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1 DEEP OCEAN CIRCUL A TION AND CONTINENTAL WEATHERING REGIMES DURING CLIMATE TRANSITIONS (LAST DEGLACIATION AND EOCENEOLIGOCENE) USING Sr, Nd, AND Pb ISOTOPES IN SEDIMENTARY ARCHIVES By CHANDRANATH BASAK 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 2011
2 2011 Chandranath Basak
3 To my mentors
4 ACKNOWLEDGMENTS I would like to take this opportunity to thank a number of people whose direct and indirect support have helped me through y ears of graduate school. First, I thank my parents f or their unconditional support towards my work and me. I am extremely thankful to my advisor Dr. Ellen E. Martin for her guidance and support throughout the Ph.D. program. Without Ellens help and advisement this work would have never been successfully completed. I would also like to thank my entire committee (Dr. Paul A. Mueller, Dr. George Kamenov, Dr. Andrea Dutton, and Dr. John Krigbaum) for their valuable advic e from time to time. Special thanks to Dr. George Kamenov for his patience while teaching the theory and instrumentation of m ass spectrometer I am grateful to my masters adviser Dr. Anthony E. Rathburn for introducing me to the field of paleoclimate research. I am also thankful to Dr. David Hodell Dr. Thomas M. Marchitto, and Dr. Paul A. Mueller for scientific discussions, which helped shape some of the research ideas presented i n this dissertation. I am very thankful for all the help and laughter that Derrick Newkirk has extended over the years. Long lab hours would have been quiet boring without Derricks presence. I sincerely thank one of our alumni, Ms. Elodie Bourbon, and sev eral undergraduate student researchers for their help with laboratory work from time to time. Continuous technical support from Mr. Ray G. Thomas and Mr. Dow Van Arnam had been very critical to continue day to day lab activity. Numerous discussion sessions with Dr. Keiji Horiwaka on myriad topics of paleoceanography often helped me evaluate old data with a new er perspective. I really appreciate all the help that I have received over the years from the entire faculty and the graduate students body with in the department. Special thanks to
5 Nita, Pam, and Susan for all their help with paper work I am thankful for the financial support extended by the National Science Foundation, the American Geochemical Society, the Geological Society of America, and the Depart ment of Geological Sciences, University of Florida. Lastly, I would like to thank my wife and dear friend, Dr. Pinki Mondal whose constant enthusiasm about my research often times, kept me goi ng when I thought there were not much to accomplish
6 TABLE OF CONTENTS page ACKNOWLEDGMENTS .................................................................................................. 4 LIST OF TABLES ............................................................................................................ 8 LIST OF FIGURES .......................................................................................................... 9 ABSTRACT ................................................................................................................... 10 CHAPTER 1 INTRODUCTION .................................................................................................... 12 Cenozoic C limatic Events ....................................................................................... 12 Research Objectives ............................................................................................... 14 2 SEAWATER Pb ISOTOPES EXTRACTED FROM CENOZOIC MARINE SEDIMENTS ........................................................................................................... 17 Overview ................................................................................................................. 17 Methods .................................................................................................................. 19 Sample Preparation .......................................................................................... 19 Column Chemistry and Isotope Analyses ......................................................... 21 Major Elements and REE Analyses .................................................................. 23 Results .................................................................................................................... 24 Pb Isotopes in Sequentially Extracted Fractions .............................................. 24 Pb Isotopes Compositions in Fossil Fish Teeth ................................................ 24 Major Elements in FeMn Extracted Fraction ................................................... 25 REE in HH Extractions ..................................................................................... 26 Discussion .............................................................................................................. 26 Integrity of Pb Isotopes from HH Extractions .................................................... 26 Fossil Fish Teeth as a Pb Isotope Archive ....................................................... 31 Phases Contributing Pb to the HH Extraction: Major Elements ........................ 33 Phases contributing Pb to the HH extraction : REEs ........................................ 35 Summary ................................................................................................................ 38 3 WEATHERING HISTORY OF ANTARCTICA DURING THE EOCENEOLIGOCENE TRANSITION .................................................................................... 55 Background ............................................................................................................. 57 Sm Nd Systematics .......................................................................................... 57 U ThPb Systematics ....................................................................................... 58 Methods .................................................................................................................. 59 Results .................................................................................................................... 61
7 Pre EOT (34 34.5 Ma) ................................................................................... 61 EOT (33.7 34 Ma) .......................................................................................... 62 Post EOT (33.3 33.7 Ma) ............................................................................... 62 Discussion .............................................................................................................. 63 Source of Radiogenic Pb During Weathering ................................................... 63 Weathering at the EOT ..................................................................................... 67 Mechanical vs. chemical weathering .......................................................... 67 Silicate vs. carbonate weathering .............................................................. 68 Residue Pb Isotopes ........................................................................................ 70 Weathering and Paleoceanogrpahic Significance ............................................ 71 Summary ................................................................................................................ 72 4 SOUTHERN SOURCE OF DEG LACIAL OLD CARBON IN THE TROPICAL NORTH PACIFIC BASED ON NEODYMIUM ISOTOPES ...................................... 80 Background ............................................................................................................. 81 Nd Isotopes Systematics .................................................................................. 81 Hydrography ..................................................................................................... 82 Core Location and Method ...................................................................................... 84 Results .................................................................................................................... 85 Discussion .............................................................................................................. 86 Summary ................................................................................................................ 90 5 CONCLUSIONS ..................................................................................................... 95 LIST OF REFERENCES ............................................................................................... 98 BIOGRAPHICAL SKETCH .......................................................................................... 113
8 LIST OF TABLES Table page 2 1 Pb isotopes in different bulk sediment fractions from ODP site 1090 ................. 41 2 2 Pb isotopes measured in fossil fish teeth ........................................................... 42 2 3 Major elements in different bulk sediment fractions from ODP Site 1090. .......... 43 2 4 REE data in different bulk sediment fractions from ODP site 1090. .................... 45 2 5 Inter and intra laboratory Pb isotope duplicates of authigenic FeMn phases. ... 47 2 7 Measured Pb, U, and Th concentrations in fossil and modern fish teeth. ........... 49 3 1 Average Pb isotopic values in seawater and residues in PreEOT, EOT, and Post EOT intervals in Southern Ocean sites 689 and 738. ................................ 74 4 1 Nd isotopes from fossil fish teeth/derbis of core MV99MC19/GC31/PC08 ........ 91
9 LIST OF FIGURES Figure page 2 1 Pb isotopes measured in the different fractions of leaching experiment ............ 50 2 2 Pb isotopes in fossil fish teeth. ........................................................................... 51 2 3 Plots showing relationship of measured major elemental concentrations with respect to Pb concentrations .............................................................................. 52 2 4 Plots showing REE behavior of the extracted fractions. .................................... 53 2 5 Compilation of plots showing the reproducibility and integrity of the extraction procedure. .......................................................................................................... 54 3 1 Site location showing ODP site 689 (2080 m) from the Atlantic sector and site 738 (2253 m) from the Indian sector of the Southern Ocean. ............................. 75 3 2 A plot of 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios versus age in seawater and detrital residues from site 689 and site 738. ................................ 76 3 3 Nd values versus age of fossil fish teeth, representing seawater, and detrital fractions for 33.3 34.5 Ma. ............................................................................... 77 3 4 A plot of 206Pb/204Pb vs. 207Pb/204Pb from a leaching experiment .............. 77 3 5 ..................................... 78 3 6 Plot showing the different proxy response to Oligocene glaciation in Antarctica ~ 34 Ma. ............................................................................................ 79 4 1 Mo dern hydrography .......................................................................................... 92 4 2 ........................................ 93 4 3 Meridional multi proxy responses to the last deglaciation (8 to 22kyr BP) ......... 94
10 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 DEEP OCEAN CIRCUL A TION AND CONTINENTAL WEATHERING REGIMES DURING CLIMATE TRANSITIONS (LAST DEGLACIATION AND EOCENEOLIGOCENE) USING Sr, Nd, AND Pb ISOTOPES IN SEDIMENTARY ARCHIVES By Chandranath Basak May 2011 Chair: Ellen E. Martin Major: Geology The global climate on earth is constantly changing and periodically the earth system crosses a climate threshold and enters a new climate state. The EoceneOligocene transition (EOT) ~33.4 Ma ago and the last deglaciation ~18 k a are two examples of major climate transitions. During the EOT the Antarctic cryos phere developed, there was global cooling of deep waters, and the carbonat e compensation depth deepened. During the last deglaciation there was massive global ice meltback, a dramatic rise in atmospheric CO2, and a decline in atmospheric 14C. These two ev ents are climatologically i mportant since the former represents the first Cenozoic transition from a greenhouse to an icehouse world, while the later represents the end of the last glacial period. M ajor climate change is frequently associated with readjustment of carbon sources and sink s. The ocean is a relatively large carbon sink that can exchange carbon with the at mosphere over geologically short time scales. Chemical weathering of continental silicate rocks also plays a key role in sequestering atmospheric carbon over geologically long time scale. Therefore, the interplay between physical and chemical
11 continental weathering can have major implications for the long term carbon budget of the Earth. In this dissertation I studied continental weathering regimes on Antarctica during the EOT and the role of carbon storage and circulation in the ocean during the last deglacial in order to improve our understanding of the relationship between climat e, ocean circulation, and continental weathering. Radiogenic isotopes of Pb, and Nd were extracted from sedimentary archives such as FeMn oxide coatings and fossil fish teeth to acquire meaningful information pert aining to weathering and ocean circulation readjustments. Pb isotopic data from this study confir m a change in the Antarctic weathering regime leading up to the EOT, which may have contributed to the drawdown of atmospheric CO2 and ice buildup on Antarctica. I also documented a r earrangement of intermediate water circulation in the Pacific using Nd isotopic data recovered from fossil fish teeth. The old carbon that was trapped in the Southern Ocean abyss by sea ice or ocean stratification during the last glaciations was transported by intermediate water s and ventilated to the atmosphere following mel tback of the ice at the end of the last glaciation.
12 CHAPTER 1 INTRODUCTION Cenozoic Climatic Events The EoceneOligocene (E/O) boundary and the last deglaciation are two important climate threshold events; the first was characterized by massive ice growth and the second by ice decay. While there are a lot of differences between these events, both are marked by dramatic changes in atmospher ic CO2 and temperature. Prior to the E/O boundary, the climate of the Paleocene and Eocene was dominated by warm global temperatures (Zachos et al., 2001) punctuated by episodes of short lived, rapid global warming known as hyperthermals (Sexton et al., 2011). Cenozoic cooling started following the midEocene climatic optima with a major step of cooling and ice growth at the E/O boundary ~34 Ma ago (Miller et al., 1991; Zachos et al., 2001). This boundary is marked by development of an Antarctic cryospher e with a volume similar to todays volume (Pekar and DeConto, 2006; Pekar et al., 2006) and is often identified as the onset of the Cenozoic icehouse. The E/O boundary is also characterized by a general decreasing trend in atmospheric CO2 (Pagani et al., 2005; Zachos et al., 2008; Pearson et al., 2009), a positive shift in benthic foraminifera 18O and 13C (Coxall et al., 2005), ~ 1 km deepening of the carbonate compensation depth (CCD) (Coxall et al., 2005), occurrences of ice rafted debris (IRD) in the Southern Oceans (Zachos et al., 1992a), and a shift in clay mineral assemblages from smectite to illite and chlorite in the Southern Ocean (Robert and Kennett, 1997; Robert et al. 2002). The development of icehouse conditions has been attributed to thermal isolation of Antarctica due to tectonic processes and opening of the Drake and Tasman gateways (Kennett, 1977; Exon, 2000), decreasing atmospheric greenhouse gas concentrations (mainly CO2) below a
13 threshold value (DeConto and Pollard, 2003), and orbitall ypaced climate cycles (Palike et al., 2006). While all of these mechanisms may have contributed to Antarctic glaciation, the decreasing trend of atmospheric CO2 has been identified as a critical factor in the observed continental ice growth (DeConto and Pollard, 2003). The large decrease in atmospheric CO2 during the Eocene and Oligocene (Pagani et al., 2005) requires a major reorganization of the carbon cycle. Following the E/O boundary glaciation the earth has experienced periodic warming episodes, such as the late Oligocene warming and midMiocene climatic optimum (Zachos et al., 2001). Throughout these climate variations continental ice volume on Antarctic waxed and waned, but the continent never returned to an icefree, pre E/O state (Pekar and Christ ie Blick, 2008). The late Pleistocene ice core data demonstrate a clear relationship between temperature and atmospheric CO2 concentrations with atmospheric CO2 fluctuations of approximately ~90 ppm between glacial and interglacial climatic intervals (Petit et al., 1999; EPICA community members, 2004). The Last Glacial Maximum (LGM) ~21 kry ago represents the peak of the last glacial event and was characterized by lower sea level (~120 m lower, Peltier and Fairbanks, 2006), extensive continental ice growt h in the Northern hemisphere, changes in the deep ocean circulation (e.g., McManus et al., 2004; LynchStieglitz et al., 2007) and extreme aridity. The last deglaciation, which began ~18 kyr ago marked another climate threshold event that was characterized by a major decline in continental ice sheet volume. The tight coupling between atmospheric CO2 and temperature observed during the E/O transition and the last deglaciation illustrates the importance of the global
14 carbon cycle on climate over millions of years and millennial time scales. The longterm global carbon cycle responds to variations in CO2 sources, such as volcanism and weathering of organic carbon, and CO2 sinks, such as silicate weathering, over millions of years (Berner, 1999). In contrast, t he ocean participate in the short term global carbon cycle through the biological pump, ventilation rates, and whole ocean alkalinity (Sigman et al., 2010), which is again controlled by the rate and pathways of deepintermediate ocean circulation. Data on continental weathering and deepintermediate ocean circulation are critical for understanding the long and short term carbon cycles and climatic threshold events. In this dissertation, I study changes in weathering and ocean circulation for two climatologi cally important global events: the transition from a greenhouse to an icehouse world in at the E/O and the end of the last glaciation and the transition to modern climatic conditions. The key questions for the E/O study were whether we could develop an arc hive based on seawater Pb isotopes to identify changes in weathering conditions on Antarctica associated with climate change. The key question for the last deglacial study was whether we could identify the sources and pathways of old carbon that accumulated in the deep ocean during the glacial period and was transported to a region in the tropical Pacific during the deglaciation. This old carbon was ultimately vented to the atmosphere, thereby dramatically increasing the atmospheric CO2 concentration and de14C. Research Objectives The initial glaciers on Antarctica at the E/O boundary were probably wet based glaciers (Kennett and Barker, 1990; Zachos et al., 1999) that introduced massive pulses of continental material into the Southern Ocean. It is important to understand the extent
15 of chemical weathering and the composition of the source rocks exposed to weathering during this process because chemical weathering of silicate rocks sequesters atmospheric CO2, while mechanical weathering and weathering of carbonate have minimal impact on the carbon cycle. Other important observations from the EoceneOligocene Transition (EOT) include deepening of the CCD and a positive shift in the 13C record from benthic foraminifera, which could al so be affected by the type of material being weathered and the total cationic flux to the Southern Ocean. While the presence of ice rafted debris (IRD) is recognized as an indicator of mechanical weathering (Zachos et al., 1992a), IRD provides little infor mation about chemical weathering rates. Thus, in order to better understand the composition of weathered material (silicate vs. chemical) and the mode of weathering (mechanical vs. chemical) at the EOT, I developed a weathering proxy (Chapter 2; Basak et al., in press ) based on Pb isotopes that compares seawater values, which are recorded in oxide phases from bulk marine sediments, to values for terrigenous residues. This proxy is then used to reconstruct the weathering history on Antarctic during the EOT ( Chapter 3). One critical climatic tipping point occurred ~18 kyr at the end of the Last Glacial Maximum (LGM). This recent deglaciation is characterized by a fifty percent increase in atmospheric CO2 and a dramatic decrease in the radiocarbon content of t he atmosphere. Conceptual models indicate that extensive seaice cover (Stephen and Keeling, 2000) and upper ocean stratification in the glacial Southern Ocean (Francois et al., 1997; Sigman and Boyle, 2000) played a critical role in the storage of a dense, saline, CO2rich, and 14C depleted water. This water could have mixed back into the main body of the ocean during deglaciation as the seaice cover melted and ventilation
16 resumed in the Southern Ocean resulting in an increase in atmospheric CO2. Evidence for saline, 14C depleted (old) water during the LGM has been reported from the deep Southern Ocean (Adkins et al., 2002; Skinner et al., 2010). Recently, evidence of 14C depleted, old water was also reported from intermediate ocean depths (~700 m) in the tropical North Pacific and the Indian Ocean, suggesting an intermediate water pathway for transferring the CO2 sequestered in the deep ocean during the glacial to the atmosphere (Marchitto et al., 2007; Stott et al., 2009; Bryan et al., 2010). Previous s14C from foraminifera to determine the age of the water mass involved, but this proxy does not provide any water mass information. In this study Nd isotopes preserved in fossil fish teeth/debris were used (Chapter 4; Basak et al., 2010) to identify the source and pathway of the water mass that delivered old carbon to the tropical Pacific during the last deglaciation.
17 CHAPTER 2 SEAWATER Pb ISOTOPES EXTRACTED FROM CENOZOIC MARINE SEDIMENTS Overview Variations in ocean circulation and continental weathering contribute to global climate change through the redistribution of heat and impacts on the carbon cycle. Seawater Nd, and Pb isotopes record information related to these processes; thus, it is critical to identify archives capable of recording and preserving these isotopic systems over geologic time scales. Owing to the short residence time of Nd (~2001000 yrs; Tachikawa et al., 1999, 2003; Arsouze, 2009) compared to the ocean mixing time (~ 1500 yrs; Broecker and Peng, 1982) and dis tinct source inputs, Nd isotopic compositions can be used to track ocean circulation (Albareda and Goldstein, 1992; Albareda et al., 1997). In particular, different water masses have distinct Nd isotopic signatures that are acquired from continental inputs in the region of water mass formation. This signal, which is carried by water masses as they circulate through the oceans, can be altered by local weathering inputs, boundary exchange (Lacan and Jeandel, 2005) and reversible scavenging (Siddall et al., 2008; Jones et al., 2008) along the pathway. For this reason, Nd isotopes are often referred to as quasi conservative tracers for water mass. In contrast the residence time of Pb in the ocean (~50200 yrs; Cr aig et al., 1973; Schaule and Patterson, 1981; H enderson and Maier Reimer, 2002) is shorter than the Nd residence time. Therefore, seawater Pb isotopic compositions carry a combined signal of advection and weathering (Abouchami and Goldstein, 1995), which tends to be dominated by local weathering inputs For this reason, Pb isotopic compositions are
18 often used to study changes in weathering provenance (Reynolds et al. 1999), as well as weathering regimes and rates (Foster and Vance, 2006). Paleoceanographic studies of Nd isotopes have been conducted us ing a range of archives including FeMn crusts, extracts of Fe Mn oxides and hydroxides, fossil fish teeth/debris, and fossil foraminifera (for review see Martin et al., 2010). Early studies of Pb isotopic compositions in past seawater focused on FeMn cru sts (e.g. Abouchami et al., 1997; Burton et al., 1997; ONions et al., 1998; Reynolds et al., 1999; van de Flierdt et al., 2004; Foster and Vance, 2006); however, due to the slow accumulation rate this archive tends to integrate seawater signals over a per iod of 104105 years (Frank, 2002). For higher resolution studies Gutjahr et al., (2007, 2009) recently used Pb isotopic compositions recorded in dispersed FeMn oxides and hydroxides of bulk decarbonated marine sediments (henceforth referred to as FeMn phases) on Holocene/ Pleistocene time scales. These FeMn phases record bottom water Nd isotopic compositions (Rutberg et al., 2000; Palmer and Elderfield, 1985; Piotrowski et al., 2005; Martin et al., 2010), thus, they are believed to form at the sediment/water interface and to record bottom water Pb isotopic compositions as well. Gutjahr et al. (2007) systematically studied sequentially leached fractions of several Pleistocene deep marine sediment samples and determined that Pb isotopic compositions obtained through reductive extraction of FeMn phases are not perturbed by detrital contamination and the signal preserves seawater values. They then applied this technique to study continental weathering during the last deglaciation (Gutjahr et al., 2009; K urzweil et al., 2010). Recently, Haley et al. (2008) extracted Pb isotopes from carbonatefree Arctic sediments (~ last 15Ma) to obtain reliable bottom water
19 values; however, these samples were unique compared to typical marine sediments because of the low carbonate contents. In this study we evaluate the seawater origin of Pb isotopic compositions obtained by reductive extraction of bulk samples on longer, Cenozoic time scales to test the integrity of this archive for paleoseawater studies. Although it is generally assumed that the Pb isotopes recovered by reductive extraction are derived from FeMn phases in the form of oxides and hydoxides, the mineralogical sources of these isotopes have never been evaluated. Thus, we also evaluated the source of the ex tracted Pb using REEs and major element data. Finally, given that fossil fish teeth are well established archives for bottom water Nd isotopic studies, we also test the suitability of this archive for Pb isotope work. Methods Sample Preparation Ten deep sea sediment samples from ODP site 1090 in the Atlantic sector of the Southern Ocean were analyzed for this study. The samples range in age from mid Eocene to early Miocene (Channell et al., 2003) and their lithologies vary from mudbearing diatom ooze and mud and diatom bearing nannofossil ooze to chalk (Gersonde et al., 1999). In order to avoid contamination and minimize bl ank contributions during leaching, we used a slightly larger sample size than previous studies (Gutjahr et al., 2007, 2009). Approximately 1 gram of bulk sample was pulverized and homogenized with an agate mortar and pestle prior to chemical extraction. Nd or Pb isotopes have been extracted from bulk marine sediments using several protocols (Chester and Hughes, 1967; Rutberg et al., 2000; Bayon et al., 2002; Tovar Sanches et al., 2003; Piotrowski et al., 2004, 2005; Gutjahr et al., 2007; Martin et al., 2010). Although each protocol is slightly different all are guided by the principle of
20 exposing authigenically formed FeMn phases to a reducing solution. The first step for all of the protocols is decarbonation of the bulk sediment. We used two decarbonation steps each with 20 mls of buffered 1M Na acetate in 2.7% optima glacial acetic acid (buffered to pH=5) followed by an additional dec arbonation step in 10 mls of the buffered solution to insure complete carbonate removal. In order to minimize the blank in this reagent, which is largely contributed by the Naacetate, we added 1012 g of Chelex 100 cation exchange resin to a liter of buff ered acetic acid and let the solution sit for at least 24 hours before filtering. The blank for this Chelex cleaned buffered acetic acid is ~0.03 ppb and the reproducibility of Pb isotopic compositions extracted from samples decarbonated with this reagent improved remarkably. Following decarbonation, the sample was centrifuged and two separate 5 ml aliquots of the supernatant were removed; one for isotopic analyses and one for elemental analyses. The remaining sample was triple rinsed in four times distil led water (4x H2O) and then treated with 10 ml of 0.02 M Hydroxylamine hydrochloride (HH) in 25% acetic acid solution in order to reduce FeMn phases. The sample and HH solution were mounted on a rotary shaker for 1.5 hrs and degassed. After centrifuging, the supernatant was divided into two 5 ml aliquots for isotopic and elemental analyses, and this fraction was designated the first HH extraction. An additional 10 ml of fresh 0.02 M HH in 25% acetic acid solution were then added and allowed to react for 24 hours to insure complete reaction and removal of any remaining FeMn phases. Again, the supernatant was divided into two 5 ml aliquots for isotopic and elemental analyses, and this fraction was designated as the second HH extraction. Unless specified, the general term HH extraction refers to the first extraction. The aliquots set aside for
21 elemental analyses were analyzed for REE, U, Th, Pb, Mg, Al, Si, P, Ca, Mn, and Fe concentrations, while the isotopic fractions were analyzed for Sr, Nd, and Pb is otopic compositions. The remaining detrital residues were triple rinsed with 4x H2O and dried. Approximately 0.05 g of dried residue was powdered in a mortar and pestle and dissolved in a 1:2 mixture of concentrated HF: HNO3 (Optima grade) in a clean tef lon vial and allowed to react on a hot plate at >100C for 24 to 48 hrs. Once the residues were completely dissolved, the solution was dried down in preparation for further analyses. As for the extracts splits were collected for isotopic and elemental analyses. Splits of each of the bulk sediment samples were also sieved using plastic sieves to avoid metal contamination. The coarse fractions (>125 m) were dried and handpicked for fossil fish teeth, which were then subjected to an oxidative/reductive cl eaning to remove FeMn oxide coatings (Boyle and Keigwin, 1985; Rosenthal et al., 1997). Two to three fossil fish teeth from each sample were combined and dissolved in a 1:1 solution of HNO3 and HCl in preparation for Pb isotopic analyses. Column Chemist ry and Isotope Analyses All sample preparation work for elemental and isotopic work was done in a class 1000 clean lab in the Department of Geological Sciences at the University of Florida (UF). Aliquots separated for isotopic work were dried down before being redissolved in 1N seastar HBr and passed through Dowex 1X 8 (100200 mesh) resin (Manhes et al., 1978). Pb was collected in 20% Optima HNO3 and the eluted 1N HBr waste contained the REEs and Sr. This waste fraction was collected in clean Teflon vial s and dried down for subsequent Sr and Nd column chemistry. Sr and REEs were separated using a primary column with Mitsubishi resin and 1.6N HCl as an eluent. The Sr cuts were dried,
22 re 3, and passed through cation exchange columns containing Sr spec resin (Eichrom Technologies,Inc.) to further purify Sr from other ions (Pin and Bassin, 1992). Nd was isolated by passing the REE aliquot through Lnspec resin with 0.25N HCl as an eluent. Collected Pb, Nd and Sr concentrates wer e dried and re dissolved in 2% optima HNO3 in preparation for analysis on the Nu Plasma Multi Collector Inductively Coupled Plasma Mass Spectrometer (MC ICPMS) at UF. Measured procedural blanks were 20 pg Pb, 14 pg Nd and 100 pg for Sr. A Tl normalizati on technique was used for Pb isotope analyses (Kamenov et al., 2004). Tl spiked Pb concentrates dissolved in 2% HNO3 were adjusted for appropriate dilution to obtain a 25 V beam on 208Pb. Samples were aspirated through a NuInstruments desolvating nebulizer (DSN 100) into the plasma source. Long term NBS 981 values analyzed over several years at UF are 206Pb/204207Pb/204208Pb/204Pb=36.69 constrain the overall external reproducibility (including the chemical extraction, column chemistry, and MC ICPMS analyses), an UF inhouse standard was prepared by mixing multiple deep sea sediments. Four runs of this UF in house standard processed through the HH extraction protocol yielded values of 39.041( +/ 0.006), 15.667(+/ 0.003), 18.862 (+/ 0.006), for 208Pb/204Pb, 207Pb/204Pb, and 206Pb/204Pb respectively. Similar to Pb isotope work, Nd samples were diluted with unspiked 2% HNO3 and aspirated into the source of the MC ICPMS using the DSN100 nebulizer. Preamplifier gain calibrations were run before the beginning of each analytical session. All Nd isotope data reported in this study were analyzed using a timeresolved analysis (TRA) method (Kamenov et al., 2008). Prior to sample introduction baseline was measured for
23 30 seconds. Data were then acquired in a series of 0.2 seconds integrations over 13 minutes. All reported 143Nd/144Nd ratios were corrected for mass fractionation using 146Nd/144Nd =0.7219. International standard JNdi 1 was analyzed between every 56 unknown samples and the average of these standard runs was compared to the long term TIMS JNdi 1 value of 0.512103 to determine a correction factor for each of the samples analyzed on that day. Long term external reproducibility of JNdi 1 is 0.000014 (0.3 Nd units). Sr isotopes were analyzed on the NuPlasma MCICPMS in wet plasma mode using the same TRA method. Each sample was normalized to 86Sr/88Sr =0. 1194. All analyses used onpeak measured zeros determined on clean 2% HNO3 to correct for isobaric interferences due to Kr impurities in the Ar gas. The longterm average value of the TRA measured 87Sr/86Sr of NBS 987 is 0.710246 (+/ 0.000030, 2 ). M ajor E lements and REE Analyses Aliquots saved for REE, U, Th, Pb, and major element concentration measurements were dried down and redissolved in 5% HNO3 (spiked with 8 ppb ReRh) prior to analysis on an Element II ICP MS at UF. All analyses were done at medium resolution with Re and Rh as an internal standard to correct for instrumental drift. All concentration measurements were based on calibration curves generated using a set of gravimetrically prepared dilutions of commercial ICP MS standard (SPEX CertiPrep, Inc.). USGS whole rock standards (AGV 1, BCR2) and USGS Fe Mn nodule standards (NOD A1, NODP1) were frequently analyzed to test the quality of concentration measurements. Reported elemental concentrations were corrected for the analyzed acid blank measured at the beginning. Long term error based on multiple AGV 1 and BCR2 analyses was +/ 5%.
24 Results Pb Isotopes in Sequentially Extracted Fractions All three Pb isotope systems produce ratios that are within error for the first and second HH extractions for Eocene to Miocene samples, with values ranging between 18.6918.81 (206Pb/204Pb), 15.6515.67 (207Pb/204Pb), and 38.7638.86 (208Pb/204Pb) ( Figure 2 1; Table 2 1). The Pb isotopic compositions from the acetic acid decarbonation fraction also agree within error with the HH extractions for all of the Miocene and many of the Eocene Oligocene samples; however, one Oligocene and one Eocene acetic acid fraction plot well above the HH data, particularly for 207Pb/204Pb and 208Pb/204Pb, while one Eocene acetic acid fraction plots well below the HH data for 208Pb/204Pb ( Figure 2 1). Residue 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb values range from 18.6518.79, 15.6415.66, and 38.7638.88 respectively. 206Pb/204Pb values of the Eocene residues tend to be less radiogenic than the HH and acetic acid fractions, while the Oligocene residues are more radiogenic, and the M iocene residues agree with other fractions within error. In contrast, the residues tend to have 207Pb/204Pb values that are slightly lower, but generally within error of the other fractions, and 208Pb/204Pb values that are within error or slightly higher t han the other fractions. Pb Isotopes Compositions in Fossil Fish Teeth 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb isotopic ratios were analyzed for fossil fish teeth handpicked from the same samples as the extractions ( Figure 2 2; Table 2 2). 206Pb/204Pb, and 208Pb/204Pb isotopic compositions measured in fossil fish teeth were corrected for ingrowth of radiogenic Pb using measured U/Pb and Th/Pb concentration ratios. Because of the low abundance of 235U in the samples, corrections
25 for ingrowth of 207Pb on 207Pb/204Pb ratios are negligible. Similar corrections were not made for the HH extractions because accurate U/Pb and Th/Pb ratios were not obtained in the extract s and data from FeMn oxide nodules (Axelsson et al., 2002) indicated that expected rat ios were too small to require correction for Eocene to Miocene material. Measured Pb isotopic compositions for corrected fossil fish teeth, and HH extractions do not agree within error or exhibi t any systematic relationship ( Figure 2 2A, B). There is a general correlation between the magnitude of the correction and the age of the fossil fish teeth with older teeth r equiring a greater correction ( Figure 2 2D, E), but in detail the correction factor depends on U/Pb and Th/Pb ratios as well as age of the fossi l fish teeth. Major Elements in Fe Mn Extracted Fraction A suite of major elements (Mg, Al, Ca, Mn, Fe, P; Table 2 3) was analyzed for all fractions (Table 2 3). Major elements measured in HH extractions were compared to Pb concentrations in order to identify the source of the Pb in the extracted fractions, although this analysis might be complicated by multiple sources. A plot of Ca vs. Pb shows no distinct relationship, and three of the samples record anomalously high Ca and Pb concentrations ( Figure 2 3A ). These three outliers have been excluded in subsequent elemental plots. Mn and Pb exhibit a strong correlation (R2 =0.94); however, much of this correlation is guided by a single high point for both elements ( Figure2 3B ). A wide range of Fe, Mg, and P concentrations are found to correspond with very narrow range of Pb concentrations, suggesting no clear relationship between these elements and Pb.
26 REE in HH Extractions Rare earth element patterns obtained from the HH extraction of NOD A1 and NODP1 (representing Atlantic and Pacific oceans respectively) exhibit pronounced negative Ce anomalies and a slight middle REE bulge ( Figure 2 4A ). The middle REE bulge is similar to patterns reported for most si te 1090 samples in this study ( Figure 2 4B ), as we ll as patterns reported in other studies that employed an extraction technique (Haley et al., 2004, Gutjahr et al., 2007; Martin et al., 2010). In comparison, REE values derived from the total digestion of nodule standard NOD A1 and NOD P1 exhibit a slight to prominent positive Ce anomaly and a slightly reduced middle REE bulge (Axelsson et al., 2002; Figure 2 4A ). Three samples in this study had middle rare earth patterns that did not match the characteristic MREE bulge for HH extractions ( Figure 2 4B ). These same samples had anomalously high Ca, Mn, and Pb concentrations and low P concentrations in the HH extractions (Table 2 4) suggesting there is something fundamentally different about these samples. The acetic acid, HH extractions (first and second), and residue fractions form individual clusters on a plot of HREE/LREE ((Tm+Yb+Lu)/(La+Pr+Nd)) vs. MREE/MREE* ((Gd+Ty+Db)/[(Tm+Yb+Lu+La+Pr+Nd)/2]) (Figure 2 4C ) indicating a common source for each fraction. The first HH extractions for the three samples w ith anomalous REE patterns plot outside of the HH cluster ( Figure 2 4C ); however, the second HH extractions for these samples plot within the cluster. Discussion Integrity of Pb Isotopes from HH Extractions To qualify as an effective archive for deepwater Pb, the Pb isotopic compositions chemically extracted from decarbonated bulk sediments must be: a) reproducible, b)
27 unaffected by detrital components present during the extraction, and c) record bottom seawater values. Two sets of seven Paleogene deep sea samples were analyzed a year apart using the same protocol and the measured Pb isotope values for each sample are w ithin reported external error ( Figure 2 5A ; Table 2 5). Three Holocene samples from the Nort h Atlantic reported in Gutjahr et al. (2009) were also analyzed at UF and these replicate Pb isotopic compositions are also wi thin error of reported values ( Figure 2 5A ). Thus, both intraand inter lab comparisons illustrate that our extraction protocol produces reproducible results. One of the concerns of any chemical extraction technique is whether the procedure specifically targets the desired phases and does not introduce contamination from other phases, such as the detrital fraction. Although we spec ifically selected weak reagents and short leaching times, it is still important to verify that there is no contamination. One test for detrital contamination is to evaluate Al concentrations in the extracts based on the assumption that all Al is sourced fr om detrital material. This technique assumes there is no Al in the leachable fraction, therefore it tends to overestimate the detrital contribution to extraction solutions. Pb isotope values from the HH extractions were adjusted for the percent contribution from the detrital fraction based on the mass balance equations below, following a technique used by Gutjahr et al., (2007) for Sr isotopes. [Pb] % contrib. of detrital Pb = [Al] HH extraction/[Al] residue [Pb] detrital contrib. in HH extraction= [Pb] % c ontrib. of detrital Pb*[Pb] residue [Pb] sw contrib. in HH extraction = [Pb] HH extraction[Pb] detrital contr. in HH extraction
28 206 Pb/ 204 Pb sw corrected = ([Pb] HH extraction*206Pb/204Pb HH extraction) ( [Pb] detrital contr. in HH extraction* 206Pb/204Pb residue) [Pb] sw contrib. in HH extraction A comparison of 206Pb/204Pb values measured in the HH extractions and values corrected assuming all Al is derived from the detrital fraction illustrates that measured and corrected values plot within error (Figure 2 5B ). Therefore, even using the worst case scenario that all Al is derived from the detrital fraction, detrital contamination did not impact the Pb isotopic compositions of our samples. We recommend testing for detrital contamination on a case by case or site by site basis, especially for samples from regions with very low carbonate contents and high clay fractions, since different lithologies may produce different results. Given that detrital contamination did not affect the measured seawater Pb isotopic values, the similarity between seawater and residue Pb values in some of our samples requires an alternative explanation. One possibility is that for the relevant time intervals continental weathering input dominated the seawater Pb isotopic values. Therefore, these isotopic comparisons may be useful for understanding past weathering regimes. Another important aspect of this study was to demonstrate that Pb isotope values obtained through chemical extraction represents bottom water Pb. This v erification is complicated by the fact that bottom water Pb isotopic compositions are not known for older samples and even modern samples are likely to be compromised by anthropogenic Pb. 87Sr/86Sr ratios have been used in several Nd extraction studies to
29 demonstrate the seawater origin of Nd isotopic compositions (Rutberg et al., 2000; Piotrowski et al., 2005; Gutjahr et al., 2007; Martin et al., 2010). However, these studies point out that this is a highly conservative test. Many samples have unaltered Nd isotopic compositions despite the fact that Sr isotopic compositions have been altered, because Sr tends to be more mobile in the sediment than Nd. Similarly, any disagreement between the Sr isotopic values of HH extractions (Table 2 6) and the seawater c urve does not necessarily imply that the Pb isotopic compositions are altered. In contrast, agreement between Sr isotopic compositions in HH extractions and the seawater curve tends to confirm the seawater origin of the extractions, assuming there is no residual carbonate with seawater Sr isotope values remaining after the decarbonation step. Fifty percent of the 87Sr/86Sr ratios in our HH extractions do agree with seawater values ( Figure 2 5C ), which suggest, but does not insure, a seawater origin for the Pb isotopic compositions in those samples. Additional support for the seawater origin of the Pb isotopic compositions in the HH extractions can be obtained by comparing the HH extraction values to another archive from the same environment. This has been done successfully for Nd isotopes by comparing HH extractions and fossil fish teeth (Martin et al., 2010), because fish teeth have been shown to record bottom water values (Martin and Scher, 2004). However, based on this study, fish teeth do not appear to be robust archives for Pb isotopic compositions. Instead, we again compare Nd isotopes in the site 1090 HH extractions (Table 2 6) and fish teeth and illustrate that both recor d the same bottom water value ( Figure 2 5D ). Pb is very particle reactive and, thus, should be even less
30 mobile than Nd; therefore, a seawater origin for Nd in the HH extraction argues for a seawater origin for Pb as well. Further evidence for a seawater origin of Pb in HH extractions comes from studies of modern sediments. It is wellknown that anthropogenic Pb has a very distinct isotopic signal that dominates modern seawater Pb (e.g., Wu and Boyle, 1997; Alleman et al. 1999). Therefore, analyses of core top FeMn oxyhydroxides should record an anthropogenic Pb signal. Hamelin et al. (1989) analyzed leachates of modern surface sediments in the Atlantic and did recover anthropogenic Pb isotopic compositions indicating that the Pb dissolved in the ambient seawater is faithfully recorded in the leachable fractions. Furthermore, a recent study of low temperature FeMn oxyhydroxide sediments precipitating in the Mediterranean Sea also recovered a strong signal of anthropogenic Pb (Kamenov et al., 2009). Overall, these studies of modern sediments indicate that FeMn phases efficiently scav enge Pb derived from ambient seawater. Finally, Pb isotopic compositions from FeMn crusts and nodules are considered to be accurate proxies for ambient seawater. These crusts/nodules have a similar mineralogy to the FeMn phases (Martin et al., 2010) that are believed to contribute Pb during the HH extraction. In addition, Pb isotope values for HH extracts from FeMn nodule standards NOD A1 and NOD P1 (Table 2 1) are similar to published values derived from total digestion of these standards (Baker et al., 2004), suggesting the extraction procedure does not introduce analytical artifacts perturbing the isotopic ratios. In the absence of direct measurements of Pb in deep water samples or another Pb isotope archive known to record bottom water, these lines of indirect evidence provide
31 the strongest argument that Pb isotopic compositions in the chemical extractions record bottom water values. The Pb isotopes measured in the acetic acid fractions are broadly in agreement with the Pb isotopic compositions measur ed in the HH extractions, which could indicate that the decarbonation step is dissolving FeMn phases on the carbonate (Gourlan et al., 2008). Alternatively, carbonates are the major contributor to the acetic acid fraction, thus Pb introduced from this fr action is likely to record a surface water signal while HH extractions are believed to record a deep water signal. Agreement between these two fractions suggests little stratification of the Pb isotopes with depth in the water column for these samples. F ossil Fish Teeth as a Pb Isotope A rchi ve Fossil fish teeth are one of the most robust archives for bottom water Nd studies (Martin and Scher, 2004) and they have been applied in numerous studies of deep ocean circulation (e.g., Scher and Martin, 2004, 2006, 2008; Via and Thomas, 2006; Thomas and Via, 2007; Newkirk and Martin, 2009; Roberts et al., 2010). However, fossil fish teeth have never been tested as a suitable archive for Pb. In order to qualify as an archive for deepwater Pb isotopes we need to understand the source of Pb, U, and Th to the fossil fish teeth and also understand whether the hydroxyflouroapatite behaves as a closed system. Concentrations of Pb reported for fossil fish teeth in this study range from 30366 ppm (Table 2 7), while modern fish teeth analyzed during the course of thi s study contained 0.52.6 ppm. Based on these modern and fossil fish teeth Pb concentrations, the initial Pb incorporated during the life of the fish represents a maximum of 8.6% of the total Pb in fossilized fi sh teeth samples. The higher concentrations of Pb in fossil
32 teeth presumably represent post mortem or early diagenetic enrichment by several orders of magnitude, similar to that observed for Nd (Martin and Haley, 2000; Martin and Scher, 2004). In that case, the post mortem or early diagenetic Pb is likely to record deep water isotopic values that dominate the Pb isotopic signal of the fossil fish teeth. Substitution and adsorption are the two most commonly discussed enrichment mechanisms for REE, Pb, Th and U in biogenic apatite (Reynard et al., 1999; Trueman and Tuross, 2002; Grun et al., 2010). For the Pb system, both daughter and parent isotopes are incorporated during early diagenesis. As a result, Pb isotopic compositions in fossil fish teeth are compri sed of a complex mixture of unknown proportions of initial Pb presumably reflecting shallow seawater incorporated in vivo early diagenetic Pb with a deep water isotopic composition and radiogenic Pb produced in situ. Given that the initial Pb concentrat ion constitutes a small percentage of the total Pb pool in the fossil fish teeth, seawater Pb isotopes incorporated during early diagenesis are likely to dominate the system, indicating that suitable corrections for radiogenic Pb ingrowth should yield pris tine bottom water Pb isotopic compositions if the teeth behave as a closed system. Unfortunately, the magnitude of correction vs. age ( Figure 2 2D E ; Table 2 2) for Pb isotopes suggests that they exhibit open system behavior with r espect to U, Th, and/or Pb. As f igure 2 2 illustrates, the relationships between teeth and HH fractions are complex. In some cases, the corrected fish teeth plot below the HH fractions indicating either recent addition of U or Pb loss, which leads to overc orrection. In other cases the corrected ratios are higher than the HH fraction, suggesting either recent U loss or Pb addition, which leads to undercorrection. However, in one sample both the measured and corrected ratios plot below the HH
33 fraction suggest ing that a high initial Pb or intial Pb with very different isotopic signature compared to the bottom water may have played a role in the Pb isotope systematics. In summary, these data demonstrate that the teeth do not behave as a suitable archive for bott om water Pb isotopes. Previous studies attempting U series dating of bones have also concluded that bones remain an open system with respect to U (Pike et al., 2002). Similar claims of open system behavior in fossil fish teeth has also been made with res pect to REE mobilization; although the open system behavior was not pervasive enough to alter the geochemically useful Nd isotopic signal acquired during early diagenesis (Kocsis et al., 2010). Our study suggests that fossil fish teeth are not a suitabl e archive for seawater Pb isotopic studies due to presence of initial Pb, variable U, Th, Pb uptake and loss leading to unconstrained proportion of radiogenic growth, and open system behavior. However, there are limited data for Pb isotopic compositions of fossil fish teeth in the marine environment and this archive should be carefully evaluated with a larger data set spanning both older and younger samples before rejecting it as a suitable archive for paleoceanographic work. Phases Contributing Pb to the HH Extraction: Major Elements Although the correlation between Mn and Pb is dominated by one high point, these two elements generate the most robust correlation measured between Pb and any major element, which supports the idea that Mn oxides are a major phase contributing Pb. Published Mn/Fe ratios for complete digestion of the nodule standards range between ~1.65.0 (Axelsson et al., 2002). Mn/Fe ratios measured on HH extracts of NOD A1 and NOD P1 during the course of this study ranged from 3.64.6. The seven
34 HH extracts from site 1090 with normal REE patterns exhibit Mn/Fe ratios that varied between 0.74 and 3.70. Thus, measured Mn/Fe ratios in HH extracts from site 1090 are similar to the total range (i.e. 1.6 5.0) obtained from HH extracts and complet e digestion of NOD A1 and NOD P1, again supporting the idea that FeMn oxides are one of the major contributing phases to the total major element pool in the HH extracts. The three HH extraction samples with anomalously high Ca and Mn concentrations and uncharacteristic REE patterns have Ca concentrations that are even higher than concentrations measured in the acetic acid fractions (Table 2 3). These three samples were also the only samples that degased during the first HH extraction step. These anomalous samples also have high Mn/Fe ratios of 9.9 to 49.5, which are well beyond measured Mn/Fe ratios for FeMn oxides, suggesting the presence of a Mn rich phase that survived the buffered acetic acid dissolution step. Pure manganoan carbonate is one possible Mn rich, dissolution resilient phase (best represented by a mixed manganesecalcium magnesium carbonate) (Mucci, 2004). Boyle (1983) documented the presence of manganoan carbonate, specifically kutnahorite, in marine sediments. Kutnahorite has a general formula of Ca(Mn,Mg(CO3)2) and commonly occurs as overgrowths on foraminiferal calcite surfaces. X ray diffraction, geochemistry, and laser ablation studies of foraminiferal calcite tests later revealed the presence of kutnahorite or a kutnahoritelike phase (known as pseudokutnahorite due to its disordered structure, but similar chemical composition) in marine sediments (Pena et al., 2005, 2008a ). Authigenic manganoan carbonates could be common in the carbonaterich environment (Boyle, 1983) of site 1090 (3580% carbonate) and precipitation in the sediment would suggest a deep water
35 geochemistry. The relative stability of this phase demonstrated by preferential dissolution in the HH extraction rather than the buffered acetic acid step, suggests it mig ht have a chemical composition closer to the structurally ordered kutnahorite, which is more resistant to dissolution than disordered or pseudokutnahorite (Mucci, 1991). Thus, data on Fe/Mn ratios and anomalously high Ca and Mn concentrations in some HH extracts suggest contributions from FeMn oxides and manganoan carbonates respectively, indicating that the Pb measured in the HH extracts is sourced from multiple phases. However, both of these phases are authigenic minerals that form in bottom waters close to the sediment water interface. Thus, both should record the same bottom water isotopic values. Phases contributing Pb to the HH extraction : REEs The middle REEs (MREE) bulge pattern observed in reductive extracts has typically been attributed to FeMn oxides (Bayon et al., 2002; Gutjahr et al., 2007), with an emphasis on Feoxides (Haley et al., 2004). With the exception of multi valant Ce (III and IV) and Eu (II and III), the REEs are a group of trivalent elements that decrease in ionic radii wit h increasing atomic number. Seawater REE patterns, which have a prominent negative Ce anomaly and relatively heavy REEs enrichment, can be explained by conversion of Ce(III) to insoluble Ce(IV) and preferential scavenging of light REEs (LREE) compared to H REE from the surface ocean by organics, and detrital particles (Elderfield and Grieves, 1982; Elderfield, 1988). Authigenically precipitated FeMn phases in the ocean do not incorporate this seawater pattern, instead preferential scavenging of MREE produces a characteristic MREE bulge in the oxide phase that is apparent in the HH fractions as well (Axelsson et al., 2002; Haley et al., 2004; Gutjahr et al., 2007; Martin et al., 2010). A similar MREE bulge pattern is also produced by
36 phosphates (Byrne et al., 1996; Martin et al., 2010). Thus, MREE patterns observed in HH extracts could indicate extraction from FeMn oxides or from phosphates. To evaluate the contribution of Pb from phosphates we compared P and Pb concentrations in the acetic acid fractions and HH extractions. Concentrations of P in the buffered acetic acid fraction (buffered at pH=5) range from 0.02 282.0 ppm (Table 2 3), while concentrations in the HH extractions (pH<5) range from 8 407 ppm. Both of these fractions tend to have higher concentrations than HH extractions of NOD A1 and NODP1, which contained 12.7 27.2 ppm P. Higher concentrations in the site 1090 samples suggest a source of P other than FeMn oxides; however, there is no positive correlation between [P] and [Pb] in the H H extracts, which argues against phosphate as a major source of Pb, as does the fact that the teeth record different Pb isotopic compositions than the HH extracts. Charette et al. (2005) extracted oxides from bulk sediments and found high P association wit h Fe (hydr)oxides, which he attributed to P adsorbed on Fe(hydr)oxides. Well defined changes in phosphate sorption by iron oxyhydroxides defined by P/Fe ratios have also been reported for the last 3 Gyr of earths ocean chemistry evolution, indicating a c lose relationship between adsorbed phosphate and Fe oxyhydrox ides (Planavsky et al., 2010). Thus, an alternative source of high P concentration in the extracts could be absorbed phosphates. Trivalent REE patterns are remarkably similar for both HH extract ions and total digest ions of NODA1 and NOD P1 ( Figure 2 4A ); however, multivalent Ce patterns yield a positive Ce anomaly in the HH extractions and a negative Ce anomaly in the com plete digestions of standards ( Figure 2 4A ). This fractionation of Ce during the HH extraction procedure suggests that the reducing solution, which remineralizes other
37 REEs effectively, must not have a strong enough reduction potential to reduce Ce(IV) oxides. These results are distinct from Gutj ahr et al (2010), who report a positive Ce anomaly for their HH extracts of NOD A1 and NOD P1. While this study and Gutjahr et al. (2010) used broadly comparable extraction methods, Gutjahr et al. (2010) employed an additional MgCl2 step and used a slightl y stronger reducing solution. They also exposed smaller amounts of sediments to similar volumes of HH solution. These small differences in technique appear to produce very different results, implying that the extraction of Ce may be sensitive to subtle dif ferences in technique and studies of Ce anomalies based on HH extractions should be viewed with caution. In order to distinguish between different potential REE sources in different fractions we have made a HREE/LREE vs. MREE/MREE* plot using the same rat ios employed by Martin et al. (2010) ( Figure 2 4C ). The first and second HH extractions and standard nodule HH extracts generally fall within a narrowly defined same field. Detrital residues and acetic acid fractions also plot in distinct fields, although the acetic acid field covers a relatively large range. Clustering of HH and standard nodule extracts indicates a common source for the REE in these fractions supporting the idea that a primary Pb archive in our extraction procedure is FeMn oxyhydroxides and that the extraction protocol does not attack the detrital fractions, which plot in a different field. Martin et al., (2010) also showed that published fossil fish teeth and pore waters occupy the same space as the HH extracts. The three anomalous first HH extraction samples that have flatter REE patterns (open circ les in Figure 2 4B ) also plot separately in HREE/LREE vs. MREE/MREE* space. Coupled with their distinct major element concentrations these data imply a unique REE source, possibly manganoan car bonate
38 as suggested by major element data. The second HH extractions from these three samples plot in the main HH field, suggesting the Mncarbonate is largely removed during the first HH extraction step. Overall, f igure 2 4C illustrates that HREE/LREE and MREE/MREE* data may be a useful tool to test for cross contamination between the major fractions that constitute a typical deep sea sample before making paleoceanographic interpretations. Summary Pb isotopes were analyzed for the acetic acid fractions, first and second HH fractions, and detrital fractions for ten samples ranging in age from mid Eocene to early Miocene from ODP site 1090 on the Agulhus Ridge. The acetic acid and HH extraction fractions all tend to produce 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb values that are within error. Agreement between Pb isotopic compositions for these fractions either indicates a common deep water source or no water column stratification with respect to Pb isotopic compositions for these samples. Excellent inter and intralab reproducibility of Pb isotope in the HH extracts illustrates that the extraction protocol used in this study is an effective method for recovering seawater Pb isotopic compositions. Additional evidence for a seawater origin of the Pb isotopic compositions obtained from HH extraction includes a number of indirect tests, such as an Al mass balance approach to evaluate the impact of detrital contaminants, comparison of Sr isotopic compositions from the HH extraction and the seawater Sr isotope curve, and comparison of Nd isotopic compositions from the HH extraction and coexisting fossil fish teeth. We recommend applying the Al mass balance test to confirm low inputs from detrital material, particularly in low carbonate settings. Nodule standard NODA1 and NODP1 are recognized to record ambient bottom water isotopic signatures and close
39 agreement between total digestion and HH extracted Pb isotopic compositions from these nodules also indicates that the HH extracted Pb isotopes represent bottom water values. Thus, all tests suggest that the measured Pb isotopic compositions in HH extractions are not only seawater, but presumably record bottom water values. Fish teeth are one of the most widely accepted archives for deep ocean Nd isotopic compos itions, but this study demonstrated that the total Pb pool in fossil fish teeth is composed of an unconstrained combination of initial Pb incorporated in the biophosphate, Pb incorporated during diagenetic enrichment, and radiogenic Pb produced from in sit u decay of U and Th. A comparison of Pb isotopes measured on fossil fish teeth, and corrected for radiogenic growth in the teeth shows little correlation with isotopes from HH extractions, suggesting that fish teeth have a tendency to behave as an open sys tem with exchange of U, Pb, and Th between the teeth and surrounding waters. It is impossible to correct for this exchange, thus fish teeth may not be effective archives for paleo Pb isotope studies. This evaluation is based on samples from one site and a limited age range, thus a more systematic evaluation of fossil fish teeth archive is warranted. Elemental data do not show any strong correlation between most major elements and Pb; however, there is a positive correlation between [Mn] and [Pb] based largely on one high data point suggesting a predominance of Mn oxides. Anomalously high Ca and Mn concentrations in HH extracts of three samples that also have anomalous REE patterns suggest the presence of a Mnbearing, refractory carbonate phase, such as manganoan carbonate. Therefore, FeMn oxides, and Mncarbonate appear to be
40 potential sources of the Pb isotope signals. Both of these phases form at the sediment water interface and record the bottom water signal. A characteristic MREE bulge which is very often attributed to the presence of oxides is observed in most of the HH extracts. In addition, clustering of most HH extracts with nodule standards in HREE/LREE vs. MREE/MREE* space indicates that FeMn oxides may be the dominant source of Pb. The t hree samples with anomalous REE patterns and elemental concentrations that plot outside the acetic acid, HH, and residue clusters presumably reflect the manganoan carbonate source. These results illustrate that the REE crossplot is an effective tool for id entifying cross contamination or unique sources. A comparison between chemical analyses of completely dissolved FeMn nodule standards and HH extractions from these same standards reveals a typical MREE bulge pattern for trivalent REEs for both fractions however, there is apparent fractionation with respect to multivalent Ce (using the extraction procedure reported in this study ). Changes in the strength of the HH solution and sediment to solution ratio may yield more quantitative extraction of Ce. Thus caution is advised when interpreting Ce anomalies from HH extractions, but these extractions from bulk deep sea sediments appear to be excellent recorders of bottom water Pb isotopes.
41 Table 2 1 Pb isotopes in different bulk sediment fractions from ODP site 1090 Age(Ma) mcd 206 Pb/ 204 Pb 207 Pb/ 204 Pb 208 Pb/ 204 Pb Acetic decarbonation 1090D,11 4A,38 41 20.31 104.91 18.777 15.658 38.843 1090E,11 5A,72 75 20.79 110.73 18.755 15.655 38.836 1090E,12 3A,71 74 21.61 116.62 18.775 15.656 38.846 1090D,17 3A,18 21 25.06 162.96 18.749 15.696 38.905 1 0 90D,19 3A,90 93 26.98 185.45 18.729 15.655 38.800 1090B,31 1W,16 19 35.58 293.85 18.705 15.646 38.758 1 0 90B,37 1W,124 126 37.68 342.99 18.727 15.651 38.754 1090B,38 2W,123 126 38.71 352.32 18.686 15.652 38.754 1090B,38 7W,42 46 39.51 361.01 18.733 15.656 38.776 1090B,40 3W,124 126 41.26 375.03 19.018 15.694 38.941 First HH Extraction 1090D,11 4A,38 41 20.31 104.91 18.782 15.651 38.838 1090E,11 5A,72 75 20.79 110.73 18.768 15.656 38.856 1090E,12 3A,71 74 21.61 116.62 18.774 15.657 38.852 1090D,17 3A,18 21 25.06 162.96 18.723 15.656 38.827 1 0 90D,19 3A,90 93 26.98 185.45 18.739 15.660 38.830 1090B,31 1W,16 19 35.58 293.85 18.714 15.655 38.808 1 0 90B,37 1W,124 126 37.68 342.99 18.730 15.650 38.759 1090B,38 2W,123 126 38.71 352.32 18.690 15.652 38.765 1090B,38 7W,42 46 39.51 361.01 18.735 15.658 38.782 1090B,40 3W,124 126 41.26 375.03 18.808 15.671 38.846 NOD A1 18.958 15.674 38.917 NOD P1 18.696 15.630 38.662 Second HH Extraction 1090D,11 4A,38 41 20.31 104.91 18.776 15.654 38.846 1090E,11 5A,72 75 20.79 110.73 18.766 15.655 38.853 1090E,12 3A,71 74 21.61 116.62 18.772 15.655 38.847 1090D,17 3A,18 21 25.06 162.96 18.725 15.659 38.828 1 0 90D,19 3A,90 93 26.98 185.45 18.733 15.656 38.811 1090B,31 1W,16 19 35.58 293.85 18.700 15.650 38.804 1 0 90B,37 1W,124 126 37.68 342.99 18.730 15.650 38.762 1090B,38 2W,123 126 38.71 352.32 18.693 15.654 38.775 1090B,38 7W,42 46 39.51 361.01 18.738 15.658 38.789 1090B,40 3W,124 126 41.26 375.03 18.819 15.670 38.852 Residues 1090D,11 4A,38 41 20.31 104.91 18.763 15.651 38.849 1090E,11 5A,72 75 20.79 110.73 18.757 15.652 38.861 1090E,12 3A,71 74 21.61 116.62 18.763 15.656 38.865 1090D,17 3A,18 21 25.06 162.96 18.768 15.659 38.878 1 0 90D,19 3A,90 93 26.98 185.45 18.763 15.655 38.852 1090B,31 1W,16 19 35.58 293.85 18.676 15.646 38.821 1 0 90B,37 1W,124 126 37.68 342.99 18.706 15.644 38.764 1090B,38 2W,123 126 38.71 352.32 18.650 15.646 38.757 1090B,38 7W,42 46 39.51 361.01 18.706 15.652 38.769 1090B,40 3W,124 126 41.26 375.03 18.791 15.664 38.837 Long term NBS 981 values analyzed over several years on the MC ICPMS at UF are 2 06Pb/204207Pb/204208Pb/204
42 Table 2 2. Pb isotopes measured in fossil fish teeth Fossil fish teeth Age(Ma)* 208 Pb/ 204 Pb 207 Pb/ 204 Pb 206 Pb/ 204 Pb Corrected 206 Pb/ 204 Pb Corrected 208 Pb/ 204 Pb 1090D,11 4A,38 41 20.31 38.805 15.643 18.776 18.742 38.793 1090E,11 5A,72 75 20.79 38.822 15.654 18.781 18.741 38.810 1090E,12 3A,71 74 21.61 38.845 15.658 18.869 18.802 38.837 1090D,17 3A,18 21 25.06 38.671 15.646 18.687 18.634 38.651 1 0 90D,19 3A,90 93 26.98 38.785 15.653 18.748 18.717 38.775 1090B,31 1W,16 19 35.58 38.776 15.633 18.744 18.640 38.717 1 0 90B,37 1W,124 126 37.68 38.749 15.649 18.790 18.701 38.697 1090B,38 2W,123 126 38.71 38.920 15.697 18.812 18.750 38.829 Age model after Channell et al. (2003) Long term NBS 981 values analyzed over several years on the MC ICPMS at UF are 206Pb/204207Pb/204Pb=15.489 208Pb/204
43 Table 2 3. Maj or elements in different bulk sediment fractions f rom ODP Site 1090. Mg Al Si P Ca Mn Fe Pb Al/Pb Mn/Fe Acetic decarbonation 1090D,11 4A,38 41 647 141 6162 614 13.54 0.51 45.37 1090E,11 5A,72 75 1181 290 24 3378 448 0.90 1090E,12 3A,71 74 1618 17 338 192 1017 200 1.15 15 1090D,17 3A,18 21 1601 11 327 282 58 396 5.00 0.57 19 79.28 1 0 90D,19 3A,90 93 1618 34 450 216 557 732 2.62 1.13 30 279.39 1090B,31 1W,16 19 1432 235 44 2470 1530 0.27 2.73 1 0 90B,37 1W,124 126 868 9 295 17 2762 1364 2.91 3 1090B,38 2W,123 126 1020 270 118 743 410 1.29 1090B,38 7W,42 46 * 1.47 1090B,40 3W,124 126 397 24 6116 403 0.47 First HH Extraction Mg Al Si P Ca Mn Fe Pb Al/Pb Mn/Fe 1090D,11 4A,38 41 51 47 45 12 18448 236 4.75 0.31 151 49.55 1090E,11 5A,72 75 128 224 119 281 135 21 12.19 0.14 1555 1.68 1090E,12 3A,71 74 159 268 158 273 63 96 25.85 0.47 571 3.71 1090D,17 3A,18 21 112 154 154 166 31 19 12.75 0.14 1127 1.53 1 0 90D,19 3A,90 93 96 153 134 319 58 24 9.61 0.19 811 2.53 1090B,31 1W,16 19 68 157 116 347 77 24 32.67 0.09 1669 0.74 1 0 90B,37 1W,124 126 41 94 161 407 70 21 16.28 0.17 543 1.27 1090B,38 2W,123 126 12 38 93 395 39 5 3.65 0.09 407 1.50 1090B,38 7W,42 46 70 29 76 21 12154 103 10.39 1.14 25 9.92 1090B,40 3W,124 126 131 14 30 8 38704 268 10.02 3.87 4 26.75 NOD A1 1918 808 400 27 349 6037 1673.09 0.12 6538 3.61 NOD P1 1291 332 683 13 133 5226 1135.63 0.07 4795 4.60 Second HH Extraction Mg Al Si P Ca Mn Fe Pb Al/Pb Mn/Fe 1090D,11 4A,38 41 65 155 97 44 718 11 23.31 0.11 1354 0.46 1090E,11 5A,72 75 350 549 261 19 777 33 42.62 0.47 1167 0.76 1090E,12 3A,71 74 491 903 403 27 739 184 140.54 1.14 792 1.31 1090D,17 3A,18 21 484 677 364 9 376 64 59.97 0.65 1036 1.07 1 0 90D,19 3A,90 93 472 510 379 7 575 56 47.17 0.80 635 1.19 1090B,31 1W,16 19 255 243 273 10 491 19 84.76 0.38 636 0.22 10 90B,37 1W,124 126 403 161 400 11 1003 46 34.50 0.52 309 1.34 1090B,38 2W,123 126 142 56 296 103 439 16 22.45 0.10 545 0.71 1090B,38 7W,42 46 92 158 244 945 2790 50 20.75 0.24 672 2.42 1090B,40 3W,124 126 45 70 116 226 1272 25 14.06 0.16 448 1.79
44 Table 23 Continued Residue Mg Al Si P Ca Mn Fe Pb Al/Pb Mn/Fe 1090D,11 4A,38 41 11165 71510 479 189 8715 899 38606 13.93 5133 0.02 1090E,11 5A,72 75 11656 67094 494 183 7498 1050 38966 13.91 4822 0.03 1090E,12 3A,71 74 11003 66095 1508 211 8992 1907 38617 14.84 4454 0.05 1090D,17 3A,18 21 12183 75618 1556 169 7461 1376 38969 14.84 5094 0.04 1 0 90D,19 3A,90 93 8081 44635 1627 101 4320 773 23011 9.49 4702 0.03 1090B,31 1W,16 19 9436 40557 786 68 7049 235 27498 4.18 9700 0.01 1 0 90B,37 1W,124 126 19728 74079 4127 99 12792 676 54202 10.27 7216 0.01 1090B,38 2W,123 126 14436 86119 546 150 17479 1056 44024 17.32 4972 0.02 1090B,38 7W,42 46 13340 73907 887 1023 21502 1409 40025 18.98 3895 0.04 1090B,40 3W,124 126 15460 71878 1020 212 13725 1687 45079 8.76 8207 0.04 no data
45 Table 2 4. REE data in different bulk sediment fractions from ODP site 1090. Acetic decarbonation La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 1090D,11 4A,38 41 1.00 1.13 0.15 0.63 0.12 0.03 0.13 0.02 0.14 0.03 0.10 0.02 0.12 0.02 1090E ,11 5A,72 75 1.08 1.85 0.20 0.94 0.22 0.06 0.26 0.04 0.29 0.07 0.24 0.04 0.33 0.06 1090E,12 3A,71 74 3.03 5.15 0.72 3.73 0.91 0.24 1.09 0.15 1.00 0.22 0.65 0.08 0.59 0.10 1090D,17 3A,18 21 4.52 9.86 1.19 6.08 1.47 0.38 1.75 0.25 1.54 0.33 0.97 0.13 0.89 0.15 1090D,19 3A,90 93 3.11 6.84 0.72 3.73 0.89 0.24 1.11 0.16 1.04 0.23 0.69 0.10 0.68 0.11 1090B,31 1W,16 19 1.69 1.68 0.22 0.90 0.18 0.06 0.20 0.03 0.25 0.06 0.21 0.04 0.29 0.05 1090B,37 1W,124 126 1.34 2.57 0.22 0.99 0.22 0.06 0.25 0.04 0.27 0.06 0.19 0.03 0.21 0.03 1090B,38 2W,123 126 1.07 2.25 0.22 1.20 0.31 0.09 0.43 0.06 0.39 0.09 0.27 0.04 0.25 0.04 1090B,38 7W,42 46 0.91 0.91 0.12 0.48 0.10 0.03 0.11 0.02 0.13 0.03 0.11 0.02 0.16 0.03 1090B,40 3W,124 126 0.96 0.92 0.12 0.47 0.09 0.03 0.09 0.02 0.11 0.03 0.09 0.01 0.12 0.02 First HH extraction La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 1090D,11 4A,38 41 1.37 1.93 0.25 0.76 0.13 0.04 0.10 0.02 0.11 0.02 0.07 0.01 0.08 0.01 1090E,11 5A,72 75 1.89 4.32 0.83 3.85 1.05 0.27 0.93 0.13 0.67 0.12 0.33 0.04 0.30 0.04 1090E,12 3A,71 74 2.34 4.93 1.00 4.62 1.22 0.30 1.03 0.15 0.80 0.14 0.38 0.05 0.36 0.05 1090D,17 3A,18 21 0.72 2.18 0.32 1.49 0.42 0.10 0.35 0.05 0.26 0.04 0.12 0.02 0.11 0.02 1090D,19 3A,90 93 1.01 2.98 0.43 2.04 0.56 0.14 0.48 0.07 0.37 0.07 0.18 0.02 0.16 0.02 1090B,31 1W,16 19 1.26 3.37 0.48 2.20 0.56 0.13 0.48 0.07 0.35 0.06 0.17 0.02 0.15 0.02 1090B,37 1W,124 126 0.51 1.80 0.24 1.14 0.33 0.08 0.27 0.04 0.20 0.03 0.09 0.01 0.08 0.01 1090B,38 2W,123 126 0.56 2.01 0.26 1.29 0.38 0.09 0.32 0.05 0.24 0.04 0.11 0.01 0.10 0.01 1090B,38 7W,42 46 1.17 2.08 0.21 0.78 0.15 0.05 0.14 0.02 0.18 0.04 0.12 0.02 0.13 0.02 1090B,40 3W,124 126 3.10 3.60 0.42 1.34 0.20 0.06 0.15 0.03 0.21 0.04 0.14 0.02 0.18 0.02 NOD A1 6.63 7.25 1.46 7.39 1.81 0.48 2.19 0.29 1.52 0.31 0.86 0.11 0.72 0.12 NOD P1 5.19 2.54 1.77 9.12 2.51 0.64 2.51 0.35 1.71 0.31 0.84 0.11 0.78 0.13 Second HH extraction La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 1090D,11 4A,38 41 0.60 1.14 0.23 1.04 0.25 0.06 0.23 0.03 0.16 0.03 0.08 0.01 0.06 0.01 1090E,11 5A,72 75 2.70 5.60 1.01 4.46 1.04 0.25 0.96 0.13 0.73 0.14 0.37 0.05 0.30 0.04 1090E,12 3A,71 74 4.00 7.64 1.47 6.36 1.48 0.35 1.32 0.19 1.04 0.19 0.53 0.07 0.46 0.06 1090D,17 3A,18 21 1.94 5.33 0.74 3.15 0.76 0.18 0.65 0.09 0.51 0.09 0.25 0.03 0.22 0.03 1090D,19 3A,90 93 2.23 6.10 0.86 3.81 0.94 0.23 0.85 0.12 0.66 0.12 0.33 0.04 0.28 0.04 1090B,31 1W,16 19 1.56 3.65 0.52 2.32 0.54 0.13 0.50 0.07 0.37 0.07 0.18 0.02 0.15 0.02 1090B,37 1W,124 126 0.86 2.64 0.35 1.65 0.44 0.10 0.37 0.05 0.26 0.05 0.12 0.02 0.10 0.01 1090B,38 2W,123 126 0.33 1.09 0.14 0.67 0.18 0.04 0.15 0.02 0.10 0.02 0.05 0.01 0.04 0.01 1090B,38 7W,42 46 3.90 8.72 1.36 6.19 1.45 0.38 1.40 0.19 1.08 0.21 0.57 0.07 0.46 0.06 1090B,40 3W,124 126 0.44 1.35 0.19 0.91 0.24 0.06 0.21 0.03 0.15 0.03 0.07 0.01 0.06 0.01
46 Table 2 4. Continued Resid ues La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 1090D,11 4A,38 41 28.86 55.35 7.26 26.74 4.94 1.37 4.63 0.67 3.76 0.73 2.11 0.28 1.80 0.25 1090E,11 5A,72 75 33.93 64.72 8.34 30.54 5.71 1.51 5.42 0.79 4.60 0.91 2.63 0.36 2.28 0.31 1090E,12 3A,71 74 30.58 58.27 7.74 28.14 5.30 1.46 4.97 0.74 4.34 0.84 2.49 0.35 2.28 0.31 1090D,17 3A,18 21 23.37 53.52 5.78 20.55 3.76 1.08 3.32 0.48 2.77 0.54 1.64 0.24 1.59 0.22 1090D,19 3A,90 93 17.42 39.70 4.41 16.37 3.07 0.85 2.92 0.43 2.50 0.50 1.46 0.20 1.31 0.19 1090B,31 1W,16 19 19.52 40.04 4.49 16.65 3.04 0.90 2.96 0.43 2.48 0.49 1.38 0.18 1.15 0.16 1090B,37 1W,124 126 26.97 69.12 7.15 27.15 5.33 1.60 4.93 0.72 4.00 0.76 2.13 0.28 1.78 0.24 1090B,38 2W,123 126 40.69 107.58 10.78 40.85 8.12 2.70 7.27 1.04 5.87 1.12 3.19 0.44 2.85 0.39 1090B,38 7W,42 46 98.08 200.44 25.80 97.72 19.44 5.24 18.75 2.75 16.16 3.10 8.89 1.16 7.26 0.96 1090B,40 3W,124 126 38.35 96.60 10.67 41.25 8.18 2.26 7.66 1.10 6.35 1.23 3.50 0.46 2.81 0.37 All d ata are normalized to initial sample bulk weight and concentrations are expressed in g/g
47 Table 2 5 Inter and intra laboratory Pb isotope duplicates of authigenic FeMn phases. mbsf 208 Pb/ 204 Pb 207 Pb/ 204 Pb 206 Pb/ 204 Pb 1139A 27R,1W,15 20 252.5 38.940 15.633 18.659 Duplicate 38.961 15.634 18.662 1139A 27R,02W,20 25 255 38.904 15.612 18.543 Duplicate 38.913 15.613 18.556 1139A,30R,01W,26 31 282 38.886 15.613 18.542 Duplicate 38.900 15.621 18.548 1139A,35R,2W,34 38 332.5 38.827 15.645 18.669 Duplicate 38.834 15.649 18.665 1139A,38R,4W,9 13 356 38.841 15.633 18.566 Duplicate 38.847 15.636 18.570 1139A,40R,3W,0 4 375 38.780 15.618 18.552 Duplicate 38.769 15.614 18.547 1139A,40R,1W,12 16 385 38.889 15.665 18.768 Duplicate 38.876 15.661 18.758 12JCP, 4250m* 50cm 39.532 15.690 19.330 77cm 39.217 15.667 19.137 251cm 38.882 15.632 18.939 All data are normalized to NIST SRM 981 values as reported in Galer and Abouchami (1998). Duplicate samples were chemically processed and analyzed one year after the initial samples. Replicate of samples from Gutjhar er al.(2009) analyzed at UF.
48 Table 2 6 Sr and Nd isotopes measured in authigenic FeMn phases. First HH extraction Age(Ma) 1 87 Sr/ 86 Sr 2 143 Nd/ 144 Nd 3 Nd(t) 1090D,11 4A,38 41 20.31 0.708093 0.512234 7.38 1090E,11 5A,72 75 20.79 0.708403 0.512199 8.05 1090D,12 3A,71 74 21.61 0.708304 0.512211 7.80 1090D,17 3A,18 21 25.06 0.708125 0.512245 6.71 1090D,19 3A,90 93 26.98 0.708052 0.512247 6.96 1090B,31 1W,16 19 35.58 0.707759 0.512275 6.20 1090B,37 1W,124 126 37.68 0.70854 0.512231 7.00 1090B,38 2W,123 126 38.71 0.707594 0.512248 6.65 1090B,38 7W,42 46 39.51 0.707576 0.512239 6.80 1090B,40 3W,124 126 41.26 0.707664 0.512164 8.22 1 Age model after Channell et al.(2003) 2Sr ratios were measured in a static mode and using time resolved analysis (TRA) on a Nu MC ICPMS and 87Sr/86Sr ratio s w ere corrected for mass bias using an exponential law and 87Sr/88Sr=0.1194. The longterm average value of the TRA measured 87Sr/86Sr of NBS 987 is 0.710246 (+/ 3All reported 143Nd/144Nd ratios were corrected for mass fractionation using 146Nd/144Nd =0.7219. International stand ard JNdi 1 was analyzed between every 56 unknown samples and the average of these standard runs were compared to a long term TIMS JNdi 1 value of 0.512103 to determine a correction factor for each of the samples analyzed on that day. Long term external reproducibility of JNdi Nd units).
49 Table 2 7 Measured Pb, U, and Th concentrations in fossil and modern fish teeth. Fossil fish teeth Age (Ma) [Pb] [U] [Th] 1090D,11 4A,38 41 20.31 191.62 28.53 31.25 1090E,11 5A,72 75 20.79 161.34 28.10 25.72 1090E,12 3A,71 74 21.61 50.95 13.86 4.81 1090D,17 3A,18 21 25.06 102.74 19.33 23.06 0190D,19 3A,90 93 26.98 366.61 37.23 36.58 1090B,31 1W,16 19 35.58 17.30 4.51 7.91 0190B,37 1W,124 126 37.68 29.96 6.24 11.38 1090B,38 2W,123 126 38.71 38.14 5.38 24.62 Modern fish teeth [Pb] [U] [Th] 13_14_Mackerel 0.71 0.24 0.15 13_14_Mackerel_Replicate 0.49 0.24 0.13 15_16_Mackerel 1.25 0.30 0.16 15_16_Mackerel_Replicate 2.03 0.27 0.14 7_8_unknown 2.61 0.25 0.13 7_8_unknown_Replicate 1.62 0.23 0.12 23_24_Halibut 1.48 0.26 0.13 23_24_Halibut_Replicate 0.55 0.24 0.12
50 Figure 21. Pb isotopes measured in the initial acetic acid decarbonation fraction, the first and second HH extraction, and residues from deep sea sediments at ODP site 1090 ranging in age from mid Eocene to early Miocene. (a) 206Pb/204Pb, (b) 207Pb/204Pb, and (c) 208Pb/204Pb.
51 Figure 22 Measured Pb isotopic values for fossil fish teeth, fossil fish teeth corrected for radiogenic growth, and for HH extractions plotted against depth (meters composite depth) (a) 206Pb/204Pb and (b) 208Pb/204Pb, (c) 207Pb/204Pb (fossil fish teeth were not corrected for this system). (d) The magnitude of the correction for radiogenic Pb ingrowth for 206Pb/204Pb and (e) 208Pb/204Pb in fossil fish teeth plotted against age illustrates a poorly defined rela tionship with age.
52 Figure 23 Plots showing relationship of measured major elemental concentrations with respect to Pb concentrations. (a) Pb vs. Ca concentrations for all ten site 1090 HH extractions. The three circled data points are wi th anomalously high [Ca]. Pb concentrations vs. (b) Mn, (c) Fe, (d) Mg, and (e) P major element concentrations for the seven first HH extractions with typical REE patterns. [Pb] vs. [Mn] is the only plot with a positive linear correlation, indicating that most of the Pb in the HH extraction is coming from Mnbearing phases. In contrast, the other [Pb] vs. [major element] plots do not illustrate a coherent increase in Pb that correlates with an increase in the major element.
53 Figure 24 Plots showing REE behavior of the extracted fractions. (a) REE patterns normalized to Post Archean Australian Shale (PAAS, Taylor and McLennan, 1985) for nodule standards NOD A1 and NOD P1. Blue open circles are published data for a total digestion of the standards (Ax elsson et al., 2002). Brown filled circles represent measured values for HH extractions. (b) REE/PAAS patterns for HH extractions from all samples from ODP site 1090. Green filled circles represent samples that exhibit a typical MREE bulge. Green open circles represent samples with unusual REE patterns that also have unusual major element concentrations. (c) A comparison of PAAS normalized HREE/LREE ((Tm+Yb+Lu)/(La+Pr+Nd)) vs. MREE/MREE* ((Gd+Ty+Db)/average of HREE and LREE) for all fractions.
54 Figure 2 5. Compilation of plots showing the reproducibility and integrity of the extraction procedure. (a) A plot of seven replicate 206Pb/204Pb ratios for HH extractions from ODP site 1139 measured one year apart at UF and three samples from Blake Nose (Gutjahr et al., 2009) that were also analyzed at UF Error bars are smaller than symbols (b) A plot of 206Pb/204Pb vs. depth comparing measured HH extraction values to values that have been corrected for detrital contamination based on [Al]. (c) 87Sr/86Sr vs. a ge for the seawater curve (blue circles; Hodell and Woodruff (1994), Farrell et al. (1995), Mead and Hodell (1995), and Martin et al. (1999)) and HH extractions from this study (red circle for HH extractions and fossil fish teeth (Scher and Martin, 2006) i llustrates excellent agreement between the two archives.
55 CHAPTER 3 WEATHERING HISTORY OF ANTARCTICA DURING THE EOCENEOLIGOCENE TRANSITION The initiation of Antarctic continental glaciation at the end of the Eocene is one of the most dramatic clim ate events in the Cenozoic. Although ephemeral glaciers have been documented during the late Eocene (Browning et al., 1996; Tripati et al., 2005), continental scale glaciers that reached the sea did not develop until the EoceneOligocene boundary ~ 34Ma ago. This interval has been intently studied due to both the rate and magnitude of this climatic phenomenon. Several mechanisms have been proposed to explain the onset of the Antarctic cryosphere. These include thermal isolation of Antarctica due to tectonic gateway processes (Kennett, 1977), and decreasing greenhouse gas concentrations (primarily CO2) below a threshold value (DeConto and Pollard, 2003) combined with orbitally paced climate cycles (Palike et al., 2006). In addition to massive continental ic e growth, the Eocene Oligocene (E/O) greenhouse to icehouse transition is characterized by a rapid two step increase in benthic foraminifera 18O (~1.5), worldwide deepening of the calcite compensation depth (CCD) by ~1 km, a positive shift in seawater 13C (Coxall et al., 2005), occurrences of ice rafted debris (IRD) in the Southern Ocean (Zachos et al., 1992), and a shift in clay mineral assemblages from smectite to illite and chlorite in the Southern Ocean (Robert and Kennett, 1997; Robert et al. 2002). While ice growth and deep water cooling can explain the shifts in oxygen isotopes, they do not account for the carbon isotope shift or deepening of the CCD. Instead these changes have been attributed to increased carbon burial (Zachos et al., 1996; Salamy and Zachos, 1999), increased silicate weathering (Ravizza and Peucker Ehrenbrink, 2003; Zachos and Kump, 2005),
56 increased export production of global siliceous plankton (Coxall et al., 2005), and a shift in CaCO3 sedimentation from the shallow shelf to deep ocean basins (Opdyke and Wilkinson, 1988; Kump and Arthur, 1997; Merico et al., 2008). Continental weathering of silicate rocks is a key component of the long term carbon cycle. Congruent chemical weathering produces a weathering solution with the sam e chemical signature as the parent rock. In contrast, incongruent chemical weathering produces a weathering solution with a distinct chemical signature compared to the parent material. Previous studies illustrated that incongruent weathering of silicate ro cks preferentially releases a labile, radiogenic Pb fraction (Erel et al., 1994; Harlavan et al., 1998). Expanding on this concept, one of the goals of this study is to test the idea that combined mechanical and incongruent weathering will yield seawater P b isotopes that record this more radiogenic fraction, while the corresponding detrital fraction of the marine sediment will maintain its original composition or a slightly less radiogenic composition. In contrast, during times of more intense chemical weat hering both seawater and the detrital fractions should record values closer to the whole source rock Pb isotopic value. Seawater Pb isotopic signatures may also distinguish between carbonate and silicate weathering. Carbonates contain high initial U/Pb and low initial Th/Pb ratios, therefore chemical weathering of carbonates should produce high 206Pb/204Pb values in the weathering solutions that are not accompanied by high 208Pb/204Pb. It is also likely that the seawater 207Pb/204Pb produced by weathering Phanerozoic carbonates would not change, because the flux of 207Pb has been low for last several hundred million years due to the short half life of 235U. In comparison, weathering of silicates is likely to
57 produce solutions with variable 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios that depend on the composition of the parent rock and the intensity of weathering. Unlike Pb isotopes, seawater Nd isotopes primarily track water mass mixing; however, during times of intense weathering Nd isotopes in seawater have also been shown to record a pulse of continental inputs (Scher et al., 2011). We analyze Pb isotopes from FeMn oxide coatings and Nd isotopes from fossil fish teeth to monitor seawater values (Martin and Scher, 2004; Martin et al., 2010; Scher et al., 2011; Basak et al., in press ), which we then compare to isotopic ratios from detrital residues that represent the bulk composition of weathered continental material. The data are derived from two intermediate depth cores; one in the Atlantic sector and one in the Indian Ocean sector of the Southern Ocean. Our ultimate goal is to use these data to evaluat e the intensity of physical and chemical weathering and to determine the source rock composition (silicate vs. carbonate) of material being weathered from continental Antarctica during the EoceneOligocene transition (EOT) in order to better understand the forcing mechanisms associated with this climatic perturbation. Background Radiogenic isotopes in seawater are often used to study continental weathering and ocean circulation. High temperature and low temperature processes that fractionate Sm Nd and U ThPb parent and daughter isotopes lead to unique geochemical signatures in rocks and sediments that can be used to track provenance and evaluate lo w temperature processes. Sm Nd Systematics Sm and Nd are light rare earth elements (LREE). They are also incompatible elements, meaning they preferentially fractionate into melt phases in the mantle. Nd is
58 slightly more incompatible than Sm. As a consequence, the concentrations of both Sm and Nd increase in the melt with increasing differentiation, but the corresponding Sm/Nd ratio typically decreases with more extensive fractionation. Once a rock forms with a particular Sm/Nd ratio, the abundance of 143Nd will increase with time due to decay of 147Sm. Because Sm and Nd have comparable ionic radii and valance states, these elements do not show appreciable fractionation during weathering processes. As a consequence, the Sm/Nd ratio and Nd isotopic composition Nd) of weathering products largely preserve the parent rock composition. Thus, Nd isotopic compositions of detrital silicates have direct implications for provenance studies. Dissolved Nd in seawater is primarily derived from the continent s via weathering and riverine inputs. Nd has a short oceanic residence time of ~6001200 yrs (Jeandel, 1993; Tachikawa et al., 1997) compared to the ocean mixing time of ~1500 yrs Nd is heterogeneously distributed in the w orlds Nd signatures derived from surrounding terranes. This acquired signal can be altered along the flowpath by weathering inputs and boundary exchange (Lacan and Jeandel, 2005). While Nd isotopes in seawater predominantly record water mass mixing and advection, transient pulses of intense weathering, such as at the E/O boundary, can temporarily dominate the water mass signal (Scher et al., 2011). U ThPb Systematics Pb has three radiogenic isotopes 206Pb, 207Pb, 208Pb, which are the products of 238U, 235U, 232Th decay respectively. There is also a non radiogenic isotope, 204Pb. Of the parent isotopes, 235U has the shortest half life (0.7038 x 109), 238U has an intermediate half life (1.551 x 1010), and 232Th has the longest half life (14.010 x 109).
59 Owing to its relatively short half life, 235U has largely decayed away and currently has a low natural abundance of 0.72%. As a result, the flux of 207Pb has been negligible for the last ~ billion year s. Therefore, increases in 207Pb imply a contribution from ancient (e.g., Archean) source rocks. Because of the long 232Th half life, 208Pb has a slower growth rate than 206Pb. U, Th and Pb are all incompatible elements during mantle melting and U is sligh tly more incompatible than Pb. As a result, U, Th, and Pb concentrate in early melts and are ultimately enriched in the continental crust compared to bulk earth. Since 238U has the highest natural abundance of U isotopes (99.3%), uranogenic Pb is enriched in the crust over time. Eolian, hydrothermal, and riverine inputs are the major sources of dissolved Pb in the modern oceans (Boyle et al., 1986; Abouchami and Goldstein, 1995; Frank, 2002; Ling et al., 2005; Klemm et al., 2007). Seawater Pb has a very s hort residence time (~50200 years; Craig et al., 1973; Schaule and Patterson, 1981: Henderson and Maier Reimer, 2002) compared to the oceanic mixing time (~1500 years; Broecker and Peng, 1982), indicating that local weathering inputs probably play a great er role in its distribution than inter basin advection (van de Fleirdt et al., 2006). Thus, variations in seawater Pb isotopes, especially from sites in close proximity to continental inputs, can be interpreted in terms of continental weathering. Methods Pb isotopes for seawater were derived from Pb extracted from FeMn oxides or similar authigenic phases in marine sediments using a multi step, selective, extraction technique applied to the bulk sediment (Basak et al., in press ; also see Chapter 2). Select ive extraction steps include decarbonation using Naacetate buffered glacial acetic acid (pH=5) followed by 1.5 hrs treatment with 0.02M Hydroxylamine
60 Hydrochloride (HH) in 25% glacial acetic acid. HH treatment reduces the FeMn oxides, which preserve the bottom water Pb isotopic signature (Basak et al., in press ). This short HH step was followed by a longer 24 hr step with a fresh aliquot of HH solution to ensure complete removal of the seawater signal. Material remaining after these dissolution and extraction steps is referred to as residues, and primarily represents the carbonatefree terrigenous silicate fraction of the sediment. About 0.05 g of residue was completely dissolved in a 1:2 mixture of HF and HNO3 prior to further chemical analyses. Aliqu ots from the short HH step and digested residues were passed through cation exchange columns (for details see Chapter 2) to separate Pb and Nd from other ions. Following column chemistry Pb and Nd isotopes were analyzed on a Nu Instruments multi collector inductively coupled mass spectrometer (MC ICPMS) at the University of Florida (UF). Pb concentration measurements were made using an Element II ICP MS at UF. Concentration measurements were performed under medium resolution using Re and Rh as internal standards to correct for instrumental drift. All concentration measurements were based on calibration curves generated using a set of gravimetrically prepared solutions of commercial ICP MS standard (APEX CertiPrep, Inc.). Site 689 (6431 S) is currently located on the Maud Rise ( Figure 3 1) within the Atlantic sector of the Southern Ocean at a water depth of 2080 m (Barker et al., 1988a). The paleodepth estimate for this site puts it at an intermediate water depth of ~1638 m (Mead and Hodell, 1995). The age model used in this study is taken from Mead and Hodell (1995) and modified to the timescale of Cande and Kent (1995). Site 738 (62
61 42 S) is currently located on the southern tip of the Kerguelen Plateau ( Figure 3 1) within the Indian Ocean sector of the Southern Ocean at a water depth of 2253 m (Barron et al., 1989). An 1853 m paleodepth has been estimated for this site (Coffin, 1989). Site 738 has a well preserved EoceneOligocene boundary; however, this site is marked by an unconformity spanning from early Oligocene to late Miocene (~3310 Ma), thereby limiting the possibility of sampling sediments younger than 33 Ma. In the absence of an appropriate magnetostratigraphic record, Bohaty (2006) developed an age model for site 738 using biostratigraphic datums (Huber, 1991; Wei and Thierstein, 1991; Huber and Quillevere, 2005) and refined the model using several unique features 1813C records of bulk sediments, which were correlated to other Southern Ocean sites, especially site 689, site 690, and site 748. This site 738 age model from Bohaty (2006) was applied to this study. Results The primary goal of this study is to understand the relationship between climate and weathering on Antarctica during the development of the Antarctic cryosphere. Specifically, this study evaluates the weathering regime and the types of rock being weathered during a one million year interval across the EOT. This interval is divided into three time slices: 1) pre EOT (3434.5 Ma) at the end of the late Eocene warmth, 2) EOT (33.734 Ma) during the dramatic climate reorganization and ice sheet growth, and 3) post EOT (33.333.7 Ma) following full establishment of the ice sheet. Pre EOT ( 34 34.5 Ma) The pre EOT timeslice is characterized by similar 206Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb values for seawater and residue samples at site 689 with values of 18.80, 15.69, and 38.85 respectively ( Figure 3 2 ; Table 31 ). Site 738 also shows similar
62 seawater and residue 206Pb/204Pb ratios of 18.80; however, seawater 207Pb/204Pb and 208Pb/204Pb values of 15.67 and 38.85 at this site are slightly less radiogenic than residue values of 15.68 and 39. Both records show similar variations through time ( Figure3 Nd values for sites 689 and 738 also show agreement between seawater, as recorded by fish teeth, and residues, with both samples from both sites recording values around 7.5 ( Figure 3 3). EOT (33.7 34 Ma) The EOT timeslice covers the twostep glacial transition defined by oxygen isotopes (Coxall et al., 2005). D uring this interval 206Pb/204Pb values for seawater show a gradual transition towards more radiogenic values at both sites, while the residues become progressively less radiogenic ( Figure 3 2). In contrast, both sites record more radiogenic seawater and residue values for 207Pb/204Pb and 208Pb/204Pb, and much of 18O step (33.7 Ma) (fig 23e, f, h, i). The sample distribution in site 689 is too sparse during this interval to determine a pattern, but like sit e 738, the residue 207Pb/204Pb value is slightly more radiogenic than the seawater value; however, the residue 208Pb/204Pb at site 689 appears to be slightly Nd values become less radiogenic and start to separate during the EOT at site 738 with a greater Nd values are ~ 10 at the end of the time slice, while residue values are ~ 15. Again, most of the changes occur during the second step of the glaciation. Post EOT (33.3 33.7 Ma) Towards the end of the post EOT timeslice, seawater 206Pb/204Pb values maintain the radiogenic value of 19 for both sites. There is a bit more structure in earliest portion
63 of the post EOT in the site 738 record. Residue 206Pb/204Pb values at site 689 also maintain their less radiogenic value of 18.8, while site 738 residue 206Pb/204Pb values document a generally decreasing trend, ranging from 18.7 to 18.3 ( Figure 3 2G ). As during the EOT, residue 207Pb/204Pb val ues remain more radiogenic than contemporaneous seawater with an average offset of ~0.03 between the two sample types at site 689 compared to ~0.01 at site 738 ( Figure 3 2 E H ). Residue 208Pb/204Pb values are slightly more radiogenic than seawater values at site 689 and much of site 738, but there is more variability in the site 738 record and the final residue 208Pb/204Pb value suggests a decreasing trend ending with a value less radiogenic than seawater (F igure. 3 2 I ). Early p Nd values continue to be less radiogenic than seawater values, and both seawater and residue samples return to values similar to the early EOT by 33.3 Ma (~ 10 for residues and 8 for seawater). Discussion Source o f Radiogenic Pb During Weathering Several studies of weathering in glacial and paleosol environments demonstrate that the radiogenic Pb fraction of a parent rock is preferentially released during early weathering of fresh material, indicating incongruent weathering (Erel et al., 1994; Harlavan et al., 1998; von B lanckenburg and Nagler, 2001; Blum and Erel, 2003). Whole rock Pb isotopic values are dominated by Pb residing within feldspar 235U/204Pb]),), while the most radiogenic Pb is concentrated in accessory phases (characterized by h as well as other minerals forming from late stage fluids. Several studies suggest the radiogenic Pb in weathering solutions is derived from the accessory phases during incongruent weathering (Harlavan et al., 1998; Blum and Erel, 2003; von Blankenburg
64 and Nagler, 2001; Frank, 2002; Foster and Vance, 2006); however, common accessory phases are relatively insoluble in weak acids and rarely constitute more than a minor (~1%) fraction of the whole rock. An alternative source of labile radiogenic Pb during incongruent weathering could be interstitial oxides and cryptocrystalline aggregates that form from late stage magmatic fluids (Brown and Silver, 1955) and are enriched in U and Th relative to Pb (Rollinson, 1998) As in the case of accessory minerals, these late stage interstitial oxides and cryptocrystalline aggregates are likely to constitute a very small fraction of the Pb in the whole rock. To reevaluate the source of this radiogenic Pb during incongruent weathering, we conducted a leaching experiment on wellcharacterized silicate rocks. Three granitic rock samples from a single pluton with well constrained whole rock Pb concentrations and isotopic values were pulverized to simulate mechanically generated rock flour. These granitic rocks contain no carbonate fraction and the leaching experiment was designed to purely evaluate the behavior of the silicate fraction. Three 1gram subsamples were taken from each rock and treated with 10% acetic acid to leach out any labile fraction. Following this step, the samples were treated with 2% HNO3. One sample from each subset was exposed to the acid for 1 day, another for 3 days and the last one for 5 days. The 10% acetic acid fraction and the 1, 3, and 5 days HNO3 leac hates for each sample were then analyzed for Pb isotopic values and concentrations. Irrespective of the reagent used and the duration of reaction, the leached fractions yielded 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb values that far exceeded the whole rock ratios ( Figure 3 4). This observation is consistent with other leaching experiments on granitic glacial moraine soil chronosequences (Harlavan et al.,
65 1998). In general, high U bearing accessory phases have very low initial Pb concentrations, thus typical accessory minerals with radiogenic Pb tend to have very high initial U/Pb ratios. If the radiogenic leachates in this experiment were produced by dissolution of such phases, the isochron defined by the leachates should pass through a value very close to the origin, indicating initial 206Pb/204Pb and 207Pb/204Pb of approximately 0. Instead, the secondary isochron in f igure 3 4 passes through 206Pb/204Pb and 207Pb/204Pb values of ~15. These data indicate the acids are not attacking typical accessory phas es. Although interstitial oxides and cryptocrystalline aggregates formed from latestage magmatic fluids can contain labile radiogenic Pb, these phases are more abundant in volcanic rocks and are practically absent in the plutonic rocks used for this exper iment. The total Pb concentration in the leachates represents 0.052% of the whole rock Pb, illustrating that the labile radiogenic Pb in the silicate fraction constitutes a very small proportion of the total Pb in the whole rock. The results also indicat e that the radiogenic Pb in the leachates comes from a phase with slightly higher U/Pb ratio compared to the mean whole rock value, but still within the range of the whole rock U/Pb. Thus, a major rock forming mineral phase with a slightly elevated U/Pb ratio would be an appropriate source for the radiogenic Pb observed in the leaching experiment. In general, major rock forming minerals are not easily leached with weaker acids; however, the metamict portions of the major minerals represent one possible readily leachable source that can contain radiogenic Pb. Metamictization is a process whereby alpha emissions and spontaneous fission products from radioactive decay of U & Th gradually damage the minerals crystal structure, making that portion of the
66 mineral more susceptible to leaching by weak acids. Pb leached from these domains would be more radiogenic than the whole rock, reflecting the radioactive source, and would plot on a secondary isochron along with the whole rocks. While our leaching experiment wit h plutonic rocks suggests the metamict portions of minerals may be an import source of radiogenic Pb in the leached solution, it would be difficult to isolate this source from radiogenic Pb leached from interstitial oxides and cryptocrystalline aggregates from volcanic rocks in the natural environment. Even with the increased surface area produced by mechanical glacial activity, mass balance calculations suggest the radiogenic contribution from metamict portions of minerals (or interstitial oxides and cry ptocrystalline aggregates) would be volumetrically too small to produce the variations in seawater 206Pb/204Pb isotopes observed during the EOT. The U concentration in seawater today is ~3.2 ppb, which is two to three orders higher than seawater concentrat ions of Pb or Th (Chen et al., 1986). Even though Pb is easily complexed with CO3 2and Cl-, its concentration diminishes rapidly due to scavenging processes that transport it to the sediment. Thorium is known to be highly insoluble and is easily removed f rom seawater through adsorption on suspended particles. Because of these distinct geochemical behaviors, seawater is highly enriched in uranium compared to thorium (U/Pb>> Th/Pb) (Jahn and Cuvellier, 1994). Carbonates formed from seawater reflect these U/P b and Th/Pb ratios. Over time these carbonates become a source of easily weatherable uranogenic Pb that is largely decoupled from thorogenic Pb. Therefore, weathering of early Paleozoic or older marine carbonate provides an alternative explanation for radi ogenic 206Pb/204Pb observed during early stages of glacial weathering that is not accompanied by
67 radiogenic 208Pb/204Pb or 207Pb/204Pb, which is limited by the low abundance of 235U in Phanerozoic seawater. Weathering at the EOT The type of weathering (mechanical vs. chemical) and the substrate that is being weathered (silicate vs. carbonate) controls the longterm drawdown of CO2. It is widely accepted that atmospheric CO2 dropped below a critical threshold as a precondition for Antarctic glaciation during the EOT (DeConto and Pollard, 2003; Pagani et al., 2005). Here, we use Pb isotopes in seawater and residues to evaluate the intensity of chemical weathering and source rock composition (silicate vs. carbonate) across the EOT. Mechanical vs. chemical weat hering The similarity between seawater and residue Pb isotopic values at the end of the Eocene, particularly for 206Pb/204Pb, is interpreted to indicate extensive transfer of the whole rock isotopic signature to seawater, suggesting extensive chemical weathering on Antarctica during this warm interval. A strong chemical weathering regime is also supported by studies of clay minerals from circum Antarctic sediments during the late Eocene (Roberts and Kennett, 1997). Immediately following the EOT increases i n both seawater and residue 207Pb/204Pb and 208Pb/204Pb ( Figure 3 18O record (Coxall et al., 2005). Synchronous changes in ice growth and residue values suggest increased mechanical weathering during the EOT, which introduced fresh rock and grains with large surface areas. Although the change in residue isotopic values can be explained by introduction of this fresh terrigenous material from various sources via mechanical weathering, associated change s in seawater isotopes require active chemical weathering as well. In this scenario, glacial mechanical weathering increased
68 the surface area/mass ratio of weathered material, which would have facilitated subsequent chemical reactions. Since the initial EO T glaciers on continental Antarctica were mostly wet based (Zachos et al., 1999), it is possible that a sufficient volume of water was available to maintain the chemical reactions needed to weather the radiogenic Pb fraction from these freshly exposed surf aces. Global cooling associated with glacial conditions would also have increased the CO2 solubility, producing stronger carbonic acid and further facilitating chemical reactions. The pattern of a gradual shift in Southern Ocean seawater towards more radiogenic 206Pb/204Pb values and the separation between seawater and residue values at the EOT ( Figure 3 2) is distinct from observed 207Pb/204Pb and 208Pb/204Pb patterns in which the isotopes for both seawater and residues rise gradually. Although, the more radiogenic 206Pb/204Pb in seawater also suggests active chemical weathering, the source rock being weathered must have been fundamentally different leading to the observed decoupling between seawater and residue Pb isotopes. Thus, the weathering regime following the development of ice on Antarctica included both mechanical and chemical weathering. Foster and Vance (2006) also argued for active chemical weathering for glaciated domain based on a Pleistocene study using Pb isotopes. Silicate vs. carbonate weathering The radiogenic Pb contribution from metamict minerals or cryptocrystalline aggregates during silicate weathering is probably volumetrically insufficient to generate the inc rease in seawater Pb observed in the Southern Ocean during the EOT. While it is difficult to estimate the total flux of radiogenic Pb from metamicts during E/O weathering, observations from the leaching experiment indicate that radiogenic Pb contributed fr om metamict minerals is less than 2% of the whole rock Pb. Therefore it is
69 reasonable to assume that metamict minerals do not play a major role in the composition of E/O weathering solutions. Thus, chemical and mechanical weathering of silicates will defin e the Pb isotopes of seawater and residues during this transition. In contrast, seawater 206Pb/204Pb values become more radiogenic while contemporaneous residue 206Pb/204Pb values become less radiogenic during the EOT and post EOT. So, in addition to the c ontributions from silicate weathering apparent in the 207Pb/204Pb and 208Pb/204Pb records, the 206Pb/204Pb data require an additional source that can account for the greater magnitude of increase in seawater 206Pb/204Pb, but had little impact on 207Pb/204P b and 208Pb/204Pb. One possible candidate is a change of source to weathering of old carbonate rocks. Due to high initial U/Pb these old carbonates would contribute radiogenic 206Pb/204Pb to seawater, but are unlikely to affect the 207Pb/204Pb ratios, because they are too young relative to the time period of active 235U decay, or the 208Pb/204Pb ratios, because of low initial Th/Pb in carbonates. The carbonates are also likely to be more susceptible to glacial mechanical and chemical weathering than silicat e basement rocks. Carbonate rocks represent 510% of the total exposed basement rock volume of the east Antarctic (Elliot, 1975). During the early Cambrian, epicontinental seas formed along the western boundary of the East Antarctic shield, leading to extensive shallow mari ne carbonate deposition under warmer conditions. Isolated outcrops of Shackleton Limestone, an ~500 Ma old carbonate sequence along the southwestern Antarctic shoreline near the Ross Sea, represent important evidence for the Cambrian platform (Elliot, 1975; Rowell and Rees, 1987; Goodge et al., 2002). Model data support extensive glacial development in the Ross Sea sector during the EOT (DeConto and
70 Pollard, 2003), providing a possible mechanism to initiate weathering of this old carbonate. Residue Pb Isotopes Pb isotopes from the residues can be used for provenance studies. Residue 206Pb/204Pb values in the Southern Ocean decrease while 207Pb/204Pb and 208Pb/204Pb values increase after the E/O boundary at both sites 689 and 738 ( Figure 3 3). The increase in 207Pb/204Pb typically indicates an ancient source rock. Detailed observations from the higher resolution records at site 738 show small positive excursions in 207Pb/204Pb and 208Pb/204Pb at ~33.9 Ma and larger positive excursions at ~33.7 Ma during the second step of the E/O glaciation. Smaller negative excursions are also detected in residue 206Pb/204Pb. The timing of these variations implies a relationship between the two step Antarctic glaciation (Coxall.et al., 2005; Lear et al., 2008) and the weathering response (Scher et al., 2011). In particular, the rapid Pb excursions observed during the second step of the E/O glaciation are consistent with the idea that intensified weathering of crustal rocks accompanied the intense interval of glacial growth. Nd values in sites 689 and 738 show ( Figure 3 3) similar values suggesting a well mixed seawater signal and efficient mixing of detrital inputs following onset of the proto ACC in the Eocene (Scher and Mart in, 2006, 2008). At the beginning of the EOT, both seawater and residue Nd isotopes at site 738 start to become less radiogenic reaching minimum values during the second step of the E/O glaciation. Rapid recovery to values similar to pre EOT values is also observed during the post EOT time window ( Figure 3 3). These observations support the interpretation deduced from Pb isotopic compositions of residues that intense glaciation increased the
71 weathering inputs. Introduction of this weathered material into seawater overwhelmed the advective signal for both the short residence time Pb and intermediate residence time Nd, producing rapid excursions in seawater isotopes for both systems. Pre EOT residue 207Pb/204Nd values plot within a limited range. The compositi on of the EOT and Post EOT continental input is defined by variable 207Pb/204Nd values in 207Pb/204Nd space. ( Figure 3 5). This overall shift suggests a source change, possibly indicating introduction of multiple new sources as the glaciers grew and weathered older silicate basement rocks from new locations. Weathering and Paleoceanogrpahic Significance Pb isotopes in seawater and residues provide important information about the mechanisms of weathering and the types of rock being w eathered during the EOT. Seawater Pb isotopes highlight the impact of glacial mechanical weathering combined with active chemical weathering associated with wet based glaciers on continental Antarctica. Observed seawater and residue Pb isotopes are best ex plained by the combination of carbonate and silicate weathering, with silicate weathering dominating both seawater and residue 207Pb/204Pb and 208Pb/204Pb signals, while carbonate weathering modulated the seawater 206Pb/204Pb isotopic record. Chemical weat hering of silicate and carbonate rocks can sequester atmospheric carbon over the long and short term respectively. Therefore, evidence for chemical weathering of silicates and carbonates from Antarctica supports the concept of a drawdown of atmospheric CO2 over the EOT based on boron isotopes (Pearson et al., 2009) and presents a possible mechanism. In addition, Pb isotopes from residues illustrate a change in the intensity of weathering and the composition (provenance) of weathered material as the East Ant arctic ice sheet grew. In particular, Nd and Pb isotopic data from seawater and
72 residues illustrate a pulse in weathering associated with each of the two steps of the EOT that have been attributed to rapid cooling and ice growth (Coxall et al. 2005; Lear e t al., 2008; Katz et al., 2008). Enhanced silicate weathering at the E/O would also deliver additional silicate to the ocean needed to support the observed increase in opal accumulation in the Southern Ocean during this transition. The carbonates weathered during the EOT are most likely old (Late Cambrian) and are believed to have formed in a shallow marine or neritic environment. Neritic 13C compared to pelagic carbonates due to high productivity rates in this shallow, margi nal environment (Stewart and Eberli, 2005). Thus 13C values observed during the EOT ( Figure 3 6) (Merico et al., 2008). Moreover, the associated influx of Ca2+ into the ocean c ould be a factor in the documented deepening of the CCD by ~ 1 km. ( Figure 3 6), although inputs associated with global lowering of sea level and weathering of exposed shelf carbonates also probably contributed. This interpretation is consistent with a modeling study by Merico et al. (2008) that suggested carbonate weathering as a source of the carbon isotope and CCD fluctuations. Summary Pb isotopes from samples representing seawater and terrigenous detrital fractions from deep sea sediments document changes in the weathering regime, intensity and source rocks during the transition from the late Eocene greenhouse into the Oligocene ice house. Similar 206Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb values between the two fractions during the late Eocene highlight an interval of intensive chemical weathering on Antarctica during a time of warm temperatures and possible weathering of well weathered s equences. Increased seawater 206Pb/204Pb and
73 decreased detrital 206Pb/204Pb during the EOT argue for incongruent weathering of freshly exposed silicate rocks that released labile, radiogenic 206Pb to seawater. Previous studies suggest this radiogenic Pb c ould be sourced from accessory minerals or interstitial oxides during incongruent weathering of silicates. A leaching experiment using isotopically well defined granitic rocks suggests the radiogenic 206Pb tapped a source with a relatively high U/Pb ratio, but a ratio that was within the range of the whole rock values and produced a secondary isochron. These data suggest the radiogenic Pb was contributed from metamict portions of major minerals in the silicate fraction. The overall Pb flux from this source probably represents a very small fraction of the whole rock Pb and would be insufficient to drive large changes the seawater Pb isotopic composition. A more abundant input that could explain the different responses observed in seawater and detrital 206Pb/204Pb is weathering of Paleozoic carbonate rocks that were initially enriched in U and depleted in Th. This source also explains the observed increase in seawater 206Pb/204Pb compared to the minimal overall change in seawater 207Pb/204Pb and 208Pb/204Pb acr oss the EOT, as well as the similarities in the values and trends of seawater and residue 207Pb/204Pb and 208Pb/204Pb. Chemical weathering of old carbonates is also consistent with the observed deepening of the CCD and positive 13C shift across the EOT. Finally, increasing residue 207Pb/204Pb isotopes during the EOT, particularly in association with the second step of the E/O glaciation, also indicates intensification and expansion of silicate weathering to new regions of silicate basement rocks, which cou ld have contributed to the drawdown of atmospheric CO2.
74 Table 31. Average Pb isotopic values in seawater and residues in PreEOT, EOT, and Post EOT intervals in Southern Ocean sites 689 and 738. 206 Pb/ 204 Pb 207 Pb/ 204 Pb 208 Pb/ 204 Pb Site 689 Site 738 Site 689 Site 738 Site 689 Site 738 Pre EOT Seawater 18.8 18.8 15.69 15.67 38.85 38.85 Residue 18.8 18.8 15.69 15.68 38.85 39.0 EOT Seawater 19.0 18.8 15.7 15.69 39.0 39.0 Residue 18.8 18.8 15.72 15.69 39.0 39.0 Post EOT Seawater 19.0 19.0 15.7 15.69 39.1 39.0 Residue 18.8 18.6 15.73 15.69 39.1 39.0 Pb isotopes for Pre EOT, EOT, and Post EOT are represented by range of Pb isotopic values. Numbers in the table represent approximate mean of these range of values.
75 Figure 31. Site location showing ODP site 689 (2080 m) from the Atlantic sector and site 738 (2253 m) from the Indian sector of the Southern Ocean.
76 Figure 32. A plot of 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios versus age in seawater and detrital residues from site 689 and site 738 from the late Eocene across the Eocene/Oligocene Transition (EOT). (ac) Records of 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios from 23 40 Ma for site 689. (d f) Records of 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204P b from 33.3 34.5 Ma for site 689 l. (gi) Records of 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios from 33.3206Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb measurements are 0.004, 0.003 and 0.008 respectively)
77 Figure 33. Nd values versus age of fossil fish teeth, representing seawater, and detrital fractions for 33.3 34.5 Ma. (a,b) Nd record of seawater and terrigenous material for sites 689 and 738 respectively. Figure 34. A plot of 20 6Pb/204Pb vs. 207Pb/204Pb from a leaching experiment. A n acetic leach, a 1 day leach in 2% HNO3, a 2 day leach in 2% HNO3, a 3 day leach in 2% HNO3, and the whole rock for 3 pulverized granite samples. The black arrows indicate the whole rock Pb isotopic values for each of the three granite samples.
78 Figure 35 A plot of 207Pb/204Nd for the detrital fractions. Green circles represent pre EOT samples, black squares represent EOT samples, and blue triangles represent post EOT samples.
79 Figure 36 Plot showing the different proxy response to Oligocene glaciation in Antarctica ~ 34 Ma. (a 18O, %CaCO3 13C data from equatorial Pacific site 1218 from Coxall et al. (2005). (d g ) 18O, Nd, 207Pb/204Pb, and 206Pb/204Pb from sit e 738. The grey bars indicate two steps of the EOT.
80 CHAPTER 4 SOUTHERN S OURCE OF DEGLACIAL O LD C ARBON IN THE TR OPICAL NORTH PACIFIC BASED ON NEODYMIUM I SOTOPES A 14C (by 190) and rise in CO2 during the Mystery Interval (17.5 to 14.5 kyr BP) are believed to have been caused by the release of 14C depleted CO2 from the ocean to the atmosphere (Broecker and Barker, 2007). The Southern Ocean has been identified as a potential source for this old CO2 (Marchitto et al., 2007) based on the idea that extensive sea ice cover (Stephens and Keeling, 2000) and/or upper ocean stratification (Francois et al., 1997; Sigman et al., 2000) in southern latitudes during the Last Glacial Maximum (LGM) isolated the underlying abyssal water mass from air sea exchange, resulting in a 14C depleted reservoir. Marchitto et al. (2007) 14C excursion in intermediate waters from the eastern North Pacific during the last deglaciation, which they attributed to injection of 14C depleted Antarctic Intermediate Water (AAIW). 14C findings clearly tag an old water mass, the source of the water remains controversial. 14C, which reflects the age or ventilation rate of a water mass, Nd isotopes can be used to track the source of the water mass; hence, to resolve this controversy I present Nd isotopes from fossil fish teeth/debris collected from the same core studied by Marchitto et al. (2007). My data exhibit a clear shi ft toward 14C excursion and Southern Hemisphere climatic signals, further supporting an AAIW source. Negative radiocarbon excursions (Marchitto et al., 2007; Stott et al., 2009) observed in the eastern tropical Pacific are contemporaneous with increased upwelling in the Southern Ocean (Anderson et al., 2009) and onset of warming and rising CO2
81 detected in Antarctic ice cores (Monnin et al. 2001) highlighting a possible cause and effect relationship between circulation in the Southern Hemisphere and radiocarbon in the eastern tropical Pacific. Despite these observations, there is no direct evidence that this older water was sourced from the Southern Ocean. An alternative possible source is the high latitude North Pacific. There is no clear way to distinguish between northern 141318O because similar oceanic and atmospheric processes could theoreticall y act to modify the isotopic characteristics of either endmember water mass. In contrast, Nd isotopes in fossil fish teeth/debris are considered quasiconservative tracers of water mass, meaning that major water masses carry distinct Nd isotopic ratios (reported here as Nd) that are generally only altered by mixing, although modification from local weathering inputs is possi ble Background Nd Isotopes Systematics Neodymium has a residence time in the ocean (~6001200 yrs) (Jeandel, 1993; Tachikawa et al., 1997) that is shorter than the mixing time of the global conveyer (~1500 yrs) (Broecker and Peng, 1982) and it is sourced primarily from the continents (Goldstein and Jacobsen, 1987, Elderfield, 1988; Jeandel, 1993; Tachikawa et al., 1999, 2003) As a res ult different ocean basins have different Nd values. For example, North Atlantic Deep Water (NADW) has a value close to 13.5 ( Piepgras and Wasserburg, 1987; Lacan and Jeandel, 2005) which reflects weathering of Precambrian shield material, while Pacific Deep Water has a value between 4 to 6 (Piepgras and Jacobsen, 1988) reflecting deep sea mixing and weathering of young volcanic ash. Major water masses carry their initial Nd isotopic signal as they circulate through the ocean; however, that signal can be altered along the route by weathering inputs and
82 boundary exchange (Lacan and Jeandel, 2005) particularly in marginal marine settings. As a result, Nd isotopes are often referred to as quasi conservative tracers of water masses. Owing to the large differences in Nd values between the major water masses in the world oceans, minor modifications of endmember values generally do not erase the differences that distinguish water masses. Hydrography Under modern hydrographic conditions AAIW is a low salin ity, high oxygen intermediate water mass. This intermediate water occupies depths of 500 to 1000 meters below sea level (mbsl) and over 50% of global AAIW is located in the Pacific Ocean. In general, northern flow of AAIW in the southern Pacific occurs along the eastern margin of the basin (Tomczak and Godfrey, 1994) with return flow southward along the western boundary. Near the Coral Sea some of the water from this gyre is deflected equatorward and passes through the Papua New Guinea (PNG) region before turning east near the equator and ultimately feeding into the E quatorial Intermediate water. At its source region in the Southern Ocean modern AAIW Nd values range between 6 to 8, representing a mixture of Atlantic and Pacific waters. Currently the intermediate water found along the equatorial Pacific is similar t o AAIW with respect to its hydrological properties, while the Nd s ignature is strongly modified. The modified AAIW in the equatorial region has an average Nd value of ~ 2.8 (Lacan and Jeandel, 2001). This dramatic increase in the Nd value of AAIW occur s through inputs from easily weathered volcanic rocks in the tropical Pacific (Lacan and Jeandel, 2001) The young volcanic rocks of PNG have an Nd value of ~+7, which is highly distinct from seawater values and thus small inputs alter AAIW Nd values to ~ 2.8 (Lacan and
83 Jenadel, 2001). M y core site is located at 23N within a shadow zone at the confluence of EqIW and NPIW. EqIW is a mixture of AAIW modified in the PNG region and upwelled PDW. There are no previously published water column or core top Nd data for EqIW; however, limited data for PDW from the deep tropical waters indicate very little variation in Nd since the LGM (Marchitto et al., 2005) Thus, variations in Nd at the Baja site are expected to primarily reflect the interaction between NPI W and modified AAIW. NPIW is a mixture of waters from the Kuroshio Current, intermediate water ventilated in the Okhotsk Sea (OSIW) and intermediate water ventilated in the western North Pacific (WNP) (You, 2003; Talley, 1999; Amakawa et al., 2004a) End m ember Nd values for these waters measured prior to ventilation and mixing are ~ 5.6 for the Kuroshio Current, 3.6 for the Okhotsk Sea, and 0.1 for the WNP (Amakawa et al., 2004b; Piepgras and Jacobsen, 1988) These three waters mix at depth to form NPI W. Varying percentages of these source waters along the flow path of NPIW account for the high spatial variability in Nd values reported within the upper 1 km of the NW Pacific water column ( 4 to 0) (Piepgras and Jacobsen, 1988; Amakawa et al., 2004a, b; Amakawa et al., 2009) Previous studies suggest that during the LGM, NPIW (Glacial NPIW) was well ventilated (Keigwin, 1998; Matsumoto et al., 2002) with probable contributions from the Bering Sea, Okhotsk Sea and/or Gulf of Alaska (Horikawa et al., 2010) introducing additional radiogenic Nd. This is consistent with my observation that NPIW at Baja has a radiogenic value of ~ 1 during the LGM. Even more pronounced intermediate and deep water production rates have been observed in the WNP during HS1 according to a drop in ventilation age of ~600 years between LGM and HS1 based on the 14C age
84 difference between benthic and planktonic foraminifera ( Okazaki et al., 2010) Increased contributions of radiogenic WNP intermediate water ( Nd ~ 0.1) to NPIW would result in an NPIW Nd value that was similar to, or more radiogenic than, LGM values. Therefore, anticipated changes in the endmember composition of NPIW during HS1 are expected to produce inc reasing Nd values at the Baja site, in direct contradiction to the observed decreasing values at the onset of HS1. Moreover, modeli ng studies by Okazaki et al. (2010) predict this young LGM/HS1 NPIW flowed out of the Pacific Ocean as a western boundary current, leaving southern sourced waters as the primary component in the eastern Pacific. Core Location and Method Our core location and depth is currently situated at the boundary between Equatorial Intermediate Water (EqIW) (a mixture of AAIW and upwelled Pacific Deep Water (PDW)) (Bostock et al., 2010) and North Pacific Intermediate Water (NPIW) within the east Pacific shadow zone ( Figure 4 1). Limited data for PDW from the deep tropical waters indicate very little variation in Nd since the LGM (Marchitto et al., 2005) Thus, the Nd value at Baja is largely controlled by the relative contributions of NPIW (higher values) and AAIW ( lower values), with AAIW values being additionally modified during transit from the Southern Ocean through the Papua New Guinea region. Fossil fish teeth/debris were handpicked from the >63 m fraction of sieved sediments from core MV99MC19/GC31/PC08 off Baja California. Picked fossil fish teeth/debris was dissolved in a 1:1 mixture of optima grade HNO3 and HCl in preparation for column chemistry. A primary column with Mitsubishi resin and 1.6N HCl as an eluent was used to separat e bulk REEs. Nd was then isolated by passing the
85 REE aliquot through Lnspec resin with 0.25 N HNO3 as an eluent. Procedural blanks were 14 pg Nd. Nd isotopes were analyzed on a Nu Plasma Multi Collector Inductively Coupled Plasma Mass Spectrometer (MC ICPMS) at the University of Florida. Nd aliquots from column chemistry were dried and redissolved in 2% NHO3 prior to aspirati on using a DSN100 nebulizer. Preamplifier gain calibrations were run at the beginning of each analytical session. All Nd isotope data reported in this study were analyzed using a time resolved analysis (TRA) method (Kamenov et al., 2008) Prior to sample introduction a baseline was m easured for 30 seconds. Data were acquired in a series of 0.2 seconds integrations over 13 minutes. All reported 143Nd/144Nd ratios were corrected for mass fractionation using 146Nd/144Nd =0.7219. International standard JNdi 1 was analyzed between every 56 unknown samples and the average of these standard runs were compared to a long term TIMS JNdi 1 value of 0.512103 +/ 0.000014 to determine a correction factor for each of the samples analyzed on that day. The long term 2 e xternal reproducibility of JNdi 1 analyses on the Nu of 0.000014 (0.3 Nd units) has been applied to all samples unless the internal error was larger. Results I report a Nd isotope time series for the last 38 kyr (Table 41 ) from core MV99 MC19/GC31/PC0 8 near southern Baja California using the same samples and age model as Marchitto et al. (2007) in order to distinguish between potential northern and southern sources for the old, 14C depleted water detected during the last deglaciation. M y record illustr ates variations in Nd values from 0.3 to 3.0. Pre LGM (before 21 kyr BP) Nd values range between 1 to 2.5, with values around 1 during the LGM. A prominent negative excursion to approximately 3.0 occurs at the beginning of the
86 deglaciation, with a subsequent increase to 1.7 during the later half of Heinrich stadial 1 (HS1). Nd isotopic values continue to increase into the Blling Allerd (BA) warm interval, where they plateau at ~ 1.4. There is limited data during the cold Younger Dryas (YD) interval; however, there is a hint of a minor negative Nd excursion during the Younger Dryas (YD) followed by a positive excursion and values ranging from 0.3 to 1.6 in the Holocene. There is some structure in the Holocene record, but the mean Nd value is comparable to LGM values and tends to be higher than preLGM values. Discussion In general, modern AAIW Nd values are ~ 6 to 8 Nd units and NPIW values are ~ 0 to 4 Nd units ( Goldstein and Hemming, 2003). Thus, the overall 0.3 to 3.0 range in Nd values observed in our data would seem to represent waters dominantly sourced from the North Pacific throughout the last 38 kyr. However, the Nd isotope signature of southernsourced AAIW is modified during its transit through the Pacific in the region of Papua New Guinea (PNG) due to interactions with volcanic inputs that introduce much higher Nd isotopic ratios ( Nd ~ +7) (lacan and Jeandel, 2001) As a result, the measured Nd value of this modified AAIW is ~ 2.8 ( Lacan and Jeandel, 2001) and it is this water that flows eastward along the equator to influence EqIW. In addition, our youngest Nd value of 1.5 may be taken as a minimum value for modern NPIW in this region based on the most conservative scenario that the site is currently bathed by pure NPIW. In the absence of any highresolution Nd reconstructions from the AAIW and NPIW source regions, we begin with the as sumption that the endmember compositions of these two water masses have remained relatively constant since the LGM. In the context of constant endmembers, Nd excursions to more radiogenic values represent an increased flux of NPIW (or less AAIW), while excursions towards less radiogenic
87 values record an increased flux of AAIW (or less NPIW). Following this logic, the preLGM period (before 21 kyr BP) in our record indicates a mixture of NPIW and AAIW, with a relatively decreased component of AAIW during the LGM, similar to Holocene values ( Figure 4 2). AAIW intensified during the deglaciation, comprising most of the water mass during the early part of the HS1 event, then decreased during the B A warming period and might have increased slightly again during the YD. Post YD the region was once again dominated by NPIW. The timing of major fluctuations in AAIW interpreted from Nd isotopes is broadly in agreement with the idea of depleted 14C orig inating in the Southern Ocean ( Figure 4 3B, C). The deglacial d ecrease in Nd values around 18 kyr BP likely represents the first pulse of AAIW into the eastern North Pacific as the Southern Ocean recovered from surface stratification (Francois et al., 1997; Sigman and Boyle, 2000) and/or extensive seaice coverage (S tephens and Keeling, 2000) during the LGM. The timing of this 13C excursion of thermocline dwelling foraminifera Neogloboquadrina dutertrei in the EEP (Spero and Lea, 2002; Pena et al., 2008b ) also in terpreted to indicate an influx of older water from the Southern Ocean ( Figure 4 3D ). In the Southern Ocean higher productivity indicated by an increased opal flux near Antarctica also supports increased upwelling of nutrient rich abyssal water and hence more AAIW activity (Anderson et al., 2009) at this time ( Figure 4 3E ). Evidence of increased deglacial AAIW has also been reported from Nd isotopic records in the Atlantic (Pahnke et al., 2008) Comparison of Northern (Grootes and Stuiver, 1997) and Sout hern Hemisphere (Monnin et al., 2001) ice core temperatures indicate that deglacial warming in the
88 Southern Ocean preceded deglaciation in the high latitude Northern Hemisphere ( Figure 4 3A ). Early warming in the Southern Hemisphere and invigorated AAIW fo rmation are well know (Marchitto et al., 2007; Toggweil er et al., 2006; Stott et al., 2009) In contrast, the deep and shallow Northern Pacific did not start responding to deglaciation until the beginning of the B A period (Galbraith et al., 2007) (~ 14.8 kyr BP), although localized intermediate water formation during HS1 in the W estern North Pacific (WNP) has recently been reported (Okazaki et al, 2010) Therefore, temporal differences in the deglacial history between the Northern and Southern Hemispheres also imply that the Nd isotopic excursion at ~18 kyr was controlled by intermediate water mass readjustment in the Southern Hemisphere. The deep ocean represents the most likely storage area for old carbon during the last glaciation due to the proposed mechanism of formation and the volume of the old carbon reservoir (Broecker and Barker, 2007; Broecker, 2009) Our Nd data imply that the old carbon reported from the Baja site (Marchitto et al., 2007) was sourced from the Southern Ocean via AAIW that was m odified in the equatorial Pacific. A recent study indicates (De PolHolz et al., 2010) no evidence for deglacial old 14C in intermediate waters off Chile, close to one of the main regions of AAIW formation today. P art of that AAIW takes a complicated route in the South Pacific before re aching the Baja site. Another portion of this AAIW passes from the Southeastern Pacific, through the Drake Passage (Talley, 1999) and circumnavigates Antarctica where it could have picked up old carbon from the deep southern abyss before circulating back into the Indian and Pacific Oceans. The Chilean site (De PolHolz et al., 2010) may be too close to the region of active AAIW formation to have been influenced by AAIW that acquired older
89 carbon during its journey around Antarctica. Recent evidence of radiocarbon depleted AAIW reported from tropical South Atlantic off Brazil during the HS1 and YD intervals (Mangini et al., 2010) and from the deep Southern Ocean (Skinner et al., 2010) during the last glacial period are consistent with acquisition of old abyssal carbon east of the Drake Passage. Half way through the HS1 event Nd values in the Baja record become more 1414C r ecord depicts a strong influence of aged water until the beginn ing of B A period ( Figure 4 3C ), whereas the Nd record illustrates increasing values during the later half of HS1, which could be interpreted as decreasing AAIW. However, no other lines of evi dence support diminished AAIW at this time (Anderson et al., 2009; Monnin et al., 2001; Spero and Lea, 2002) ( Figure 4 3). Alternatively, an increased NPIW flux could potentially explain the observed discrepancy during the mid H S 1. Although the WNP is known to be a region of active intermediatedeep water formation during H S 1 these waters were relatively enriched in 14C and do not appear to have reached the e astern Pacific (Okazaki et al., 2010) In the absence of any compelling evidence in favor of diminished AAIW or increased NPIW we must appeal to modified endmember Nd values of AAIW or NPIW to explain the 14C and Nd. Given that Nd values suggest the water at the Baja site is dominated by AAIW by mid HS1, we focus on this endmember. Changes in weathering in the PNG region, possibly due to a southward shift of the Intertropical Convergence Zone (Krebs and Timmerman, 2007) and increased rainfall around PNG during deglaciation, could have introduced more positive PNG type Nd values to AAIW.
90 Changes in Nd during the B A and YD are more subtle but appear to be consistent with anticipated variations in NPIW and AAIW. During B A opal flux data from the Southern Ocean indicates reduced nutrient upwelling and AAIW formation (Andersen et al., 2009) A relative increase in the proportion of NPIW versus AAIW during the B A period is consistent with the small increase observed in Nd at the Baja site. A minor negative Nd excursion (albeit a onepoint low) during the YD is consistent with 14C and other indicators of increased AAIW from the Southern Ocean (figs. 4 3b, 4 3c). Post YD values are comparable to LGM values, reflecting increased influenc e of NPIW during the Holocene. This scenario is c onsistent with rearrangements in Holocene deep and intermediate circulation patterns leading to a modern system that is dominated by NPIW at this site. Summary Evidence of old carbon at intermediate depths in the eastern North Pacific has been one of the most intriguing paleoceanographic findings establishing the mechanism of atmospheric CO2 leakage at the end of the last glaciation. Direct and indirect evidence in favor of an old carbon reservoir and increased AAIW have been reported from the Souther n Oc ean (Spero and Lea, 2002; Anderson et al., 2009; Skinner et al., 2010), but there is no direct proof that these Southern waters reached the tropical North Pacific during deglaciation. This coupled study of Nd 14C reports variations in the relative do minance of more radiogenic NPIW and less radiogenic AAIW over the past 38 ky and supports the concept of an increased contribution of AAIW at the beginning of the deglaciation, suggesting that the old carbon was transported from the deep southern abyss to the shallow tropics via AAIW.
91 Table 41 Nd isotopes from fossil fish teeth/derbis of core MV99MC19/GC31/PC08 Sample Composite Calendar Nd error depth (m) Age (kyr BP) (+/ ) MC19, 10 11cm 0.100 0.32 1.4 0.4 MC19, 25 26cm 0.250 0.81 0.7 0.2 MC19, 40 41cm 0.400 1.29 0.9 0.2 GC31 3, 22 23cm 0.475 1.54 0.9 0.2 GC 31 3, 50 51cm 0.755 2.44 0.3 0.2 GC31 3, 70 71 cm 0.950 3.08 0.4 0.3 GC31 3, 100 101 cm 1.255 4.07 1.6 0.2 GC 31 3, 123 124 cm 1.480 4.79 0.8 0.2 GC31 2, 0 1cm 1.755 5.69 1.1 0.2 GC31 2, 50 51cm 2.255 7.31 1.0 0.1 GC31 2, 100 101cm 2.755 9.01 1.3 0.2 GC31 1, 8 9 cm 3.255 10.79 0.8 0.2 PC08 9, 60cm 3.410 11.32 0.3 0.5 PC08 9, 85cm 3.660 12.12 1.7 0.2 PC08 9, 110cm 3.910 12.88 1.4 0.2 PC08 9, 130cm 4.110 13.57 1.3 0.2 PC08 9, 140 142cm 4.220 13.94 1.5 0.3 PC08 9, 140 142cm 4.220 13.94 1.7 0.3 PC08 9, 150 cm 4.310 14.2 2.0 0.3 PC08 8, 16 17cm 4.475 14.65 1.8 0.2 PC08 8, 25 26 cm 4.565 14.96 1.7 0.2 PC08 8, 35 cm 4.660 15.3 1.7 0.1 PC08 8, 45 46cm 4.765 15.67 2.0 0.2 PC08 8, 65 66 cm 4.965 16.37 3.0 0.3 PC08 8,105 106 5.365 17.78 1.9 0.1 PC08 8,115 117cm 5.470 18.12 1.3 0.1 PC08 8,125cm 5.560 18.41 0.9 0.2 PC08 7,60 61cm 6.415 21.14 1.0 0.1 PC08 7,80 81cm 6.615 21.77 1.9 0.2 PC08 7,132 133cm 7.135 23.43 1.5 0.2 PC08 6,20 21 7.585 25.15 1.8 0.2 PC08 5,25cm 9.010 30.7 2.2 0.2 PC08 5,130cm 10.060 34.55 2.3 0.2 PC08 4,15cm 10.410 35.89 1.8 0.2 PC08 4,65cm 10.910 37.92 0.9 0.2 Nd =(( 143 Nd/ 144 Nd) measured /( 143 Nd/ 144 Nd) CHUR ) 1) 10 4 All reported 143 Nd/ 144 Nd ratios were corrected for mass fractionation using 146Nd/144Nd =0.7219. International standard JNdi 1 was analyzed between every 5 6 unknown samples and the average of these standard runs were compared to a long term TIMS JNdi 1 value of 0.512103 +/ 0.000014 to determine a correction factor for each of the samples analyzed on that day. Long term external reproducibility of JNdi 1 analyses on the Nu is 0.000014 (0.3 Nd units) Error bars for a ll figures in the text represent external reproducibility unless the internal error is larger.
92 Figure 41. Modern hydrography. Depth profile of annual mean salinity along the continental margin of the eastern North Pacific. Cores MV99 MC19/GC31/PC08 (23.5oN 111.6oW ) and VM2130 (1oS, Eastern Equatorial Pacific (EEP ) ) are bathed by EqIW which is largely a combination of NPIW AAIW and PDW AAIW: Antarctic Intermediate Water; CC: California Current; CU: California Undercurrent; EqIW: Equatorial Intermediate Water; EUC: Equatorial Undercurrent; NPIW: North Pacific Intermediate Water; PDW: Pacific Deep Water
93 Figure 42. Nd record from Baja California for the last 38kyr BP. Nd of fossil fish teeth/debris plotted against calendar age. [ Nd= ( ( (143Nd/144Nd)measured/(143Nd/144Nd)CHUR) 1) 104] Vertical error bars represent long term external reproducibility of the JNdi 1 standard (0.3 Nd units) except for samples with a higher internal error. The yellow, blue, and pink shaded areas indicate Heinrich Stadial (HS1), B lling Aller d (B A), and Younger Dryas (YD) climate episodes. The period between 14.5 and 17.5 kyr BP has been referred to as the Mystery Interval.
94 Figure 43. Meridional multi proxy responses to the last deglaciation (8 to 22kyr BP). (a) Greenland ice cores temperatures, (b) Nd from MV99 MC19/GC31/PC08, 14C from MV99 13C of thermocline dwelling foraminifera Neogloboquadrina dutertrei in the Eastern Equatorial Pacific (EEP), (e) opal flux data from the Antarctic Polar Front, (f) atmospheric CO2 ) from Antarctic Dome C plac ed on the GISP 2 time scale. The blue shaded area highlights the initial deglaciation.
95 CHAPTER 5 CONCLUSIONS The over arching goals of this dissertation have been to better understand the climatic response to changing weathering regimes and reorganization of ocean circulations during two climatically important threshold events; namely the onset of large Cenozoic ice sheets at the EoceneOligocene boundary (~34 Ma) and the end of the last major Northern Hemisphere glaciation during the last deglacial (~18 kry) ago. In order to achieve these goals the research design included developing a geochemical proxy for weathering, application of the newly established proxy to develop time series records, and application of an existing water mass tracer proxy to understand water mass reorganization during a threshold climatic event. The first step uses Pb isotopes from seawater and deep sea residues to demonstrate that reliable seawater Pb isotopes could be extracted from deep sea sediments. Bulk sediment f rom ten samples ranging in age from mid Eocene to early Miocene fro m ODP site 1090 on the Agulhus Ridge was subjected to a sequential leaching protocol that involved exposure to acetic acid, followed by a first and second HH extraction. Agreement of Pb isotopes between the three fractions indicates a common deep water source. During the course of this study, excellent inter and intralab reproducibility of Pb isotopes in the HH extracts were observed, which emphasize the reliability of the developed extraction protocol. Several lines of evidence provide additional support for the idea that measured Pb isotopes represent seawater These include an Al mass balance approach to evaluate the impact of detrital contaminants, comparison of Sr isotopes from the HH extraction and the seawater Sr isotope curve, and comparison of Nd isotopes from the HH extraction and coexisting fossil fish teeth.
96 FeMn oxide nodule standards NOD A1 and NOD P1 are recognized to record ambient bottom water isotopic signatures and close agreement between total digestion and HH extracted Pb isotopes from these nodules also indicates that the HH extracted P b isotopes represent bottom water values. Thus, all tests suggest that the measured Pb isotopes in HH extractions not only represent seawater, but presumably record bottom water values. During the course of this study, the fossil fish teeth were tested as another potential archive for bottom water Pb isotopic signals. Unconstrained combinations of initial Pb incorporated in the biophosphate, Pb incorporated during diagenietic enrichment, and radiogenic Pb produced from in situ decay of U and Th renders this archive unsuitable for Pb isotopic studies. Major element data from the leaching experiment indicates the presence of manganon carbonate, a refractory carbonate phase, in addition to FeMn oxides. Cross plots between HREE/LREE vs. MREE/MREE* indicate that such plots can be an effective tool to identify cross contamination during the leaching experiment. Once the Pb isotopes in FeMn oxides and similar authigenic phases extracted from bulk sediments were established as archives for bottom water Pb isotopes, this proxy was used to understand the weathering history of Antarctica during the EoceneOligocene (E/O) transition. The E/O transition marks the beginning of largescale glaciation on continental Antarctica. Distinct radiogenic changes in residue 206Pb /204Pb, 207Pb/204Pb and 208Pb/204Pb during the E/O boundary indicate weathering intensification of fresh, continental silicate rocks, while decoupling of seawater and residue 206Pb/204Pb indicate that one of the major contributors to Antarctic weathering f luxes during the EOT glaciation was old carbonate.
97 Seawater Nd isotopes preserved in fossil fish teeth are a well recognized water mass tracer that is frequently used in paleoceanography reconstruction of water mass sources and flowpaths. It is believed that the glacial Southern Ocean played a critical role in the storage of a high density, saline, CO2rich, 14C depleted water mass during t he last glacial period. Recent study of 14C from foraminifera revealed evidence of 14C depleted old water from intermediate ocean depths (~700 m) in the tropical North 14C from foraminifera provide the age of the water mass involved, but they provide no information about source. In this study I used Nd isotopes in fossil fish teeth/debris to demons trate that Antarctic Intermediate Water most likely carried this old carbonrich water from the deep Southern abyss during the last deglaciation.
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113 BIOGRAPHICAL SKETCH Chandranath Basak was born in Calcutta, India. He went to Presidency College, Calcutta, India for his bachelor s degree and became interested in geology during his first geochemistry course. Following graduation from college, he pursued a masters degree at Jadavpur University, Calcutta, India where he studied an inverted Barrovian sequence in the eastern Himalayas and was exposed to a range of problems related to metamorphic petrology. Following successful completion of the mast ers program, Chandranath joined Indiana State University (ISU), Terre Haute, Indiana to pursue a second Master of S cience degree. For this research he was involved in a modern proxy calibration study using stable isotopes in live benthic foraminifera from Alaska and the Australian margin under the guidance of Dr. Anthony E. Rathburn. Towards the end of his second masters he decided he wanted to continue in academics and pursue a PhD using the calibrated proxy to understand past climate. At that point he j oined the Ph.D. program at the University of Florida, Gainesville working on research in paleoclimatology under the guidance of Dr. Ellen E. Martin. Chandranths research at UF focused on establishing new paleoceanographic proxies as well as using those pr oxies to reconstruct paleoclimatic histories across dramatic climatic events on a range of time scales. After his Ph.D. he plans on continuing research in the field of paleoceanography/paleoclimatology and intends to remain dedicated to the world of academ ia.