. Oligocene to Holocene low latitude planktonic foraminifera.In: Bolli, H.M., Saunders, J.B., Perch-Nielsen, K. (Eds.). Plankton Stratigraphy,Cambridge, Cambridge University Press, pp. 155-262.Bolli, H.M., Saunders, J.B., Perch-Neilsen, K., 1985b. Introduction to the foraminiferal chapters.In: Bolli, H.M., Saunders, J.B., Perch-Nielsen, K., (Eds.). Plankton Stratigraphy,Cambridge, Cambridge University Press, pp. 11-16.Bond, G.C., Showers, W., Elliot, M., Evans, M., Lotti, R., Hajdas, I., Bonani, G., Johnson, S.,1999. The North Atlantic's 1-2 kyr climate rhythm: relation to Heinrich Events,Dansgaard/Oeschger Cycles and the Little Ice Age. In: Clark, P.U., Webb R.S., Keigwin,L.D., (Eds.). Mechanisms of global climate change at millennial time scales. GeophysicalMonograph 112, 35-58.Bowers, N.E., Cande, S.C., Gee, J., Hildebrand, J.A., Parker, R.L., 2001. Fluctuations of thepaleomagnetic field during chron C5 as recorded in near bottom marine magnetic anomalydata. J. Geophys. Res. 106, 26,379-26,396.Bowles, J., Tauxe, L., Gee, J., McMillan, D., Cande S., 2003. Source of tiny wiggles in ChronC5: A comparison of sedimentary relative intensity and marine magnetic anomalies.Geochem. Geophys. Geosyst. 4, 1049, doi:10.1029/2002GC000489.Bown, P.R., 2005. Cenozoic calcareous nannofossil biostratigraphy, ODP Leg 198 Site 1208(Shatsky Rise, northwest Pacific Ocean). In: Bralower, T.J., Premoli-Silva, I., Malone,M.J. (Eds.), Proceedings of the ODP, Sci. Results 198, 1 [Online].

PAGE 193

193Bown, P.R.,Young, J.R., 1998. Techniques. In Bown, P.R. (Ed.), Calcareous nannofossilbiostratigraphy: Dordrecht, The Netherlands (Kluwer Academic Publ.), pp. 16.Bralower, T.J., Premoli Silva, I., Malone, M.J., The Shipboard Scientific Party, 2002.Proceedings of the ODP, Init. Repts. 198 [CD-ROM]. Ocean Drilling Program, TexasA&M University, College Station, TX.Broecker, W., Denton, G.H., 1989. The role of ocean-atmosphere reorganizations in glacialcycles. Geochimica et Cosmochimica Acta 53, 2465-2501.Bukry, D., 1973. Low-latitude coccolith biostratigraphic zonation. In: Edgar, N.T., Saunders,J.B., and Shipboard Scientists. Init. Repts. DSDP, 15, Washington (U.S. Govt. PrintingOffice), 685.Bukry, D., 1975. Coccolith and silicoflagellate stratigraphy, northwestern Pacific Ocean, DeepSea Drilling Project Leg 32. In: Larson, R.L., Moberly, R., Shipboard Scientists. Init.Repts. DSDP, 32. Washington (U.S. Govt. Printing Office), 677.Cande, S.C., Kent, D.V., 1992. A new geomagnetic polarity timescale for the late Cretaceousand Cenozoic. J. Geophys. Res. 97, 13917-13951.Cande, S.C., Kent, D.V., 1995. Revised calibration of the geomagnetic polarity timescale for thelate Cretaceous and Cenozoic, J. Geophys. Res. 100, 6093-6095.Censarek, B., Gersonde, R., 2002. Miocene diatom biostratigraphy at ODP sites 689, 690, 1088,1092 (Atlantic sector of the Southern Ocean). Marine Micropaleontology 45, 309-356.Channell, J.E.T., 1999. Geomagnetic paleointensity and directional secular variation at OceanDrilling Program (ODP) Site 984 (Bjorn Drift) since 500 ka: Comparisons with ODP Site983 (Gardar Drift). J. Geophys. Res. 104, 22,937-22,951.Channell, J.E.T., D.A. Hodell, B. Lehman, 1997. Relative geomagnetic paleointensity and "18Oat ODP Site 983 (Gardar Drift, North Atlantic) since 350 ka. Earth Planet. Sci. Lett. 153,103-118.Channell, J.E.T., Labs, J., Raymo, M.E., 2003. The Reunion subchronozone at ODP Site 981(Feni Drift, North Atlantic). Earth Planet. Sci. Lett. 215, 1-12.Channell, J.E.T., Mazaud, A., Sullivan, P., Turner, S., Raymo, M.E., 2002. Geomagneticexcursions and paleointensities in the 0.9-2.15 Ma interval of the Matuyama Chron at ODPSites 983 and 984 (Iceland Basin). J. Geophys. Res. 107, doi:10.1029/2001JB000491.Channell, J.E.T., Kanamatsu, T., Sato, T., Stein, R., Alvarez Zarikian, C.A., Malone, M.J.,Expedition 303/306 Scientists, 2006. Proceedings IODP, 303/306. College Station TX(Integrated Ocean Drilling Program Management International, Inc.).doi:10.2204/iodp.proc.303306.104.

PAGE 194

194Charles, C.D., Lynch-Stieglitz, J., Ninnemann, U.S., Fairbanks, R.G., 1996. Climate connectionsbetween the hemispheres revealed by deep sea sediment core/ice core correlations. EarthPlanet. Sci. Lett. 142, 19-28.Chough, S.K., Hesse, R., 1985. Contourites from Eirik Ridge, south of Greenland. SedimentaryGeology 41, 185-199.Clemens, S.C., 1999. An astronomical tuning stratigraphy for Pliocene sections: implications forglobal-scale correlation and phase relationship. In: Shackleton, N.J., McCave, I.N.,Weedon, G.P., (Eds.). Phil. Trans. R. Soc. Lond. A, 1949-1973.Clement, B. M., Robinson, F., 1986. The magnetostratigraphy of Leg 94 sediments, In:Ruddiman, W.F., Kidd, R. B., Thomas, E., Shipboard Scientists, Init. Repts. DSDP 94:Washington (U.S. Government Printing Office), 635-650Day, R., Fuller, M., Schmidt, V.A., 1977. Hysteresis properties of titanomagnetites: grain-sizeand compositional dependence. Phys. Earth Planet. Int. 13, 260-267.Evans, H.F., Channell, J.E.T., Sager, W.W., 2005. Late MioceneHolocene magnetic polaritystratigraphy and astrochronology, ODP Leg 198, Shatsky Rise. In: Bralower, T.J., PremoliSilva, I., Malone, M.J. (Eds.), Proceedings ODP, Sci. Results 198 [CD-ROM]. OceanDrilling Program. Texas A&M University College Station, TX, 1-39.Evans H.F., Westerhold, T., Channell, J.E.T., 2004. ODP Site 1092: revised composite depthsection has implications for Upper Miocene "cryptochrons". Geophys. J. Inter. 156, 195-199.Evans H.F., Channell, J.E.T., 2003. Upper Miocene magnetic stratigraphy at ODP Site 1092(sub-Antarctic South Atlantic): recognition of cryptochrons in C5n.2n. Geophys. J. Inter.153, 483-496.Evans, H.F., Channell, J.E.T., Stoner, J.S., Hillaire-Marcel, C., Wright, J.D., Neitzke, L.C.,Mountain G.S., submitted. Paleointensity-assisted chronostratigraphy of detrital layers onthe Eirik Drift (North Atlantic) since marine isotope stage 11. Geochem. Geophys.Geosyst. Submitted.Expedition 303 Scientists, 2006. Site U1304. In: Channell, J.E.T., Kanamatsu, T., Sato, T., Stein,R., Alvarez Zarikian, C.A., Malone, M.J., Expedition 303/306 Scientists. ProceedingsIODP, 303/306: College Station TX (IODP Management International, Inc.).doi:10.2204/iodp.proc.303306.104.Expedition 306 Scientists, 2006. Expedition 306 summary. In: Channell, J.E.T., Kanamatsu, T.,Sato, T., Stein, R., Alvarez Zarikian, C.A., Malone, M.J., Expedition 303/306 Scientists.Proceedings IODP, 303/306: College Station TX (IODP Management International, Inc.).doi:10.2204/iodp.proc.303306.109.

PAGE 195

195Funder, S., Hjory, C., Landvik, J. Y., Nam, S-I., Reeh, N., Stein, R., 1998. History of a stable icemarginEast Greenland during the Middle and Upper Pleistocene. Quat. Sci. Rev. 17, 77-123.Gradstein, F., Ogg, J., Smith A., 2005. A Geologic time scale 2004. Cambridge, CambridgeUniversity Press, pp. 589.Grootes, P.M., Stuvier, M., 1997. Oxygen 18/16 variability in Greenland snow and ice with 1033to 105-year time resolution. J. Geophys. Res. 102, 26,455-26,470.Gubbins, D., 1999. The distinction between geomagnetic excursions and reversals. Geophys. J.Inter. 137, F1-F3.Guyodo, Y., Channell, J.E.T., Thomas, R.G., 2003. Deconvolution of u-channel paleomagneticdata near geomagnetic reversals and short events. Geophys. Res. Letters 29, 1845,doi:10.1029/2002GL014963.Hagelberg. T.K., Pisias, N.G., Shackleton N.J., Mix, A.C., Harris S., 1995. Refinement of a high-resolution, continuous sedimentary section for studying equatorial Pacific Oceanpaleoceanography, Leg 138. In: Pisias, N.G., Mayer, L.A., Janecek, T.R. Palmer-Julson,A., van Andel, T.H., (Eds.). Proceedings ODP Sci., Res. 138. Ocean Drilling Program,College Station, TX, 31-46.Harland, W.B., Cox, A.V. Llewellyn, P.G., Pickton, C.A.G., Smith, A.G., Walters, R., 1982. AGeologic Time Scale. Cambridge Univ. Press, Cambridge.Harland, W.B., Armstrong, R., Cox, A.V., Craig, L., Smith, A., Smith, D., 1990. A GeologicTime Scale 1989. Cambridge Univ. Press, Cambridge.Hemming, S., 2004. Heinrich Events: Massive Late Pleistocene detritus layers of the NorthAtlantic and their global imprint. Rev. Geophys. 42, RG1005,doi:10.1029/2003RG000128.Heirtzler, J.R., Dickson, G.O., Herron, E.M., Pittman, W.C., III, LePichon, X., 1968. Marinemagnetic anomalies, geomagnetic field reversal and motions of the ocean floor andcontinents. J. Geophys. Res. 73, 2119-2136.Hilgen, F.J., 1991a. Astronomical calibration of Gauss to Matuyama sapropels in theMediterranean and implication for the Geomagnetic Polarity Time Scale. Earth Planet. Sci.Lett. 104, 226-244.Hilgen, F.J., 1991b. Extension of the astronomically calibrated (polarity) time scale toMiocene/Pliocene boundary. Earth Planet. Sci. Lett. 107, 349-368.Hilgen, F.J., Krijgsman, W., Langereis, C.G., Lourens, L.J., Santarelli, A., Zachariasse, W.J.,1995. Extending the astronomical (polarity) time scale into the Miocene. Earth Planet. Sci.Lett. 136, 495-510.

PAGE 196

196Hilgen, F.J., Krijgsman, W., Raffi, I., Turco, E., Zachariasse, W.J., 2000. Integrated stratigraphyand astronomical calibration of the Serravallian/Tortonian boundary section at MonteGibliscemi (Sicily, Italy). Marine Micropaleontology 38, 181-211.Hilgen, F.J., Abdul Aziz, H., Krijgsman, W., Raffi, I., Turco, E., 2003. Integrated stratigrpahyand astronomical tuning of the Serravallian and lower Tortonian at Monti dei Corvi(Middle-Upper Miocene, northern Italy). Palaeogeography, Palaeoclimatology,Palaeoecology 199, 229-264.Hillaire-Marcel, C., Bilodeau, G., 2000. Instabilities in the Labrador Sea water mass structureduring the last climatic cycle. Can. J. Earth Sci. 37, 795-809.Hillaire-Marcel, C., De Vernal, A. Bilodeau, G., Wu, G., 1994. Isotope stratigraphy,sedimentation rates, deep circulation, and carbonate events in the Labrador Sea during thelast ~200 ka. Can. J. Earth Sci. 31, 63-89.Hiscott, R.N., Aksu, A.E., Mudie, P.J., Parsons, D.F., 2001. A 340,000 year record of ice rafting,paleoclimatic fluctuations, and shelf crossing glacial advances in the southwesternLabrador Sea. Global and Planetary Change 28, 227-240.Hodell, D.A., Charles, C.D., Sierro, F.J., 2001. Late Pleistocene evolution of the ocean'scarbonate system. Earth Planet. Sci. Lett. 192, 109.Hodell, D.A., Kennett, JP., 1986. Late Miocene-early Pliocene stratigraphy andpaleoceanography of the South Atlantic and southwest Pacific Oceans: A synthesis.Paleoceanography 1, 285-311.Hollerbach, R., Jones, C.A., 1995. On the magnetically stabilizing role of the Earths inner core,Phys. Earth Planet. Inter. 87, 171-181.Iaccarino, S., 1985. Mediterranean Miocene and Pliocene planktic foraminifera. In: Bolli, H.M.,Saunders, J.B., Perch-Nielsen, K. (Eds.). Plankton Stratigraphy, Cambridge, CambridgeUniversity Press, pp. 283-314.Jenkins, D.G., 1985. Southern and mid-latitude Paleocene to Holocene planktic foraminifera. In:Bolli, H.M., Saunders, J.B., Perch-Nielsen, K. (Eds.). Plankton Stratigraphy, Cambridge,Cambridge University Press, pp. 263-282.Jouzel, J., Lorius, C., Petit, J.R., Genthon, C., Barkov, N.I., Kotlyakov, V.M., Petrov, V.M.,1987. Vostok ice core: a continuous isotope temperature record over the last climatic cycle(160 000 years). Nature 329, 402-408.Kawase, M., Sarmiento, J.L., 1986. Circulation and nutrients in middepth Atlantic waters, J.,Geophys. Res. 91, 9748-9770.Keller, G., 1979a, Late Neogene planktonic foraminiferal biostratigraphy and paleoceanographyof the northwest Pacific DSDP Site 296. Palaeogeography Palaeoclimatology,Palaeoecology 27, 129-154.

PAGE 197

197Keller, G., 1979b. Late Neogene paleoceanography of the North Pacific DSDP Sites 173, 310and 296. Marine Micropaleontology 4, 159-172.Keller, G., 1979c. Early Pliocene to Pleistocene planktonic foraminiferal datum levels in theNorth Pacific: DSDP Sites 173, 310, 296. Marine Micropaleontology 4, 281-294.Kennett, J.P., Srinivasan, M., 1983. Neogene Planktonic Foraminifera: A phylogenetic atlas:Stroudsberg, Hutchinson Ross Publishing Co. pp. 265.King, J.W., Banerjee, S.K., Marvin, J., 1983. A new rock-magnetic approach to selectingsediments for geomagnetic paleointensity studies: application to paleointensity for the last4000 years. J. Geophys. Res. 88, 5911-5921.Kirschvink, J.L., 1980. The least squares lines and plane analysis of palaeomagnetic data.Geophys. J. R. Astr. Soc. 62, 699-718.Kissel, C., Laj, C., Labeyrie, L., Dokken, T., Voelker, A., Blamart, D., 1999. Rapid climaticvariations during marine isotope stage 3: magnetic analysis of sediments from the NordicSeas and North Atlantic. Earth Planet. Sci. Lett. 171, 489-502.Kok, Y.S., Tauxe, L., 1996a. Saw-toothed pattern of relative paleointensity records andcumulative viscous remanence. Earth Planet. Sci, Lett. 137, 95-99.Kok, Y.S., Tauxe, L., 1996b. Saw-toothed pattern of sedimentary paleointensity recordsexplained by cumulative viscous remanence. Earth Planet. Sci. Lett. 144, 9-14.Krijgsman W., Kent, D.V. 2004. Non-uniform occurrences of short-term polarity fluctuations inthe geomagnetic field? New results from Middle to Late Miocene sediments from theNorth Atlantic. In: Timescales of the paleomagnetic field. Channell, J.E.T., Kent, D.V.,Lowrie, W., Meert, J. G. (Eds.). AGU Geophysical Monograph 145, 161-174.Laj, C., Kissel, C., Beer, J., 2004. High resolution global paleointensity stack since 75 kyr(GLOPIS-75) Calibrated to absolute values. In: Timescales of the paleomagnetic field.Channell, J.E.T., Kent, D.V., Lowrie, W., Meert, J. G. (Eds.). AGU GeophysicalMonograph 145, 255-265.Laj, C., Kissel, C., Mazaud, A., Channell, J.E.T., Beer, J., 2000. North Atlantic paleointensitystack since 75 ka (NAPIS-75) and the duration of the Laschamp event. Phil. Trans. Roy.Soc. 358, 1009-1025.Laskar, J., Joutel, H., Boudin, F., 1993. Orbital, precessional and insolation quantities for theEarth from Myr to +10 Myr. Astron. Astrophys. 270, 522-533.Laskar, J., Robutel, P., Joutel, F., Gastineau, M., Correia, A.C.M., Levrard, B., 2004. A longterm numerical solution for the insolation quantities of the Earth. Astron. Astrophys. 428,261-285.

PAGE 198

198Lisiecki L., Raymo, M., 2005. A Pliocene-Pleistocene stack of 57 globally distributed benthic"18O records. Paleoceanography 20, doi:10.1029/2004PA001071.Lourens, L.J., Hilgen, F.J., Laskar, J., Shackleton, N.J., Wilson D., 2004. The Neogene Period.In: Gradstein, F.M., J.G. Ogg, A.G. Smith, (Eds.). Geologic Time Scale 2004. CambridgeUniv. Press, pp. 409-440.Lucotte, M., Hillaire-Marcel, C., 1994. Identification et distribution des grandes masses d'eaudans les mers du Labrador et d'Irminger. Can. J. Earth Sci. 31, 5-13.Lund, S.P., Acton, G., Clement, B., Hastedt, M., Okada, M., Williams, T., Shipboard ScientificParty, 1998. Geomagnetic field excursions occurred often during the last million years.Trans. Am. Geophys. Union (EOS) 79, 178.Maenaka, K., 1983. Magnetostratigraphic study of the Osaka Group, with special reference to theexistance of pre and post-Jaramillo episodes in the Late Matuyama polarity epoch. Mem.Hanazono Univ. 14, 1-65.Martinson, D.G., Pisias, N. G., Hays, J.D., Imbrie, J., Moore, T.C., Shackleton, N.J., 1987. Agedating and the orbital theory of the Ice Ages: development of a high-resolution 0 to300,000-year chronostratigraphy. Quat. Res. 27, 1-29.Mazaud, A., 1996. 'Sawtooth' variation in magnetic intensity profiles and delayed acquisition ofmagnetization in deep sea cores. Earth Planet. Sci. Lett. 139, 379-386.McCartney, M.S., 1992. Recirculating components to the deep boundary current of the northernNorth Atlantic. Prog. Oceanog. 29, 283-383.McCave, I.N., Tucholke, B.E., 1986. Deep current-controlled sedimentation in the western NorthAtlantic. In: The geology of North America: The Western Atlantic region. Vogt, P.R.,Tucholke, B.E., (Eds.). Geol. Soc. Am., DNAG Ser., Boulder, CO., Vol., M Spec. Publ.,pp. 451-468.McCave, I. N., Manighetti, B., Robinson, S.G., 1995. Sortable silt and fine sediment slicing:Parameters for paleocurrent speed and paleoceanography. Paleoceanography 10, 593-610.Meese, D.A., Gow, A. J., Alley, R.B., Zielinski, G.A., Grootes, P.M., Ram, M., Taylor, K.C.,Mayewski, P.A., Bolzan, J.F., 1997. The Greenland Ice Sheet Project 2 depth-age scale:Methods and results. J. Geophys. Res. 102, 26,411-26,423.Meynadier, L., Valet, J-P., Bassinot, F., Shackleton, N.J., Guyodo, Y., 1994. Asymmetrical saw-tooth pattern of the geomagnetic field intensity from equatorial sediments in the Pacificand Indian Oceans. Earth Planet. Sci. Lett. 126, 109-127.Miller, K.G., Tucholke, B.E., 1983. Development of Cenozoic abyssal circulation south of theGreenlandScotland Ridge. In: Bott, M.H.P. (Ed.), Structure and Development of theGreenlandScotland Ridge: New Methods and Concepts. Plenum Press, New York, pp.549.

PAGE 199

199Moreno, E., Thouveny, N., Delanghe, D., McCave, I.N., Shackleton, N.J., 2002. Climatic andoceanographic changes in the Northeast Atlantic reflected by magnetic properties ofsediments deposited on the Portuguese Margin during the last 340 ka. Earth Planet. SciLett. 202, 465-480.Muscheler, R., Beer, J., Kubik, P.W., Synal, H.-A., 2005. Geomagnetic field intensity during thelast 60,000 years based on 10Be and 36Cl from the Summit ice cores and 14C. Quat. Sci.Reviews 24, 1849-1860.Ness, G., Levi, S., Couch, R., 1980. Marine magnetic anomaly timescales for the Cenozoic andLate Cretaceous: A precis, critique and synthesis. Reviews of Geophysics and SpacePhysics 18, 4, 753-770.Okada, H., Bukry, D., 1980. Supplementary modification and introduction of code numbers tothe low-latitude coccolith biostratigraphic zonation (Bukry, 1973; 1975). MarineMicropaleontology 5, 321.Opdyke, N.D., Glass, B., Hays, J.P., Foster, J., 1966. Paleomagnetic study of Antarctica deep-seacores. Science 154, 349-357.Opdyke, N.D., Channell, J.E.T., 1996. Magnetic stratigraphy. Academic Press, San Diego, Calif.,346 pp.Paillard, D., Labeyrie, L., Yiou, P., 1996. Macintosh program performs time-series analysis,Trans. Am. Geophys. Union, EOS 77, 379.Paulsen, H., Westerhold, T., Bickert, T., in press. Middle to Late Miocene oxygen isotopestratigraphy of the Southern Ocean. Geology, in press.Raffi, I., Flores, J-A., 1995. Pleistocene through Miocene calcareous nannofossils from easternequatorial Pacific Ocean (ODP Leg 138). In: Pisias, N.G., Mayer, L.A., Janecek, T.R.Palmer-Julson, A., van Andel, T.H., (Eds.). Proceedings ODP, Sci., Res. 138. OceanDrilling Program, College Station, TX, 59-72.Raymo, M.E., Ruddiman, W.F., Backman, J., Clement, B.M., Martinson, D.G., 1989. LatePliocene variation in Northern Hemisphere ice sheets and North Atlantic Deep Watercirculation. Paleoceanography 4, 413.Rea, D.K., Basov, I.A., Janecek, T.R., Palmer-Julson, A., Shipboard Scientific Party 1993.Proceedings ODP, Init. Repts., 145, College Station, TX, ODP.Rio, D., Raffi, I., Villa, G., 1990. Pliocene-Pleistocene calcareous nannofossil distributionpatterns in the western Mediterranean. In: Kastens, K.A., Mascle, J. Proceedings of theODP, Sci. Res., Leg 107. Ocean Drilling Program, College Station, TX, ODP, 513-533.Roberts, A.P., Lewin-Harris, J.C., 2000. Marine magnetic anomalies: evidence that tinywiggles represent short-period geomagnetic polarity intervals. Earth Planet. Sci. Lett. 183,375.

PAGE 200

200Rhl, U., Abrams, L.J., 2000. High-resolution, downhole, and non-destructive coremeasurements from Sites 999 and 1001 in the Caribbean Sea: application to the LatePaleocene Thermal Maximum. In: Leckie, R.M., Sigurdsson, H., Acton, G.D., Draper G.(Eds.). Proceedings ODP, Sci. Res. 165. Ocean Drilling Program, College Station, TX,191-203.Ruddiman, W.F., Raymo, M., McIntyre, A., 1986. Matuyama 41,000-year cycles: North AtlanticOcean and northern hemisphere ice sheets. Earth Planet. Sci. Lett. 80:117.Ruddiman, W.F., Kidd, R.B., Thomas, E., Shipboard Scientific Party 1987. Init. Repts. DSDP,94 (Pts. 1 and 2), Washington (U.S. Govt. Printing Office).Schneider, D.A., 1995. Paleomagnetism of some Leg 138 sediments: detailing Miocenemagnetostratigraphy, In: Pisias, N.G., Mayer, L.A., Janecek, T.R. Palmer-Julson, A., vanAndel, T.H., (Eds.). Proceedings ODP Sci., Res. 138. Ocean Drilling Program, CollegeStation, TX, 59-72.Shackleton, N.J., Fairbanks, R.G., Chiu, T.-C., Parrenin, F., 2004. Absolute calibration of theGreenland time scale: implications for Antarctic time scales and for #14C. Quat. Sci.Reviews 23, 1513-1522.Shackleton, N.J., Crowhurst, S., 1997. Sediment fluxes based on an orbitally tuned time scale 5Ma to 14 Ma, Site 926. In: Shackleton, N.J., Curry, W.B., Richter, C., Bralower, T.J.(Eds.). Proceedings ODP Sci., Res. 154. Ocean Drilling Program. College Station, TX 69-82.Shackleton, N.J., Crowhurst, S., Hagelberg., T., Pisias, N.G., Schneider, D.A., 1995. A new LateNeogene time scale: application to Leg 138 sediments. In: Pisias, N.G., Mayer, L.A.,Janecek, T.R. Palmer-Julson, A., van Andel, T.H., (Eds.). Proceedings ODP Sci., Res. 138.Ocean Drilling Program, College Station, TX, 73-101.Shackleton, N.J., Baldauf, J.G., Flores, J-A., Iwai, M., Moore, T.C., Raffi, I., Vincent, E., 1995a.Biostratigraphic summary for Leg 138. In: Pisias, N.G., Mayer, L.A., Janecek, T.R.Palmer-Julson, A., van Andel, T.H., (Eds.). Proceedings ODP Sci., Res. 138. OceanDrilling Program, College Station, TX, 73-101.Shackleton, N.J., Berger, A., Peltier, W.R., 1990. An alternative astronomical calibration of thelower Pleistocene timescale based on ODP site 677. Trans. R. Soc. Edinburgh 81, 251-261.Shackleton, N. J., Hall, M. A., Boersma, A., 1984. Oxygen and carbon isotope data from Leg 74foraminifers, Init. Rep. DSDP 74, Washington (U.S. Govt. Printing Office), 599.Shipboard Scientific Party, 1999. Site 1092. In: Gersonde, R., Hodell, D.A., Blum, P., ShipboardScientific Party. Proceedings ODP Init. Repts. 177, Ocean Drilling Program, CollegeStation, TX, 1-82.

PAGE 201

201Shipboard Scientific Party, 2002a. Leg 198 Summary. In: Bralower, T.J. Premoli-Silva I.,Malone M., Shipboard Scientific Party. Proceedings ODP, Init. Repts. 198 [CD-ROM],College Station, TX (Ocean Drilling Program) 1-148.Shipboard Scientific Party, 2002b. Site 1208. In Bralower, T.J., Premoli Silva, I., Malone, M.J.,Shipboard Scientific Party. Proc. ODP, Init. Repts., 198: College Station, TX (OceanDrilling Program), 1. doi:10.2973/odp.proc.ir.198.104.2002.Singer. B.S., Hoffman, K.A., Chauvin, A., Coe, R.S., Pringle, M.S., 1999. Dating transitionallymagnetized lavas of the late Matuyama chron: Toward a new 40Ar/39Ar timescale ofreversals and events. J. Geophys. Res. 104, 679-693.Snowball, I., Sandgren, P., 2004. Geomagnetic field intensity changes in Sweden between 9000and 450 cal BP: extending the record of archaeomagnetic jerks by means of lake sedimentsand the pseudo-Thellier technique. Earth Planet. Sci Lett. 227, 361-376.Snowball, I., Moros, M., 2003. Saw-tooth pattern of North Atlantic current speed duringDansgaard-Oeschger cycles revealed by the magnetic grain size of Reykjanes Ridgesediments at 59 N. Paleoceanography 18, doi:10.1029/2001PA000732.Srinivasan, M.S., Sinha, D.K., 1993. Late Neogene planktonic foraminiferal events of thesouthwest Pacific and Indian Ocean: A comparison. In: Tsuchi, R., Ingle, J.C., Eds. PacificNeogene environment, evolution and events: Tokyo, University of Tokyo Press, pp. 203-220.Srivastava, S.P., Tapscott, C.R., 1986. Plate kinematics of the North Atlantic. In: The Geology ofNorth America: The Western Atlantic Region. Vogt, P.R., Tucholke, B.E., (Eds.), Geol.Soc. Am., DNAG Ser., Boulder, CO., Vol., M, Spec. Publ., p. 589-604.Stoner, J.S., Channell, J.E.T., Hodell, D.A., Charles, C., 2003. A 580 kyr paleomagnetic recordfrom the sub-Antarctic South Atlantic (ODP Site 1089). J. Geophys. Res. 108,doi:10.1029/2001JB001390.Stoner, J.S., Laj, C., Channell, J.E.T., Kissel, C., 2000. South Atlantic (SAPIS) and NorthAtlantic (NAPIS) geomagnetic paleointensity stacks (0-80 ka): implications for inter-hemispheric correlation. Quat. Sci. Reviews 21, 1141-1151.Stoner, J.S., Channell, J.E.T., Hillaire-Marcel, C., 1996. The magnetic signature of rapidlydeposited detrital layers from the deep Labrador Sea: Relationship to North AtlanticHeinrich layers. Paleoceanography 11, 309-325.Stoner, J.S., Channell, J.E.T., Hillaire-Marcel, C., 1995a. Late Pleistocene relative geomagneticpaleointensity from the deep Labrador Sea: Regional and global correlations. Earth Planet.Sci. Lett. 134, 237-252.Stoner, J.S., Channell, J.E.T., Hillaire-Marcel, C., 1995b. Magnetic properties of deep-seasediments off southwest Greenland: Evidence for major differences between the last twodeglaciations. Geology 23, 241-244.

PAGE 202

202Tauxe, L., 1993. Sedimentary records of relative paleointensity of the geomagnetic field: Theoryand practice. Rev. Geophys. 31, 319-354.Tauxe, L., Hartl, P., 1997. 11 million years of Oligocene geomagnetic field behaviour. Geophys,J. Int. 128, 217-229.Tauxe, L., Pick, T., Kok, Y.S., 1995. Relative paleointensity in sediments: a pseudo-Thellierapproach. Geophys. Res. Lett. 22, 2885-2888.Tauxe, L., Shackleton, N.J., 1994. Relative palointensity records from the Ontong-Java Plateau.Geophys. J. Int. 117, 769-782.Thibal, J., J. P. Pozzi, V. Barthe`s, Dubuisson, G., 1995. Continuous record of geomagnetic fieldintensity between 4.7 and 2.7 Ma from downhole measurements. Earth. Planet. Sci. Lett.136, 541.Thomas, R., Guyodo, Y., Channell, J.E.T., 2004. U-channel track for susceptibilitymeasurements. Geochem. Geophys. Geosyst. 1050, doi: 10.1029/2002GC000454Tian, J., Wang, P., Cheng, X., Li, Q., 2002. Astronomically tuned Plio-Pleistocene benthic "18Orecord from South China Sea and Atlantic-Pacific comparison. Earth Planet. Sci. Lett. 203,1015-1029.Turon, J.-L., Hillaire-Marcel, C., Shipboard Participants, 1999. IMAGES V mission of theMarion Dufresne. Leg 2, 30 June to 24 July 1999. Geol. Surv. Canada, Open File 3782.Valet, J-P. Time variations in geomagnetic intensity, 2003. Rev. Geophys. 41,doi:10.1029/2001RG000104.Valet, J-P., Meynadier, L., 1993. Geomagnetic field intensity and reversals during the past fourmillion years. Nature 366, 234-238.Van Kreveld, S.A., Knappertsbusch, M., Ottens, J., Ganssen, G., van Hinte, J., 1996. Biogeniccarbonate and ice-rafted debris (Heinrich layer) accumulation in deep-sea sediments from aNortheast Atlantic piston core. Mar. Geol. 131, 21-46.Venti, N.J., 2006. Revised Late Neogene mid-latitude planktic foraminiferal biostratigraphy forthe Northwest Pacific (Shatsky Rise), ODP Leg 198. MS Thesis, University ofMassachusetts.Venz, K.A., Hodell, D.A., Stanton, C., Warnke, D.A, 1999. A 1.0 Myr record of Glacial NorthAtlantic Intermediate Water Variability from ODP Site 982 in the northeast Atlantic.Paleoceanography 14, 42-52.Voelker, A., Sarnthein, M., Grootes, P. M., Erlenkeuser, H., Laj, C., Mazaud, A., Nadeau, M.J.,Schleicher, M., 1998. Correlation of marine 14C ages from the Nordic sea with GISP2isotope record: implication for 14C calibration beyond 25 ka BP. Radiocarbon 40, 517-534.

PAGE 203

203Westerhold T., Roehl, U., Raffi, I., Bowles, J., Evans, H.F., in preparation. The first completeorbital chronology for the Paleocene: Implications for the Geomagnetic Polarity TimeScale and the age of the K-T boundary.Yamazaki, T., Oda, H., 2005. A geomagnetic paleointensity stack between 0.8 and 3.0 Ma fromequatorial Pacfic sediment cores. Geochem. Geophys. Geosys. 11,doi:10.1029/2005GC001001.Yang, S., Odah, H., Shaw, J., 2000. Variations in the geomagnetic dipole moment over the last12,000 years. Geophys. J. Int. 140, 158-162.Young, J.R., 1998. Neogene. In Bown, P.R. (Ed.). Calcareous Nannofossil Biostratigraphy.Dordrecht, The Netherlands (Kluwer Academic Publ.), pp. 225.

PAGE 204

204BIOGRAPHICAL SKETCHHelen F. Evans was born in Swansea, South Wales, in 1977 to Terry and Eryl Evans.Growing up in an area of outstanding natural beauty she gained an interest at a young age ingeology and natural sciences. She earned three A-levels in geology, chemistry and biology fromGowerton Comprehensive School in 1995. The same year she began an undergraduate career atImperial College of Science, Technology and Medicine, in London. She graduated with honorsin 1998 with a BSc. in geology with paleontology and Associateship of the Royal School ofMines (ARSM). After taking a year out she moved to the University of Florida to continue hereducation in geology. She gained a MS in Geology from the University of Florida in 2001 with athesis entitled "Late Miocene to Pleistocene Magnetic Stratigraphy at ODP Site 1092(subantarctic South Atlantic)". During her stay at the University of Florida, Helen was awarded aMcLaughlin Dissertation Fellowship, the University Womens Club Graduate StudentScholarship and an outstanding academic achievement award from the College of Liberal Artsand Sciences. She also spent seven weeks at sea aboard the JOIDES Resolution during a researchcruise to the North Atlantic. Helen is a member of the American Geophysical Union and theGeological Society of America. She has been an author on sixteen abstracts presented atinternational meetings and six peer-reviewed journal articles.