UBC Theses and Dissertations

UBC Theses Logo

UBC Theses and Dissertations

Palaeoceanography of the Northeastern Pacific Ocean off Vancouver Island, Canada McKay, Jennifer Lynn 2003

You don't seem to have a PDF reader installed, try download the pdf

Item Metadata

Download

Media
ubc_2003-860140.pdf [ 11.73MB ]
Metadata
JSON: 1.0052403.json
JSON-LD: 1.0052403+ld.json
RDF/XML (Pretty): 1.0052403.xml
RDF/JSON: 1.0052403+rdf.json
Turtle: 1.0052403+rdf-turtle.txt
N-Triples: 1.0052403+rdf-ntriples.txt
Original Record: 1.0052403 +original-record.json
Full Text
1.0052403.txt
Citation
1.0052403.ris

Full Text

Palaeoceanography of the Northeastern Pacific Ocean off Vancouver Island, Canada by Jennifer Lynn McKay H.B.Sc. The University of Western Ontario, 1988 M.Sc., The University of Western Ontario, 1992 A thesis submitted in partial fulfillment of the requirements for the degree of Doctor of Philosophy in The Faculty of Graduate Studies (Department of Earth and Ocean Sciences) We accept this thesis^s conforming to the required standards THE UNIVERSITY OF BRITISH COLUMBIA September 2003 © Jennifer Lynn McKay, 2003 UBC Rare Books and Special Collections - Thesis Authorisation Form 5/29/03 12:25 PM In presenting this thesis in partial fulfilment of the requirements for an advanced degree at the University of British Columbia, I agree that the Library shall make it freely available for reference and study. I further agree that permission for extensive copying of this thesis for scholarly purposes may be granted by the head of my department or by his or her representatives. It is understood that copying or publication of this thesis for financial gain shall not be allowed without my written permission. The University of British Columbia Vancouver, Canada Date http://www.library.ubc.ca/spcoll/thesauth.html Page 1 of 1 Abstract Marine sediment cores from the continental margin off Western Canada (48° to 50°N, 125 to 128°W) yield evidence of dramatic changes in oceanographic conditions over the last 16 kyr. During the late Glacial the accumulation of marine organic matter was greatly reduced. This reflects low primary and export production because glacial-mode atmospheric circulation did not drive coastal upwelling. At the start of the Boiling warm period (-14.3 kyr B.P.), and coincident with the retreat of glaciers from the continental shelf, there was a substantial increase in the burial of organic matter. However, much of this material was "old" terrestrial organic detritus derived in part from the erosion of shelf sediments. It was not until the Allerod (-13.5 to 12.6 kyr B.P.) that the accumulation of marine organic matter increased substantially. Other paleoproductivity proxies (i.e., % biogenic Ba, % opal and alkenone abundances) also indicate very high marine productivity at this time. Since primary production in the region is controlled by the upwelling of nutrient-rich subsurface waters, this suggests upwelling was enhanced during the Allerod. A decrease in benthic-planktonic age differences (i.e., older planktonic foram ages) supports this interpretation. There was a brief return to glacial conditions (i.e., lower primary and export production) during the Younger Dryas, followed by a slight rise in organic carbon burial between -11 and 10 kyr B.P.. In general, the accumulation of organic carbon was low throughout the Holocene despite high primary productivity. Low sedimentation rates and the resulting long oxidant exposure times as well as extensive biological recycling appear to be the primary causes of low organic carbon burial, although lower productivity relative to the Allerod cannot be ruled out. Changes in the vertical settling flux of organic carbon to the sediment had a direct impact on sedimentary redox conditions at Site JT96-09 (920 m water depth). At present, near-surface sediments become suboxic within millimetres of the sediment-water interface and Re enrichment is observed below the depth of bioturbative mixing. However, anoxic conditions do not develop despite relatively low bottom water oxygen concentrations (0.3 ml/1) and relatively high organic carbon flux to the sea floor. The situation was much different during the deglacial. ii During the Allerod, when the organic flux to the sediment peaked, Mo enrichment is observed (>2 ug/g) implying that near-surface sediments were anoxic at this time. The benthic foraminifera assemblage (i.e., the dominance of Bolivina spp.) likewise suggests that sediments were more oxygen-depleted. These results imply that the bottom waters at Station JT96-09 were more oxygen-depleted during the Allerod and since JT96-09 is located within the oxygen minimum zone (OMZ), it follows that the OMZ was more intense. The increase in marine organic carbon accumulation and the lack of evidence of decreased ventilation suggests that OMZ intensification was a response to increased primary production and carbon export. Deglacial intensification of the OMZ is documented along the entire length of the California Current System, but it appears to have been delayed by ~1500 years off Vancouver Island. It is probable that the presence of the Laurentide and Cordilleran ice sheets continued to influence atmospheric circulation, and in turn oceanic circulation, in the region until at least the Allerod. It is possible to use redox-sensitive trace metals as palaeo-proxies in the deglacial sediments of Core JT96-09 because sedimentation rates were quite high (>100 crn/kyr). In contrast, in the slowly deposited Holocene sediments (5 cm/kyr) the trace metal record is corrupted due to the precipitation of trace metals decimetres below the sediment-water interface. This enrichment obscures the original palaeo-signal, which formed shortly after sediment deposition. Rapid fluctuations in the sedimentation rate can also complicate trace metal records. In the deglacial deposits of Core JT96-02 the formation of Mo and Re concentration spikes is the direct consequence of episodic deposition of turbidites that restricted oxygen influx into the sediment and thus allowed anoxic conditions to develop. Intriguing results were obtained from the measurement of sedimentary Ag concentration in near-surface sediments. Unlike the other redox-sensitive trace metals, for example Cd which has a similar geochemical behaviour, the primary control on Ag concentration is not sedimentary redox conditions. Rather, it seems that Ag accumulates in, and is transferred to the sediment by, settling organic particles in much the same way that Ba is. Hence, measuring the Ag concentration of sediments may provide another means of assessing palaeoproductivity. iii Table of Contents Abstract ii Table of Contents v List of Figures viList of Tables ix Acknowledgements xi 1. Palaeoceanography of the Northeastern Pacific Ocean off Vancouver Island Canada 1 1.1 Introduction1.2 The study area 3 1.3 Objectives 7 1.4 Presentation of results1.5 References 11 2. Organic Carbon Accumulation over the Last 16 kyr off Vancouver Island, Canada: Evidence for Increased Marine Productivity during the Deglacial 14 2.1 Introduction 12.2 Materials and methods 16 2.3 Results.. 21 2.3.1 Radiocarbon data for bulk organic carbon 22.3.2 Organic carbon content 22.3.3 Organic carbon/nitrogen ratios 3 2.3.4 Carbon- and nitrogen-isotope data 22.3.5 Biogenic barium and opal data 9 2.3.6 Biomarker data 22.4 Discussion 31 2.4.1 Organic carbon2.4.2 Biogenic barium and opal accumulation 44 2.4.3 Alkenones 45 iv 2.4.4 Nitrogen-isotope record 47 2.5 Summary 51 2.6 References 3 3. Intensification of the Oxygen Minimum Zone in the Northeast Pacific during the Last Deglaciation: Ventilation or Export Production? 65 3.1 Introduction 63.2 Materials and methods 71 3.3 Results 75 3.4 Discussion 83.4.1 Evidence of OMZ intensification 81 3.4.2 Ventilation changes? 86 3.4.3. Changes in productivity? 8 3.5 Summary 93 3.6 References 5 4. Accumulation of Redox-sensitive Trace Metals in Continental Margin Sediments off Western Canada 101 4.1 Introduction4.2 Methods 104 4.3 Results Ill 4.4 Discussion 119 4.4.1 C and N isotope data 114.4.2 Iodine and manganese 124.4.3 Rhenium 130 4.4.4 Uranium 2 4.4.5 Cadmium 3 4.4.6 Molybdenum 134 4.4.7 Barium 5 4.4.8 Silver 8 4.5 Summary 141 4.6 References 3 v 5. Geochemical Response to Pulsed Sedimentation on the Western Canadian Continental Margin: Implications for the Use of Mo as a Palaeo-oxygenation Proxy 150 5.1 Introduction 155.2 Geochemical variations 153 5.3 Origin of the sulphide layers 4 5.4 Trace metal geochemistry 161 5.5 Summary 165 5.6 References 7 6. Summary 170 6.1 Sedimentation of the Vancouver Island Margin 170 6.2 The sea-surface temperature record 176.3 The marine productivity record 1 6.4 Redox-sensitive trace metals and their palaeo-applications 173 6.5 Fluctuations in OMZ intensity 174 6.6 Biogenic Ag? 175 6.7 Future work6.8 References 8 Appendices 180 vi List of Figures Fig. 1.1. The study area off the west coast of Vancouver Island, British Columbia, Canada (Inset) showing the locations where cores were collected 4 Fig. 2.1. The study area is located off the west coast of Vancouver Island, British Columbia, Canada (Inset). Sediment cores were collected from Station JT96-09 that is located on the continental slope at a water depth of 920 m 17 Fig. 2.2. Geochemical data for Core JT96-09 24 Fig. 2.3. Downcore profiles of the organic carbon/total nitrogen ratio (OC/TN) and C29 n-alkane/organic carbon ratio (C29/OC) 28 Fig. 2.4. N-isotopic composition of the bulk sediment (515N) versus the C-isotopic composition of organic matter (813Corg) 30 Fig. 2.5. C29 n-alkane normalized to organic carbon (C29/OC) versus a) the C-isotopic composition of the organic matter (513Corg), and b) the N-isotopic composition (515N) of the bulk sediment 32 Fig. 2.6. The percentage of terrigenous organic carbon (i.e., the terrigenous fraction) in Core JT96-09 estimated using both the 813Corg and 515N data 37 Fig. 2.7. a) Terrestrial organic matter concentration (open circles) and mass accumulation rate (MAR, solid circles) over the past 16 kyr B.P. b) Marine organic matter concentration (open squares) and mass accumulation rate (MAR, solid squares) over the past 16 kyr B.P 40 Fig. 2.8. The measured concentration of alkenones, a marine biomarker, versus the estimated marine organic carbon concentration, b) The measured concentration of C2Q n-alkane, a terrestrial biomarker, versus the estimated terrestrial organic carbon concentration 41 Fig. 2.9. The downcore Zr/Al profile for core JT96-09 3 Fig. 3.1. The study area is located off the west coast of Vancouver Island, British Columbia, Canada (Inset). Sediment cores were collected from Station JT96-09 that is located on the continental slope at a water depth of 920 m 70 Fig. 3.2. Plot of 14C ages of planktonic and benthic foraminifera versus sample depth in Core JT96-09 76 Fig. 3.3. Benthic-planktonic age differences are plotted against calendar age 79 Fig. 3.4. Redox-sensitive trace metal concentrations (symbols) and metal/Al ratios (thick lines) in Core JT96-09 80 Fig. 3.5. Various palaeo records for Core JT96-09 (8 to 16 kyr B.P.) 82 vii Fig. 3.6. Marine organic carbon concentration (open squares) and mass accumulation rate (MAR, solid squares) at Station JT96-09 from 8 tol6 kyr B.P 90 Fig. 4.1. The study area off the west coast of Vancouver Island, British Columbia, Canada (Inset) showing the locations where multicores (mc) and box cores (be) were collected. Exact water depths and core descriptions are provided in Table 4.1 105 210 Fig. 4.2. Concentrations of "excess" Pb in near-surface sediment cores Olmc, 02bc and 09mc 114 Fig. 4.3. Concentrations of a) organic carbon, b) carbonate, and c) opal in near-surface sediment cores 5 13 Fig. 4.4. a) The C-isotopic composition of organic matter (8 Corg) and the N-isotopic composition of bulk sediment (815N) in near-surface sediment cores 117 Fig. 4.5. a) I/Corg, b) Mn/Al, and c) Ba/Al ratios in near-surface sediment cores 118 Fig. 4.6. a) Re and b) U concentrations in the near-surface sediment cores 120 Fig. 4.7. The concentrations of a) Cd, b) Mo, and c) Ag in the near-surface sediment cores 121 Fig. 4.8. Concentration of Zr versus the concentration of total Ba in near-surface sediments from the Vancouver Island Margin 137 Fig. 4.9. Plots of Ag concentrations versus a) % opal, b) % organic carbon, c) Zr/Al ratios, and d) biogenic Ba content, in near-surface sediment cores 140 Fig. 5.1. The study area is located off the west coast of Vancouver Island, British Columbia, Canada (Inset). Piston core JT96-02 was collected from the continental slope at a water depth of 1340 m 152 Fig. 5.2. Downcore profiles of a) total sulphur (Stot) and organic carbon (Corg) contents, b) carbonate content, and c) 813Corg and 815N values 155 Fig. 5.3. Downcore profiles of the a) Si/Al ratio and b) Zr/Al and Ti/Al ratios 156 Fig. 5.4. a to e) Downcore profiles of trace metal concentrations (symbols) and metal/Al ratios (thick black lines). Shaded areas indicate the locations of sulphide layers. Background (i.e., lithogenic) concentrations of the various metals are marked by the dashed lines (see Table 5.1 for references), f) Downcore profile of the sedimentary Re/Mo ratio. The dashed line represents both the Re/Mo ratio of seawater and the estimated lithogenic Re/Mo ratio 157 Fig. 5.5. Plot of Mo concentration versus total sulphur content in the sulphide layers (solid circles) and the intervening sediments (open circles) 159 Fig. 6.1. A comparison of Re and Ag concentrations over the last 16 kyr in Core JT96-09 176 viii List of Tables Table 1.1. List of the cores studied, their geographic locations and water depths, as well as various types of data provided in the appendices 5 Table 2.1. Radiocarbon data and estimates of the input of organic matter of infinite radiocarbon age for Composite Core JT96-09 22 Table 2.2. Geochemical data for Composite Core JT96-09 26 Table 3.1. Radiocarbon ages of plantonic and benthic foraminifera in Core JT96-09 77 Table 3.2. Oxygen exposure times for sediments in Core JT96-09 92 Table 4.1. General data for sampling locations and core descriptions 106 Table 4.2. Results of biogenic barium dissolution tests 109 Table 4.3. Radiocarbon data and estimated sedimentation rates 112 Table 4.4. General geochemical data for Vancouver Island Margin multicores (mc) and box cores (be) 12Table 4.5. Major, minor and trace element data for Vancouver Island Margin multicores (mc) and box cores (be) 125 Table 5.1. Seawater and lithogenic concentrations of the redox-sensitive trace metals 164 Table Al. List of the cores studied, their geographic locations and water depths, as well as various types of data provided in the appendices 185 Table A2. Description of Sediment Cores 186 Table A3. CTD data for Stations JT96-01, 2, 4, 6 and 9 187 Table A4. Dissolved oxygen data for Stations JT96-01, 2,4, 5, 6 and 9 192 Table A5. Magnetic susceptibility data for sediment cores from Stations JT96-02, 5, 6 and 9 193 Table A6. Radiocarbon data for planktonic and benthic foraminifera, and bulk organic carbon 195 Table A7. 210-Pb data for multicores JT96-01, 02 and 09 196 Table A8. Stable isotope data for foraminifera from Cores JT96-02, 06pc, 09mc and 09pc, and Tul96-05pc, 03tc and 03pc 197 Table A9. Geochemical data for Multicore JT96-01 201 Table A10. Geochemical data for Multicore and Piston Core JT96-02 203 Table Al 1. Geochemical data for Multicore JT96-04 210 ix Table A12. Geochemical data for Box Core and Trigger Core JT96-05 212 Table A13. Geochemical data for Multicore, Box Core and Piston Core JT96-06 216 Table A14. Geochemical data for Multicore and Piston Core JT96-09 220 Table A15. Geochemical data for Triger Core and Piston Core Tul96-03 228 Table A16. Geochemical data for Trigger Core and Piston Core Tul96-05 230 x Acknowledgements Yes IT is finally done, but it would not have been possible without the help of many people. First and foremost to my advisor Dr. Tom Pedersen. Thank you for taking the risk of accepting me as a PhD student. I'm sure you had your doubts about the sanity of that decision but I hope in the end that you think it proved sagacious. Thank you for your guidance, patience and understanding and your research money. Thank you for helping me to discover the world of palaeoceanography and the world in general. To Dr. Steve Calvert. I'm at a loss for words. You are an incredible scientist in every sense. Your love for, and dedication to, science is truly inspiring. It has been an honour to get to know you and to learn from you. To my thesis committee (Steve Calvert, Kristen Orians and Kurt Grimm). I am indebted to you for keeping me on track and for reviewing my thesis. I hope it was not too trying. To Bente, Bert, Kathy and Maureen. Your help in the lab was tremendous and your friendship equally as important. Thank you for everything. To the members of the geochemistry group. Thank you for the stimulating discussions and feedback on my work. A special thank to Stephanie Kienast without whom there would be no SST record on which so much of my work relies. Stephanie, it has been a pleasure to work with you and if I'm very fortunate I will have the opportunity to do so again in the future. xi Endless thanks to my office mates in the NCE building for answering all of my biology, chemistry and physical oceanography questions, and for putting up with my hyper and stressed-out self. A special thanks to the Aussies. Your unique sense of humor never failed to cheer me up when things were going to heck in the lab. As a special note to Andrew, people watching is just not the same without you. To the office staff of Earth and Ocean Sciences. I greatly appreciate all of your help throughout my stay at UBC. Special thanks to Alex and Carol for keeping me in touch when I ran off to Montreal. Lastly to my friends and family. Thank you for your love and support over the seemingly endless years of schooling. I may not have said it often enough but I love you all very much. Now just a warning. The thesis may be done but the science continues. If you haven't guessed by now I actually like the endless hours in the lab, the stress when things don't work and the thrill when they do, the field work and of course the conferences (Hawaii...need I say more). For better or for worse, I'm in this for life. Just one last thing. It is now OK to use the "T" word. xii 1. Palaeoceanography of the Northeastern Pacific Ocean off Vancouver Island, Canada 1.1 Introduction With the recognition that the activities of mankind are impacting Earths climate (IPCC, 2001) there has been renewed interest in the study of natural climate change, its magnitude, the speed at which it occurs and the response of various earth systems, for example the biosphere and cryosphere, to such change. Some of the best records of climate variability over the past 500,000 years have been obtained from the Greenland and Antarctica ice sheets (e.g., Barnola et al., 1987; Dansgaard et al., 1993; Petit et al., 1999). Marine sediment cores are another source of information and their analysis provides a better global coverage, although obtaining high resolution data equivalent in quality to information obtained from ice cores is often hindered by bioturbation coupled with slow sedimentation that characterize the majority of sedimentary deposits on the sea floor. Many initial palaeoceanographic studies focused on the Atlantic Ocean and in particular the North Atlantic, in part because this is a principal location of deepwater formation at present. In rather stark contrast, less work has been conducted in the Pacific Ocean and very little in the Subarctic Northeast Pacific. Only four palaeoceanographic studies exist for the Gulf of Alaska (Zahn et al., 1991; Sabin and Pisias, 1996; de Vernal and Pedersen, 1997; McDonald et al., 1999), and even fewer exist for the region off western Canada. Only two cores have been studied in any detail (TT39-PC17 and TT39-PC12; Sabin and Pisias, 1996) and these were collected from sites located hundreds of kilometres offshore. There have been no palaeoceanographic studies of the upwelling region off Vancouver Island, that is until very recently (Kienast and McKay, 2002). The northeast Pacific Ocean is a region of primary interest for a number of reasons: 1 i) The North Pacific is the terminus of deep water circulation. As such, the deep waters in the region are a reservoir of substantial amounts of C02. Depending on various factors (e.g., primary productivity, upwelling, sea-surface temperature, wind mixing) the North Pacific could be a source or sink for atmospheric CO2. Recent work has suggested that the North Pacific was a significant C02 sink from 1979 to 1992 (Wong et al., in review), although other studies suggest the exact opposite (e.g., Stephens et al., 1995). In the Pleistocene, periods of high diatom productivity caused reduced [C02](aq) in mixed-layer (McDonald et al., 1999) and could potentially have driven the uptake of atmospheric C02. ii) The North Pacific is a region of intermediate water formation and ventilation. Changes in the ventilation rate of intermediate waters could have influenced the intensity (i.e., degree of oxygen depletion) of the oxygen minimum zone (OMZ) throughout the North Pacific. This has important implications for the oceanic nitrogen cycle (i.e., nitrate availability) and in turn ocean productivity (Ganeshram et al., 1995). iii) The California Current System (CCS), which extends from Vancouver Island, Canada to Baja, California, is one of four major eastern boundary current systems. Upwelling of nutrient-rich waters causes these regions to be important locations of organic matter production, export and burial. However, because upwelling is primarily wind-driven, it is highly susceptible to even minor changes in wind strength and atmospheric circulation patterns. It is thus reasonable to assume that the large scale changes in atmospheric circulation patterns that accompanied glaciations (COHMAP, 1988) would have had a significant impact on upwelling and thus ocean productivity. Such was the case off Oregon and northern California during the last glacial (Lyle et al., 1992; Dean et al., 1997; Ortiz et al., 1997; Mix et al., 1999). In turn, there is the potential that atmospheric C02 concentrations could have been affected. 2 iv) Changes in sea surface temperatures in the north Pacific, for example the 6°C increase observed off Vancouver Island during the last deglacial (Kienast and McKay, 2002), could impact the hydrological budget of North America and more generally that of the northern hemisphere (Peteet et al., 1997). This thesis presents the first comprehensive palaeogeochemical investigation of the late Pleistocene to Holocene evolution of the northeast Pacific Ocean in the area west of Vancouver Island, Canada. As such, it addresses directly the existing dearth of information on the palaeoceanography of this important but understudied region. 1.2 The study area Sediment cores were collected during a 1996 cruise of the Canadian Coast Guard research vessel the John Tully. Box- and multicores were collected at six stations (01, 02, 04, 05, 06 and 09; Fig. 1.1). Piston cores were also collected at stations 02, 05, 06 and 09 (Fig. 1.1). A second suite of piston cores from the same area were made available by Dr. George Spence of the University of Victoria. Two of these cores (Tul96-03 and Tul96-05; Fig. 1.1) were sampled for palaeoceanographic analysis. Core locations and a summary of the data available for each core are provided in Table 1.1. Preliminary analyses revealed that only piston cores JT96-09pc and JT96-02pc were suitable for in depth palaeoceanographic analysis of the last deglaciation and Holocene. Piston Core JT96-05pc contained too many sandy turbidites while Piston Core JT96-06pc did not yield a recognizable glacial-interglacial oxygen-isotope stratigraphy that is critical for palaeoceanographic studies. Analysis of Piston Core Tul96-03 ceased when it became apparent from the 8180 measurements made on benthic foraminifera that a significant portion of the core top, including all of the Holocene, had been lost during collection of the core. Piston Core Tul96-05 also contained a large time gap from -10.9 to 27.9 calendar kyr B.P., and this limited its usefulness. 3 127°W 126°W 125°W Fig. 1.1. The study area off the west coast of Vancouver Island, British Columbia, Canada (Inset) showing the locations where cores were collected. Further details can be found in Table 1.1. 4 J3 00 (N •S E Os H S o vo H D. ON O o\ vo *o ON Ov H H 5 The study area is located at the northern end of the California Current System (CCS) where the eastward flowing Subarctic and North Pacific currents split into the northward flowing Alaska Current and southward flowing California Current (Fig. 1.1). It is situated in a transitional zone where even minor changes in atmospheric and oceanic circulation might influence sea surface temperature (SST), salinity, water column stratification, as well as primary productivity via changes in nutrient supply. Given the appropriate conditions, these changes may be recorded in the underlying sediments. However, palaeoceanographic studies in the region have been hindered by the occurrence of turbidites and the scarcity of foraminifera which makes the development of a robust chronostratigraphy difficult (Zahn et al., 1991). Furthermore, the character of sedimentation on the Vancouver Island Continental Margin has changed dramatically over time. During the last glacial piedmont-type glaciers extended on to the shelf as far as the shelf edge west of the Strait of Juan de Fuca and almost to the shelf edge west of Barkley Sound (Herzer and Bornhold, 1982). These glaciers reached their maximum extent between -15 and 14 14C kyr and then began to rapidly retreat (Clague and James, 2002). Coincident with deglaciation, sedimentation rates on the slope (Station JT96-09) increased substantially (i.e., from -47 cm/kyr up to 169 cm/kyr). Ice had retreated from the shelf by -13 14C kyr (Blaise et al., 1990; Josenhans et al., 1995). This was a period of global sea level rise; however, as a result of isostatic rebound the land rose and relative sea level was -150 m lower than at present between at -10 14C kyr B.P. (Josenhans et al., 1995; Barrie and Conway, 1999). Modern sea level was not established until -9.1 14C kyr (Josenhans et al., 1995). There was a substantial decrease in the sedimentation rate on the shelf as sea level rose due to trapping of sediments within the fjords of Vancouver Island and winnowing by waves and currents (Bornhold and Yorath, 1984). As a result, Holocene sediments on the inner and mid-shelf typically comprise sand and gravel lag deposits, while on the outer shelf Pleistocene clays are mantled by thin layers (< 30 cm) of sandy sediments (Bornhold and Barrie, 1991; Bornhold and Yorath, 1984). Modern sedimentation rates on the upper slope are also low (<2 cm/kyr; Stations JT96-04 and JT96-06). 6 1.3 Objectives A number of specific question are addressed in this thesis: i) How has primary productivity, export productivity, and carbon burial on the Vancouver Island Continental Margin varied on glacial-interglacial and shorter timescales? ii) Has nutrient availability and utilization varied over the last 16 kyr? If so, can these variations be related to changes in upwelling and/or climatic conditions? iii) At present, the region is characterized by a moderately intense oxygen minimum zone (OMZ) (~ 0.3 ml/1 02 at 920 m water depth). Has the intensity (i.e., degree of oxygen-depletion), thickness and/or depth of the OMZ varied over time. If so, have denitrification within the water column and organic carbon burial been affected? Are variations in OMZ intensity related to changes in intermediate water mass hydrography (i.e., ventilation) and/or changes in the settling flux of organic matter (i.e., primary productivity)? iv) How does the palaeoceanography of the study area compare with other regions within the California Current System? What do the similarities and differences tell us about the response of the northeast Pacific to climate change? 1.4 Presentation of results This thesis has been written in paper format. Each chapter, excluding the introduction and summary (i.e., Chapters 1 and 6, respectively), represents one paper that deals with specific palaeoceanographic questions. Chapter 2 (McKay et al., in press) presents an investigation of the nature of organic matter accumulation on the continental slope (Piston Core JT96-09, 920 m water depth) off 7 Vancouver Island. Both concentration and mass accumulation rates of organic carbon in marine sediments are commonly used as palaeoproductivity proxies (e.g., Pedersen, 1983; Sarnthein et al., 1988). However, marine deposits often contain substantial quantities of terrestrial organic matter, particularly along the continental margin (e.g., Prahl et al., 1994). It is therefore necessary to quantify the fraction of marine organic matter if variations in primary and export production are to be inferred. Furthermore, the relative abundances of 13 marine and terrestrial organic matter must be known if other data, for example the 8 C of 13 15 organic carbon (8 Corg) and 8 N of bulk sediments, are to be correctly interpreted. The identification and quantification of terrestrial organic matter are made in Chapter 2 using a variety of data (i.e., Corg/N ratios, C and N isotope data, n-alkane concentrations and radiocarbon ages of bulk organic matter). The total organic carbon data are then corrected for the terrigenous input to yield the accumulation rate of marine organic carbon for the past 16 kyr B.P.. These results are compared with other palaeoproductivity records based on the mass accumulation rates of biogenic barium and opal, as well as the abundance of C37 alkenones. It should be noted that the analysis of biomarkers (i.e., alkenones and n-alkanes) was conducted in co-operation with S. Kienast at the University of British Columbia. Chapter 3 investigates whether the intensity (i.e., degree of oxygen-depletion) of the oxygen minimum zone (OMZ) off Vancouver Island has changed in the past 16 kyr B.P. and if so, why. Piston Core JT96-09, that comes from the most intense portion of the modern OMZ (920 m water depth), is studied. Along the southern portion of the California Current System the OMZ has fluctuated on glacial-interglacial and shorter timescales (Keigwin and Jones, 1990; Behl and Kennett, 1996; Dean et al., 1997; Ganeshram et al., 1995; Cannariato and Kennett, 1999; Zheng et al., 2000). Redox-sensitive trace metal and benthic foraminifera species data presented in this chapter suggest that the OMZ off Vancouver Island has also fluctuated. Two possible causes are discussed: i) a change in ventilation of intermediate waters, and/or ii) a change in primary productivity and the subsequent export and degradation of the organic matter. Radiocarbon dating of coeval benthic and planktonic foraminifera is 8 employed to assess whether fluctuations in OMZ intensity are the result of ventilation changes. This method assumes that the apparent age of surface waters, recorded by planktonic foraminifera, have not changed while the age of intermediate waters, which is recorded by benthic foraminifera, might have varied. The accumulation rate of marine organic carbon, determined in Chapter 2, is employed as a proxy for palaeoproductivity. The possibility that organic carbon accumulation is influenced by OMZ intensity is also discussed. In Chapter 4 the accumulation of redox-sensitive trace metals (Re, U, Cd, Mo and Ag) in near-surface sediments deposited on the continental margin off Vancouver Island is investigated. Data for multicores and boxcores from six stations (01, 02, 04, 06, 05 and 09; Fig. 1.1) are presented. The sedimentary concentrations of such trace metals can be influenced by redox conditions within the sediment. Sedimentary redox conditions are in turn influenced by organic carbon flux to the sediment and oxygen concentrations in the overlying bottom water. Changes in the accumulation of redox-sensitive elements are thus commonly used to identify changes in palaeoproductivity and/or bottom water oxygenation over time (e.g., Crusius et al., 1999; Yarincik et al., 2000; Adelson et al., 2001; Pailler et al., 2002). However, a variety of other factors (e.g., sedimentation rate, metal source and post-depositional migration) can also influence metal concentrations. An understanding of the key processes that govern such accumulation is required if trace metal palaeo-records are to be correctly interpreted. In Chapter 5 data for Piston Core JT96-02 (1340 m water depth) are presented. The deglacial clay in this core contains numerous centimetre-scale, organic- and sulphide-rich layers. A 16 cm interval containing three such layers was sampled at very high resolution (3 mm spacing) and a variety of geochemical measurements were made. Three issues are addressed: i) what was the nature of deglacial sedimentation on the Vancouver Island Continental Margin, ii) why did the sulphide-rich layers form, and iii) what factors controlled the accumulation of redox-sensitive trace metals in these layers. The final conclusion is that 9 an abrupt increase in sedimentation rate due to the episodic deposition of thin turbidites was ultimately responsible for the development of anoxic conditions and formation of the sulphide-layers. There was no apparent decrease in bottom water oxygen concentration or increase in organic carbon accumulation. Thus, this chapter highlights the dangers of using trace metals as palaeo-proxies for bottom water ventilation and organic carbon flux. Each of chapters 2 through 5 has been submitted as a stand-alone paper, and four different journals have been targeted as venues for publication. This approach unavoidably leads to some overlap in the chapters, which is primarily limited to description of the study area and the hydrography of the region. 10 1.5 References Adelson, J.M. Helz, G.R., Miller, C.V., 2001. Reconstructing the rise of recent coastal anoxia; molybdenum in Chesapeake Bay sediments. Geochimica et Cosmochimica Acta 65, 237-252. Barnola, J.M., Raynaud, D., Korotkevich, Y.S., Lorius, C, 1987. Vostok ice core provides 160,000-year record of atmospheric C02. Nature 329, 408-414. Barrie, J.V., Conway, K.W., 1999. Late Quaternary glaciation and postglacial stratigraphy of the Northern Pacific margin of Canada. Quaternary Research 51, 113-123. Behl, R.J., Kennett, J.P., 1996. Brief interstadial events in the Santa Barbara Basin, NE Pacific, during the past 60 kyr. Nature 379, 243-246. Blaise, B., Clague, J.J., Mathewes, R.W., 1990. Time of maximum Late Wisconsin glaciation, west coast of Canada. Quaternary Research 34, 282-295. Bornhold, B.D., Barrie, J.V., 1991. Surficial sediments on the western Canadian continental shelf. Continental Shelf Research 11,685-699. Bornhold, B.D., Yorath, C.J., 1984. Surficial geology of the continental shelf, northwestern Vancouver Island. Marine Geology 57, 89-112. Cannariato, K.G., Kennett, J.P., 1999. Climatically related millenial-scale fluctuations in strength of California margin oxygen-minimum zone during the past 60 k.y.. Geology 27, 975-978. Clague, J.J., James, T.S., 2002. History and isostatic effects of the last ice sheet in southern British Columbia. Quaternary Science Reviews 21, 71-87. COHMAP Members, 1988. Climatic changes of the last 18,000 years: Observations and model simulations. Science 241, 1043-1052. Crusius, J., Pedersen, T.F., Calvert, S.E., Cowie, G.L., Oba, T., 1999. A 36 kyr geochemical record from the Sea of Japan of organic matter flux variations and changes in intermediate water oyxgen concentrations. Paleoceaography 14, 248-259. Dansgaard, W., Johnsen, S.J., Clausen, H.B., Dahl-Jensen, D., Gundestrupt, N.S., Hammer, C.U., Hvidberg, C.S., Steffensen, J.P., Sveinbjornsdottier, A.E., Jouzel, J., Bond, G., 1993. Evidence for general instability of past climate from a 250-kyr ice-core record. Nature 364, 218-220. Dean, W.E., Gardner, J.V., Piper, D.Z., 1997. Inorganic geochemical indicators of glacial interglacial changes in productivity and anoxia on the California continental margin. Geochemica et Cosmochimica Acta 61, 4507-4518. de Vernal, A., Pedersen, T.F., 1997. Micropaleontology and palynology of core PAR87A-10: A 23,000 year record of paleoenvironmental changes in the Gulf of Alaska, northeast North Pacific. Paleoceanography 12, 821-829. 11 Ganeshram, R.S., Pedersen, T.F., Calvert, S.E., Murray, J.W., 1995. Large changes in oceanic nutrient inventories from glacial to interglacial periods. Nature 376, 755-758. Herzer, R.H., Bornhold, B.D., 1982. Glaciation and post-glacial history of the continental shelf off southwestern Vancouver Island, British Columbia. Marine Geology 48, 285-319. Josenhans, H.W., Fedje, D.W., Conway, K.W., Barrie, J.V., 1995. Post glacial sea levels on the Western Canadian continental shelf: evidence for rapid change, extensive subaerial exposure and early human habitation. Marine Geology 125, 73-94. Keigwin, L.D., Jones, G.A. 1990. Deglacial climatic oscillations in the Gulf of California. Paleoceanography 5, 1009-1023. Kienast, S.S., McKay, J.L., 2001. Sea surface temperature in the subarctic northeast Pacific reflect millennial-scale climate oscillations during the last 16 kyrs. Geophysical Research Letters 28, 1563-1566. Lyle, M., Zahn, R., Prahl, F. Dymond, J., Colier, R. Pisias, N., Suess, E., 1992. Paleoproductivity and carbon burial across the California Current: The mulitracers transect, 42°N. Paleoceanography 7, 251-272. McDonald, D., Pedersen, T.F, Crusius, J., 1999. Multiple late Quaternary episodes of exceptional diatom production in the Gulf of Alaska. Deep Sea Research II 46, 2993-3017. McKay, J.L., Pedersen, T.F., Kienast, S.S., in press. Organic carbon accumulation over the last 16 kyr off Vancouver Island, Canada: Evidence for increased marine productivity during the deglacial. Quaternary Science Reviews. Mix, A.C., Lund, D.C., Pisias, N.G., Boden, P., Bornmalm, L., Lyle, M., Pike, J., 1999. Rapid climate oscillation in the northeast Pacific during the last deglaciation reflect northern and southern hemisphere sources. In: Mechanisms of Global Climate Change at Millennial Time Scales, Geophysical Monograph 112,127-148. , Ortiz, J., Mix, A., Hostetler, S., Kashgarian, M., 1997. The California Current of the last glacial maximum: Reconstruction at 42°N based on multiple proxies. Paleoceanography, v. 12, p. 191-205. Pailler, D., Bard, E., Rostek, F., Zheng, Y., Mortlock, R., van Geen, A., 2002. Burial of redox sensitive metals and organic matter in the equatorial Indian Ocean linked to precession. Geochimica et Cosmochimica Acta 66, 849-865. Pedersen, T.F., 1983. Increased productivity in the eastern equatorial Pacific during the last glacial maximum (19,000 to 14,000 yr B.P.). Geology 11, 16-19. Peteet, D., Genio, A.D., Lo, K.K.W., 1997. Sensitivity of the northern hemisphere air temperatures and snow expansion to North Pacific sea surface temperature in the Goddard Institute for space studies general circulation model. Journal of Geophysical Research 102, 23781-23791. Petit, J.R., Jouzel, J., Raynaud, D., Barkov, N.I., Barnola, J.-M., Basile, I., Benders, M., Chappellaz, J., Davis, M., Delaygue, G., Delmotte, M., Kotlyakov, V.M., Legrand, M., Lipenkov, V.Y., Lorius, C, Pepin, L., Ritz, C, Saltzman, E., Stievenard, M., 1999. Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica. Nature 399, 429-413. 12 Prahl, F.G., Ertel, J.R., Goni, M.A., Sparrow, M.A., Eversmeyer, B., 1994. Terrestrial organic carbon contributions to sediments on the Washington margin. Geochimica et Cosmochimica Acta 58, 3035-3048. Sabin, A.L., Pisias, N.G., 1996. Sea surface temperature changes in the northeastern Pacific Ocean during the past 20,000 years and their relationship to climate change in northwestern North America. Quaternary Research 46,48-61. Sarnthein, M., Winn, K., 1990. Reconstruction of low and mid latitude export productivity, 30,000 years B.P. to present: Implications for control of global carbon reservoirs. In: Climate-Ocean Interaction (ed. M.E. Schlesinger), Klewer Academic Publishers, p. 319-342. Stephens, M.P., Samuels, G., Olson, D.B., Fine, R.A., Takahashi, T., 1995. Sea-air flux of C02 in the North Pacific using shipboard and satellite data. Journal of Geophysical Research 100,13571-13583. Wong, C.S., Chan, Y.-H., Feely, R.A., Goyet, C, Inoue, H., in review. North Pacific Ocean as a significant C02 sink during 1979-1992. Science. Yarincik, K.M., Murray, R.W., Lyons, T.W., Petersen, L.C., Haug, G.H., 2000. Oxygenation history of bottom waters in the Cariaco Basin, Venezuela, over the past 578,000 years: Results from redox-sensitive metals (Mo, V, Mn, and Fe). Paleoceanography 15, 593-604. Zahn, R., Pedersen, T.F., Bornhold, B.D., Mix, A.C., 1991. Water mass conversion in the glacial subarctic Pacific (54°N, 148°W): Physical constraints and the benthic-planktonic stable isotope record. Paleoceanography 6, 543-560. Zheng, Y., van Geen, A., Anderson, R.F., Gardner, J.V., Dean, W.E., 2000. Intensification of the northeast Pacific oxygen minimum zone during the Bolling-Allerod warm period. Paleoceanography 15, 528-536. 13 2. Organic Carbon Accumulation over the Last 16 kyr off Vancouver Island, Canada: Evidence for Increased Marine Productivity during the Deglacial 2.1 Introduction The history of palaeoproductivity, or more correctly export production, can be inferred using a variety of proxies such as the abundance and/or mass accumulation rate of organic carbon (e.g., Pedersen, 1983; Sarnthein et al., 1988), biogenic barite (e.g., Schmitz, 1987; Dymond et al., 1992) and opal (e.g., Lyle et al., 1992; Gardner et al., 1997) in marine sediments, as well as the concentrations of various marine biomarkers (e.g., Prahl et al., 1989; Jasper and Gagosian, 1993; Ohkouchi et al., 1997; Schubert et al., 1998; Werne et al., 2000). However, each proxy has its own particular set of problems and limitations. For example, proxies may be preferentially preserved when sedimentation rates are high. Palaeoproductivity estimates based on the amount and/or accumulation rate of organic carbon are further complicated by the presence of terrestrial organic matter which can constitute a large fraction of nearshore, non-deltaic marine sediments (> 50%; Prahl et al., 1994). Even deep ocean sediments in some areas can have a significant terrestrial component (e.g., 20%) due to aeolian transport (Prahl et al., 1989; Ohkouchi et al., 1997) and ice-rafting (Villanueva et al., 1997a). It is therefore necessary to estimate the relative abundance of marine and terrestrial organic matter when using organic carbon as a palaeoproductivity proxy. This information is also required if bulk, rather than compound specific, 8 C values of organic matter are to be used in the estimation of dissolved C02 concentrations (e.g., Rau et al., 1991; Rau, 1994), and if sedimentary 815N data are to be correctly interpreted. Little is known about palaeoproductivity and marine organic matter accumulation on the Western Canadian Continental Margin, in part because it has been extremely difficult to collect cores that preserve a continuous record of hemipelagic sedimentation. In this paper, 14 we present the first estimates of marine and terrestrial organic matter accumulation on the continental slope west of Vancouver Island for the last 16 kyr. These results are part of a larger IMAGES-program study of productivity history for the Late Quaternary along the western margin of the Americas. The study area is located at the northern end of the California Current System (CSS) off Vancouver Island, British Columbia, Canada (48° 54' N, 126° 53' W; Fig. 2.1). This region sits within a transition zone where the eastward-flowing Subarctic and North Pacific currents split into the northward flowing Alaska Current and southward flowing California Current (Fig. 2.1; Thomson, 1981). At present, the area immediately west of Vancouver Island is characterized by high primary productivity driven by seasonal (spring - summer) wind-induced upwelling which is related to the northerly position (38°N) of the North Pacific High (Thomas et al., 1994; Hickey, 1998). The strength of this upwelling is determined by the strength of the northerly, alongshore wind that is in turn controlled by the atmospheric pressure gradient between the land and ocean (Bakun, 1990). In winter, when the North Pacific High shifts southward to ~28°N and the Aleutian Low influences the region there is a reversal of the wind direction and cessation of upwelling. Palaeo-evidence suggests that during the last glacial maximum (LGM) the North Pacific High was weaker and displaced south of its present position (COHMAP, 1988; Thunell and Mortyn, 1995; Mortyn et al., 1996; Sabin and Pisias, 1996; Doose et al, 1997). As a result, upwelling was substantially reduced or non-existent along the central and northern portions of the CCS and in turn export productivity was lower (Dymond et al., 1992; Lyle et al., 1992; Sancetta et al., 1992; Ortiz et al., 1997; Dean and Gardner, 1998; Mix et al., 1999). However, in this paper we show that off Vancouver Island the mass accumulation rate of total organic carbon was higher during the late glacial in comparison to present. To explain the contrast with regions directly to the south a variety of geochemical data are used to characterize the type of organic matter and then the terrigenous input is then quantified using C- and N-isotopic data. The concentration and mass accumulation rate of 15 total organic carbon are corrected for the input of terrestrial organic matter to yield a palaeo-record of marine organic matter accumulation over the last 16 kyr. Finally, to circumvent concerns that variable preservation may have influenced the organic carbon record a variety of other palaeoproductivity proxies (i.e., biogenic barium, opal and C37 alkenone contents) are presented for comparison. 2.2 Materials and methods Sediment cores were collected from the continental slope off the west coast of Vancouver Island, British Columbia, Canada during a 1996 Canadian Joint Global Ocean Flux Study (CJGOFS) cruise. Data are presented here for a 374 cm long piston core (JT96-09pc) and a corresponding 40 cm long multicore (JT96-09mc), both collected at the same site (48° 54' N, 126° 53' W; Fig. 2.1) from a water depth of 920 m. The multicore and upper 51 cm of the piston core are composed of a homogeneous, carbonate-poor, olive-green mud. In the piston core this mud is underlain by 85 cm of gray-green clay at the base of which is a 16 cm thick sandy turbidite. The remainder of the core (222 cm) is composed of a dense, gray clay. The sediment-water interface was intact in the multicore and a comparison of geochemical data from the multicore and piston core suggests that ~12 cm, not 20 cm as stated in Kienast and McKay (2001), were lost off the top of the latter during its collection. As a result, piston core depths have been corrected for this loss (+12 cm) and for the presence of the turbidite (-16 cm). Records from the two cores were then merged to yield a composite record for site JT96-09. The age model for JT96-09 is based on nine accelerator mass spectrometry (AMS) radiocarbon dates for mixed assemblages of planktonic foraminifera (N. pachyderma right-and left-coiling and G. bulloides; Table 2.1). The radiocarbon age of bulk organic carbon in eight samples was also measured, following removal of carbonate material with dilute phosphoric acid and preheating to 150°C to drive off adsorbed water. All radiocarbon dating 16 127°W 126°W 125°W Fig. 2.1. The study area is located off the west coast of Vancouver Island, British Columbia, Canada (Inset). Sediment cores were collected from Station JT96-09 that is located on the continental slope at a water depth of920m. 17 was conducted at the Lawrence Livermore National Laboratory. Results for the planktonic foraminifera were converted from radiocarbon to calendar years using CALIB 4.3 (Stuiver et al., 1998) assuming a reservoir age of 800 years for radiocarbon ages younger than 12 kyr and 1100 years for those dates that are older (Kienast and McKay, 2001). It should be noted that the age model for the Holocene used here is slightly different than that in Kienast and McKay (2001) as a result of the new estimate of sediment loss during coring. Unless otherwise stated all dates presented in this paper are in calendar years B.P. Alkenone-based palaeothermometry conducted on Core JT96-09 yielded evidence of rapid sea-surface temperature (SST) fluctuations of during the last deglacial (Fig. 2.2a; Kienast and McKay, 2001). The pattern of deglacial SST fluctuations is remarkably similar to temperature fluctuations in the GISP-2 ice core record suggesting that the Bolling-Allerod and Younger Dryas events are recorded in Core JT96-09 (Kienast and McKay, 2001). More importantly these events appear to be nearly synchronous with those in GISP-2 and this allows the palaeo-records from Core JT96-09 to be placed in a more global context. However, the match between the GISP-2 and JT96-09 records is not perfect (see Kienast and McKay, 2001). Offsets between the two records, in some instances in the order of hundreds of years, most probably reflect errors in the age model of Core JT96-09 resulting from: i) problems inherent to radiocarbon dating (e.g., 14C plateaus); ii) a poorly known reservoir age; iii) bioturbation, although the effects of this are limited by the high sedimentation rates during the deglacial; and iv) large errors associated with the radiocarbon dating of small samples. Sedimentation rates, that were calculated by assuming linearity between age picks, range from 5 cm/kyr during the Holocene to between 14 and 169 cm/kyr during the deglacial. The sedimentation rate in the late glacial is ~47 cm/kyr. Mass accumulation rates (MARs) were calculated as the product of linear sedimentation rate (cm/kyr) and sediment dry bulk density (g/cm3). The latter was determined using chlorinity data and a pore water salinity of 35 to estimate porosity, and by assuming an average grain density of 2.5 g/cm3. 18 Core JT96-09 was subsampled at 5 cm intervals for geochemical analysis, yielding a resolution of 1200 years for the Holocene and from <100 to 350 years for the period from 10 to 16 kyr B.P. Samples were prepared for geochemical analysis by freeze-drying followed by hand grinding with an agate pestle and mortar. Total carbon (TC) and total nitrogen (TN) were measured by high temperature combustion using a Carlo Erba 1500 CNS analyzer. Two marine sediment standards (PACS-1 and MESS-1; National Research Council of Canada) were analyzed with each batch of samples and yielded relative standard deviations (RSD, la) of 3 % and 5 % for carbon and nitrogen, respectively. The amount of carbonate carbon (i.e., inorganic carbon, IC) was measured by coulometry. Results for a calcium carbonate standard yielded an RSD of ~1 % (i.e., 11.93 ± 0.13, n=141). The amount of organic carbon (OC) was calculated by difference (OC = TC - IC) and the aggregate RSD for these data is ~4 %. Stable isotope data were obtained by continuous-flow mass spectrometry using a Fisons NA 1500 elemental analyzer attached to a VG Prism mass spectrometer. Samples for carbon isotopic analysis of organic matter (513Corg) were pretreated with 10% HC1 to remove carbonate material and then dried at 50°C overnight. These samples were not washed with distilled water prior to drying. Nitrogen isotope results (815N) were obtained for untreated bulk sediment samples. Data are reported in the standard 8-notation relative to VPDB for carbon and atmospheric N2 for nitrogen. The isotopic results for an in-house sediment standard are ± 0.1 %o for carbon and ± 0.2%o for nitrogen. Repeat analyses of samples are generally better than this. Total Ba and Al concentrations were measured by X-ray fluorescence (XRF) according to the method of Calvert (1990). The concentration of biogenic barium was calculated using Equation 2.1, following the method of Dymond et al. (1992): Bio-barium = Total Ba - (Al x Ba/A\m) (2.1) 19 A Ba/Al lithogenic ratio (Ba/Aliith) of 0.0027 was employed. This value is significantly lower than the average Ba/Aliith ratio of crustal rocks (0.0075; Dymond et al., 1992); however, the latter can vary greatly depending on the composition of the sediment source (e.g., 0.005 to 0.010; Taylor and McLennan, 1985). The geology of Vancouver Island is complex and includes a mixture of metamorphic, igneous and sedimentary rock types (Yorath and Nasmith, 1995); thus it is not possible to chose a Ba/A\m ratio representative of any particular rock type. Rather, the Ba/Aluth ratio used here (i.e., 0.0027) was determined from an exponential regression of the Ba/Al ratios of surface sediments versus water depth (Chapter 4). This method assumes that the fraction of biogenic Ba increases seaward and that close to land (i.e., ~0 m water depth) all of the Ba is terrigenous in origin. The low Ba/Allith estimated using this regression method was confirmed by chemically extracting the bio-barium using a 2M solution of NH4CI and then measuring the barium content of the residue (i.e., the lithogenic Ba) by XRF (Chapter 4). While the average Ba/Aliith ratio is slightly higher (0.0033) using this method, possibly due to incomplete dissolution of the bio-barium, the results show that there is no change in the ratio over the last 16 kyr B.P. Biogenic silica (i.e., opal) content was determined by alkaline dissolution following the procedure of Mortlock and Froelich (1989). The concentration of Si was measured by spectrophotometry and then converted to wt.% opal by multiplying by 2.4 (this assumes 10% water in opal; Mortlock and Froelich, 1989). The RSD (la) for two in-house standards (SNB and JV5) and repeat samples was <4 %. However, the precision of the method is much poorer at opal concentrations <10 % (R. Ganeshram, pers. comm.), possibly due to the dissolution of volcanic glass and clay minerals. Organic geochemical analysis was carried out using the solvent extraction method of Villanueva et al. (1997b). The organic fraction was extracted from 3 to 10 mg of freeze-dried sediment using dichloromethane (CH2C1) followed by saponification with a mixture of 6 % KOH in methanol. The organic fraction was then extracted with «-hexane and run through a silica column. The fraction containing the rc-alkanes and C37 alkenones was eluted 20 from the column with a mixture of dichloromethane and n-hexane. Analyses were made by manual-injection on an HP 5880 gas chromatograph. The n-alkanes and C37 alkenones were identified using an internal standard (i.e., mixture of C]9, C36 and C40), and the identification of C37 alkenones was confirmed by GC-MS. The concentrations of C29 «-alkanes and C37 alkenones were calculated by assuming that their concentrations were proportional to the chromatogram peak area and that the response factor was the same as for the internal standard. Only the di- and tri-unsaturated C37 alkenones were quantified, as the amount of tetra-unsaturated C37 alkenone was generally below the detection limit of 10 ng. 2.3 Results 2.3.1 Radiocarbon data for bulk organic carbon The age of bulk organic matter ranges from 9920 to 24500 14C years (Table 2.1). If most of the organic carbon in a sample is of marine origin and formed contemporaneously with planktonic foraminifera shells, the difference between the organic carbon and shell ages should be small. The greater the input of old organic carbon, the larger this difference will be. For the two youngest samples the radiocarbon age difference is less than 2000 14C years, but for all other samples it is substantially higher (3435 to 12370 14C kyrs; Table 2.1). These results suggest that between 12.3 and 16.0 kyr B.P. there was a large amount of old organic matter being deposited off Vancouver Island. Assuming that all of the old organic carbon is of infinite radiocarbon age we estimate that between 22 and 51 % of the total organic carbon present in late glacial and early deglacial sediments is reworked material. These are however minimum estimates because a portion of the old organic carbon may not be of infinite age. 2.3.2 Organic carbon content 21 Table 2.1. Radiocarbon data and estimates of the input of organic matter of infinite radiocarbon age for Composite Core JT96-09. Sample 1 Calendar Age Radiocarbon Age (yrs) Radiocarbon % Old Organic % Terrigenous (kyrs B.P.) Age Difference4 Matter5 Organic Matter6 Measured Age Measured Age Estimated Age Measured Age of of Planktonic Forams of Benthic Forams of Planktonic Forams Bulk Organic Matter 38-40 mc 8.23 - 9830± 110 9030 J 9920 890 9 14 35-36 10.03 2 9760 ±60 - - - - -36-37 10.10 - - 9760 11380 1620 14 17 65-66 12.24 2 11210 ± 120 - - - -66-67 12.29 - - 11210 14790 3580 24 35 75-76 12.73 2 11500 ± 110 - - - - - -90-91 12.84 2 11600 ± 80 - - - - - -96-97 13.04 - - 12225 15660 3435 22 24 100-101 13.172 12460 ± 120 - - - - - -130-131 13.43 2 12640 ±90 - - - - - -165-166 13.55 13210± 150 166-167 13.55 - - 12110' 24480 12370 51 54 265-266 14.142 13410±80 -266-267 14.15 - - 13410 24500 11090 45 55 290-291 14.30 2 13520 ±70 - - - - - -350-351 15.57 2 14140 ±70 - - - - -351-352 15.59 - - 14150 21440 7290 34 57 370-371 16.01 - - 14525 20320 5795 29 59 1 Samples taken from the multicore are labelled with mc. 2 Calendar ages used in the creation of an age model for Composite Core JT96-09. These data were first published in Kienast and McKay (2001). ' Planktonic ages for samples 38-40mc and 166-167 were estimated by subtracting the reservoir age from the benthic foraminfera age. For samples younger than 12000 MC years a reservoir correction of 800 years was applied while for older ages a correction of 1100 years was used. All other planktonic ages were estimated using the measured planktonic age of the overlying sample. 4 The radiocarbon age difference is the difference between the estimated age of planktonic forams and the age of bulk organic carbon. 5 Old organic matter refers to organic matter of infinite radiocarbon age. It is calculated using the following formula: % Old carbon = (Radiocarbon age difference / Radiocarbon age of bulk organic carbon)*100 6 % Terrigenous Organic matter is calculated using Equation 2.2 (i.e., 8"C„^ data). See the text for a further discussion. 22 Organic carbon (OC) values range from 0.07 to 3.07% (Table 2.2). Sediments deposited between -11 kyr B.P. and the present have the highest OC content while late glacial and deglacial sediments are characterized by concentrations <1 % (Fig. 2.2b). Due to the large variations in sedimentation rate it was necessary to calculate the mass accumulation rate (MAR) of organic carbon. MAR values are relatively low (0.09 to 0.21 g/cm2/kyr) prior to 14.3 kyr B.P. and after 12.7 kyr B.P., and substantially higher (>0.5 g/cm2/kyr) from 14.3 to 12.7 kyr (Fig. 2.2b). High organic carbon accumulation occurred during a period of high sea-surface temperature off Vancouver Island that has been correlated to the Boiling-Allerod (Fig. 2.2a; Kienast and McKay, 2001). The data also suggest a slight increase in the MAR of total organic carbon at the Pleistocene-Holocene boundary (10.8 to 9.2 kyr B.P.), also a time of warmer sea-surface temperatures (Kienast and McKay, 2001). Interestingly, MAR values during the late glacial are generally higher than MAR values in the Holocene. 2.3.3 Organic carbon/nitrogen ratios Organic carbon to total nitrogen weight ratios (OC/TN) range from 7.0 to 11.2 (Table 2.2) with an average value of 9.4. The highest ratios (>10) are observed in deglacial sediments (i.e., 14.6 to 11.6 kyr B.P.; Fig. 2.3). In general, the relatively low OC/TN values suggest the dominance of marine organic carbon; however, as will be discussed later these data are misleading. 2.3.4 Carbon- and nitrogen-isotope data 13 The carbon isotopic composition of organic matter (8 Corg) in Core JT96-09 varies by 4 %o (-25.2 to -21.1 %o; Table 2.2) while the nitrogen isotopic composition of the bulk sediment (815N) varies by almost 6 %o (+2.9 to +8.9 %o; Table 2.2). As seen in Figure 2.2c, both 813Corg and 815N values are low during the late glacial to early deglacial, particularly 23 Fig. 2.2. Geochemical data for Core JT96-09. a) Alkenone-derived sea-surface (SST) temperature data from Kienast and McKay (2001). SST data are provided in Table 2.1. The precision for SST measurements is ±1.5 °C. Arrows indicate the position of radiocarbon dates used in the creation of the age model, b) Total organic carbon concentration (OC, open squares) and mass accumulation rate of organic carbon (MAR, solid squares). The precision of organic carbon measurements is 4 %. Note, multicore samples are labeled with the letters mc. c) Carbon-isotopic data for organic matter (813Corg, solid circles) and nitrogen-isotopic data for the bulk sediment (815N, open circles). The errors are ±0.1 %o and ±0.2 %o (C- and N-isotopic results, respectively), d) Biogenic barium concentration (crosses) and mass accumulation rate of biogenic barium (MAR, squares). The error associated with the estimation of biogenic barium concentrations is relatively large (RSD = 30%). e) Opal concentration (open triangles) and mass accumulation rate of opal (MAR, solid triangles). Opal concentration data have a RSD of 4 %. f) Concentration profile of the terrestrial biomarker C2a n-alkane normalized to total organic carbon (C2Q/OC, solid diamonds) and concentration profile of the marine biomarker C37 alkenone normalized to total organic carbon (C37/OC, open diamonds). The linear sedimentation rate during the Boiling (-14.3 to 13.5 kyr B.P.) is 169 cm/kyr, during the Allerod (-13.5 to 12.7 kyr B.P.) it ranges from 20 to 143 cm/kyr, and by the Holocene it has decreased to 5 cm/kyr. 24 2.2a) 2.2b) 2.2c) Sea Surface Temperature (°C) Organic Carbon (OC; wt.%) 813Corg (permil vs PDB) -28 -26 -24 -22 -20 MAR Organic Carbon (g/cm2/kyr) 8 N (permil vs air) 2.2d) 2.2e) 2.20 Bio-barium (ppm) 0 100 200 300 Opal (wt. %) 2 4 6 8 10 12 10 12 C29n-alkane/organic carbon (C29/OC; ug/g) 0 50 100 150 200 MAR bio-barium (g/cm2/kyr) MAR Opal (g/cm2/kyr) 0 20 40 60 80 alkenone/organic carbon (C37/OC; ug/g) 25 Table 2.2. Geochemical data for Composite Core JT96-09. Sample 1 Corrected Calendar Age SST! 5"N 5"Coig TN' TC IC OC OC/TN4 Total Al Total Zr Total Ba Bio-barium Opal CJC/I -alkalies Alkenones Depth (cm) (kyrs B.P.) TO (petmil) (permil) (wt.%) (wt.%) (wt %) (wt. %) (%) (ppm) (ppm) (ppm) (wt. %) ng/g dry sed ng/g dry sed. 0-1 mc 0.5 0.11 6.74 -21.28 0.35 3.20 0.13 3.07 8.7 13.61 146 624 257 7.0 0-2 mc 1.0 0.21 9.83 290.3 1110.3 1-2 mc 1.5 0.32 0.36 3.16 0.11 3.05 8.4 2-3 mc 2.5 0.53 0.37 3.17 0.10 3.07 8.3 2-4 mc 3.0 0.63 10.19 309.7 636.9 3-4 mc 3.5 0.74 0.36 3.07 0.08 2.99 8.4 4-5 mc 4.5 0.95 0.36 3.09 0.06 3.03 8.4 4-6 mc 5.0 1.06 10.23 5-6 mc 5.5 1.16 7.03 -21.19 0.35 3.03 0.07 2.96 8.5 14.74 139 624 226 6.7 6-8 mc 7.0 1.48 0.35 3.01 0.04 2.97 8.5 8-10 mc 9.0 1.90 10.27 0.35 3.01 0.04 2.97 8.5 357.8 979.9 10-12 mc 11.0 2.32 10.83 6.66 -21.19 0.31 2.80 0.02 2.78 8.9 15.38 151 617 202 6.8 318.6 682.5 12-14 mc 13.0 2.74 10.67 0.33 2.84 0.04 2.80 8.6 203.6 465.9 14-16 mc 15.0 3.17 9.72 6.81 -21.17 0.31 2.75 0.04 2.71 8.7 14.92 153 633 230 5.8 265.1 548.9 16-18 mc 17.0 3.59 10.34 0.32 2.79 0.04 2.75 8.6 214.1 357.9 6-7 18.5 3.91 10.34 7.33 -21.13 0.30 2.67 0.10 2.57 8.5 14.94 160 652 249 6.6 237.7 534.7 18-20mc 19.0 4.01 9.92 0.23 2.11 0.04 2.07 8.9 249.4 294.1 20-22 mc 21.0 4.43 6.50 -21.39 0.20 1.96 0.05 1.91 9.4 14.01 155 598 220 4.3 22-24 mc 23.0 4.86 9.69 0.19 1.79 0.09 1.70 9.1 276.4 561.3 11-12 23.5 4.96 6.95 -21.55 0.21 2.02 0.09 1.94 9.2 15.10 160 576 168 4.7 24-26 mc 25.0 5.28 7.71 -21.67 0.18 1.79 0.14 1.65 9.3 15.64 156 596 174 4.4 26-28 mc 27.0 5.70 9.50 0.18 1.83 0.22 1.61 9.0 319.9 367.1 16-17 28.5 6.02 10.96 8.17 -21.56 0.20 2.08 0.26 1.82 9.1 4.5 169.2 326.5 28-30 mc 29.0 6.12 10.53 0.18 1.90 0.25 1.64 9.0 214.9 406.3 30-32 mc 31.0 6.54 10.45 7.35 -21.64 0.18 1.91 0.28 1.63 9.2 15.81 155 608 181 4.5 280.9 660.8 19-20 31.5 6.65 10.16 412.3 1025.4 32-34 mc 33.0 6.97 0.19 1.99 0.29 1.70 9.1 21-22 33.5 7.07 7.77 -21.92 0.21 2.29 0.28 2.01 9.4 15.67 148 631 208 5.7 34-36 mc 35.0 7.39 9.77 7.93 -21.63 0.18 2.01 0.28 1.72 9.4 15.72 147 591 167 4.9 383.4 349.9 36-38 mc 37.0 7.81 10.81 0.20 2.08 0.29 1.79 9.1 265.6 492.0 26-27 38.5 8.13 8.33 -21.75 0.22 2.26 0.22 2.04 9.2 15.90 4.7 38-40 mc 39.0 8.23 7.45 -21.85 0.19 2.07 0.32 1.76 9.2 16.16 151 618 182 4.8 31-32 43.5 9.18 11.11 8.38 -22.04 0.20 2.23 0.33 1.90 9.3 14.83 138 640 240 5.9 322.8 483.0 36-37 48.5 10.10 11.94 8.61 -22.02 0.20 2.46 0.59 1.87 9.1 15.18 5.9 228.2 314.0 39-40 51.5 10.32 10.20 349.4 625.2 41-42 53.5 10.47 12.38 8.34 -21.96 0.20 2.55 0.65 1.90 9.6 15.20 135 642 232 6.8 164.5 287.6 44-45 56.5 10.69 10.85 456.3 570.6 46-47 58.5 10.84 12.01 8.89 -21.99 0.19 2.41 0.53 1.87 9.7 14.87 7.1 229.3 317.5 48-49 60.5 10.99 9.79 352.3 667.4 48.5-50 61.2 11.04 10.38 215.5 382.9 51-52 63.5 11.21 6.93 7.10 -22.19 0.12 1.47 0.30 1.17 9.4 15.41 136 618 202 4.8 111.1 182.5 53-54 65.5 11.35 6.46 325.7 421.8 56-57 68.5 11.57 5.93 6.86 -22.67 0.07 0.93 0.22 0.71 10.0 15.64 134 600 178 4.8 108.1 84.3 58-59 70.5 11.72 5.% 180.5 381.5 61-62 73.5 11.94 6.63 6.23 -23.11 0.06 0.80 0.27 0.53 8.7 15.90 140 599 170 3.7 108.7 169.5 66-67 78.5 12.29 6.99 5.89 -23.08 0.05 0.77 0.23 0.55 11.2 14.51 4.1 96.0 126.2 71-72 83.5 12.54 7.07 6.21 -23.53 0.04 0.73 0.33 0.39 10.8 13.92 212 555 179 4.3 116.3 109.8 76-77 88.5 12.74 8.00 7.41 -22.54 0.09 1.20 0.35 0.85 9.8 16.39 127 718 275 4.7 209.7 178.7 81-82 93.5 12.78 8.89 7.75 -22.75 0.10 1.38 0.48 0.90 9.3 15.93 120 715 285 5.2 169.2 477.6 86-87 98.5 12.81 8.80 7.20 -22.42 0.09 1.21 0.32 0.89 9.6 16.27 120 720 281 6.1 180.1 196.9 91-92 103.5 12.87 7.54 7.52 -22.91 0.09 1.30 0.48 0.82 9.1 16.97 118 724 266 4.8 228.0 215.3 96-97 108.5 13.04 8.41 7.38 -22.44 0.09 1.28 0.45 0.84 9.8 16.23 121 704 266 5.7 209.0 288.1 101-102 113.5 13.18 9.10 7.36 -22.75 0.10 1.34 0.49 0.85 8.9 16.37 122 736 294 170.8 200.5 106-107 118.5 13.22 8.50 7.28 -22.69 0.08 1.24 0.41 0.83 9.9 16.30 121 677 237 5.3 247.8 262.0 111-112 123.5 13.27 8.81 7.10 -23.16 0.08 1.18 0.42 0.76 9.1 14.67 121 676 280 4.6 158.9 164.4 116-117 128.5 13.31 8.95 6.77 -22.87 0.08 1.15 0.41 0.74 9.7 16.48 123 676 231 5.8 165.4 172.3 121-122 133.5 13.36 9.06 7.37 -23.01 0.09 1.16 0.34 0.82 9.0 16.03 124 683 250 5.3 261.1 212.7 126-127 138.5 13.40 8.67 6.75 -23.01 0.08 1.10 0.36 0.74 9.5 16.42 125 696 253 4.7 173.2 276.3 131-132 143.5 13.44 7.42 7.23 -23.26 0.09 1.12 0.36 0.77 9.0 16.28 128 608 168 5.6 183.4 304.5 136-137 148.5 13.47 8.98 4.92 -23.75 0.02 0.45 0.23 0.22 9.1 14.19 172 539 156 79.7 64.2 141-142 153.5 13.47 7.02 -23.39 0.04 0.76 0.31 0.45 10.2 15.49 226 598 180 206.2 273.8 146-147 158.5 13.47 2.93 -23.32 0.01 0.17 0.11 0.07 6.8 13.26 165 486 128 16.6 6.5 151-152 163.5 13.47 4.79 -24.24 0.05 0.77 0.21 0.56 11.0 14.57 131 512 119 752.3 71.3 156-157 152.5 13.49 7.9.1 4.84 -24.22 0.05 0.74 0.23 0.52 10.1 17.03 131 569 109 4.0 343.2 174.5 161-162 157.5 13.52 4.22 -25.20 0.04 0.73 0.41 0.32 8.3 15.94 173 635 205 4.2 224.0 42.3 166-167 162.5 13.55 6.82 5.00 -24.26 0.05 0.78 0.28 0.50 9.7 16.57 138 556 109 265.8 119.3 171-172 167.5 13.58 6.91 4.81 -24.64 0.05 0.71 0.20 0.51 9.7 16.68 124 526 76 3.6 297.2 52.2 176-177 172.5 13.61 6.86 3.98 -24.62 0.05 0.67 0.19 0.49 10.2 16.61 127 522 74 466.3 34.5 181-182 177.5 13.64 4.33 -25.05 0.04 0.75 0.48 0.27 6.9 16.55 167 673 226 4.6 285.6 31.1 186-187 182.5 13.67 7.30 4.57 -24.51 0.05 0.70 0.21 0.48 10.5 16.74 126 547 95 240.7 62.7 191-192 187.5 13.70 7.52 4.44 -24.86 0.05 0.65 0.18 0.46 9.7 16.70 125 520 69 3.4 192.7 40.2 196-197 192.5 13.73 8.41 4.21 -24.79 0.05 0.67 0.15 0.51 11.0 16.68 119 513 63 273.3 71.6 201-202 197.5 13.76 8.37 4.11 -24.68 0.05 0.67 0.16 0.51 9.8 16.70 121 516 65 3.5 282.4 58.1 206-207 202.5 13.79 9.06 3.72 -24.52 0.04 0.64 0.24 0.40 10.2 16.53 137 561 115 251.1 74.4 211-212 207.5 13.82 8.73 4.38 -24.69 0.05 0.68 0.18 0.50 9.5 16.93 126 494 37 3.6 218.0 93.6 216-217 212.5 13.85 10.29 3.50 -24.71 0.05 0.66 0.15 0.51 10.9 16.86 120 496 41 249.5 80.2 221-222 217.5 13.88 9.70 4.06 -24.81 0.06 0.71 0.16 0.55 9.9 16.93 125 490 33 3.7 206.8 78.8 226-227 222.5 13.91 9.99 3.81 -24.60 0.05 0.66 0.15 0.50 10.7 16.92 119 500 43 270.7 63.6 231-232 227.5 13.94 10.12 4.03 -24.68 0.05 0.63 0.17 0.46 9.5 16.44 133 486 42 3.4 286.9 67.3 236-237 232.5 13.97 7.87 3.85 -24.47 0.05 0.73 0.16 0.57 11.1 16.82 119 523 69 258.2 69.6 241-242 237.5 14.00 10.34 4.29 -24.62 0.05 0.68 0.14 0.53 10.1 15.92 124 464 34 3.4 242.1 44.9 246-247 242.5 14.03 9.75 4.31 -24.42 0.05 0.74 0.25 0.49 10.2 16.85 131 565 110 420.2 106.1 continued on next page 26 Table 2.2. (continued) Sample1 Corrected Calendar Age SST! 8"N 8"Corg TN' TC' IC' OC OOTN" Total Al Total Zr Total Ba Bio-barium Opal Q0 n -alkane! Alkenones Depth (cm) (kyrs B.P.) (°C) (permil) (permil) (wt.%) (wt. %) (wt. %) (wt. %) (%) (ppm) (ppm) (ppm) (wt %) ng/g 251-252 247.5 14.06 9.59 3.96 -24.59 0.05 0.59 0.16 0.44 9.7 16.83 135 466 12 3.3 231.2 45.9 256-257 252.5 14.09 7.71 3.85 -24.49 0.05 0.69 0.17 0.52 10.4 16.89 122 524 68 221.1 86.6 261-262 257.5 14.12 6.38 5.08 -24.49 0.05 0.75 0.20 0.55 10.3 16.55 136 515 68 3.6 205.4 96.7 266-267 262.5 14.15 6.51 4.72 -24.27 0.05 0.76 0.20 0.56 10.5 16.93 127 539 82 218.0 139.2 271-272 267.5 14.18 8.03 4.91 -24.74 0.05 0.74 0.33 0.41 • 8.9 16.34 156 572 131 4.1 146.0 70.0 276-277 272.5 14.21 4.50 -24.41 0.05 0.78 0.29 0.49 10.0 16.39 142 572 129 281-282 277.5 14.25 5.99 4.85 -24.71 0.06 0.81 0.27 0.53 9.5 16.21 139 525 87 3.8 258.0 64.5 286-287 282.5 14.28 5.03 -24.24 0.05 0.92 0.37 0.55 10.3 16.20 138 556 119 291-292 287.5 14.33 6.34 5.44 -24.37 0.06 0.91 0.40 0.51 9.2 16.27 150 561 122 4.4 271.2 124.7 296-297 292.5 14.43 5.14 -24.15 0.06 0.96 0.39 0.57 10.2 16.38 133 544 102 301-302 297.5 14.54 5.70 4.93 -24.60 0.05 0.88 0.37 0.50 9.2 16.37 144 547 105 3.5 249.5 60.0 306-307 302.5 14.60 4.91 -24.43 0.05 0.84 0.27 0.57 10.3 16.58 128 536 88 311-312 307.5 14.75 5.48 5.58 -24.28 0.05 0.93 0.50 0.43 8.8 15.12 172 523 115 4.1 166.5 82.2 316-317 312.5 14.85 5.38 -24.07 0.04 0.83 0.46 0.36 8.8 15.54 185 618 198 321-322 317.5 14.96 8.58 5.31 -24.30 0.04 0.80 0.41 0.39 9.0 15.85 184 595 167 4.9 180.1 79.4 326-327 322.5 15.06 5.16 -24.23 0.04 0.78 0.42 0.36 9.1 16.04 178 644 211 331-332 327.5 15.17 9.69 4.45 -24.42 0.04 0.75 0.46 0.29 7.7 15.05 203 562 156 4.8 138.0 30.1 336-337 332.5 15.27 5.08 -24.36 0.04 0.75 0.45 0.31 8.7 15.53 187 646 227 341-342 337.5 15.38 6.42 5.59 -24.09 0.04 0.82 0.44 0.38 8.6 16.25 186 573 134 4.4 184.4 53.8 346-347 .342.5 15.48 5.07 -24.09 0.04 0.89 0.52 0.37 9.4 15.41 351-352 347.5 15.56 10.39 4.98 -24.41 0.04 0.74 0.45 0.29 7.7 15.32 202 587 17.3 4.6 184.9 43.2 356-357 352.5 15.69 4.58 -24.64 0.03 0.65 0.41 0.24 8.7 14.99 209 602 197 361-362 357.5 15.80 7.35 5.81 -24.10 0.05 0.80 0.39 0.40 8.7 14.07 185 581 201 4.7 134.7 68.4 366-367 362.5 15.90 5.65 -24.22 0.04 0.74 0.42 0.32 9.0 15.68 190 628 205 371-372 367.5 16.01 8.30 4.66 -24.52 0.03 0.67 0.41 0.27 7.7 14.84 195 551 150 4.7 1.36.5 44.5 1 Sample taken from the multicore are labelled with mc. ; Sea-surface temperature (SST) record is from Kienast and McKay (2001). 3 TN = total nitrogen; TC = total carbon; IC = inorganic carbon; OC = organic carbon. * OC/TN are weight, not atomic, ratios. 27 Organic Carbon/Total Nitrogen (OC/TN, weight ratio) 5 7 9 11 13 f'T I 'I I | I I I I | I I 1 I | I I I I | 0 50 100 150 200 C n-alkanes/Organic carbon (C /OC;ug/g) Fig. 2.3. Downcore profiles of the organic carbon/total nitrogen ratio (OC/TN) and C29 n-alkane/organic carbon ratio (C29/OC). 28 between 14.3 and 13.5 kyr B.P. (i.e., the Boiling). Isotopic values abruptly increase at 13.5 kyr B.P. (i.e., the beginning of Allerod), remain high throughout the Allerod, decrease in the Younger Dryas (-12.7 to 11.0 kyr B.P.) and rise again at the close of the Younger Dryas (Figs. 2.2a and 2.2c). The 815N values peak at -10.8 kyr B.P. and then slowly decrease throughout the Holocene while 813Corg values increase slightly through the Holocene. 813Corg and 815N exhibit a very strong positive correlation in late Pleistocene sediments, but are negatively correlated in Holocene deposits (Fig. 2.4). 2.3.5 Biogenic barium and opal data The concentration of biogenic barium (i.e., bio-barium) ranges from 11 to 294 ppm (Table 2.2). Values are lowest (generally <100 ppm) between 14.3 and 13.5 kyr B.P., when sedimentation rate is highest, and are highest (>200 ppm) between 13.5 and 12.7 kyr B.P. (Fig. 2.2d). Interestingly, Holocene and late deglacial sediments have a similar bio-barium content (Fig. 2.2d). The highest bio-barium mass accumulation rates (> 0.02 g/cm2/kyr) occur between 14.3 and 12.5 kyr B.P. (Fig. 2.2d) contemporaneously with a time of high SST that is believed to be the Bolling-Allerod. Opal content ranges from 3.3 to 7.1 wt.% (Table 2.2). Concentrations are highest at -13.0 kyr B.P., between 11.0 and 10.0 kyr B.P., and in late Holocene sediments (Fig. 2.2e). Opal mass accumulation rates, which range from 0.16 to 10.17 g/cm /kyr are relatively low during the late glacial, high between 14.3 and 13.5 kyr B.P. and very low during the Holocene (Fig. 2.2e). 2.3.6 Biomarker data Long chain n-alkanes (C25 to C3i) with an odd-carbon number predominance are present in all samples analyzed and the distribution is similar to that observed in Columbia 29 9-i'-6-5: 4-• Holocene • Late Pleistocene • /Il y =-26.048+ -1.5538x R2 = 0.54 • y = 40.248 + 1.4535x R2 = 0.89 -26 -25 -24 -23 -22 -21 -20 5 Corg (permil, PDB) Fig. 2.4. N-isotopic composition of the bulk sediment (515N) versus the C-isotopic composition of organic matter (513Corg). Note the positive correlation between S^N and S^Cc-g for Late Pleistocene samples and the negative correlation for Holocene samples. Two clusters of Late Pleistocene samples are observed. Those samples characterized by S^N values of < -24%0 and 5i3Corg values of < +6%o are older than 13.5 kyr B.P. 30 River sediments with a peak at C29 (Prahl et al., 1994). The concentrations of C29 «-alkane in Core JT96-09 range from 17 to 46 ng/g dry sediment (Table 2.2). To circumvent the problem of varying amounts of dilution by inorganic material, a reflection of the highly variable sedimentation rate over the last 16 kyr, the concentration of C29 n-alkanes has been normalized to the organic carbon concentration. The C29/OC ratios, which range from 9 to 132 p.g/gC, are high during late glacial and early deglacial, particularly between 14.3 and 13.5 kyr B.P. (i.e., the Boiling; Fig. 2.2f). At 13.5 kyr B.P. C29/OC ratios abruptly drop, then decrease more slowly until ~11 kyr B.P. and remain relatively low after this time (Fig. 2.2f). 13 15 High C29/OC ratios correlate very well with low 8 C and 8 N values (Figs. 2.5a and b). The concentration of C37 alkenones ranges from 7 to 1110 ng/g dry sediment (Table 2.2) and the C37/OC ratio ranges from 7 to 61 ug/gC. C37/OC ratios are low in the glacial and early deglacial, when C29 n-alkane content is high, and then increase abruptly at 13.5 kyr B.P. and remain high until ~11.9 kyr (i.e., the Allerod and early part of the Younger Dryas; Fig. 2.2f). C37/OC ratios decrease from 11.6 to 10.1 kyr B.P. and only rise slightly in the Holocene (Fig. 2.2f). In general, Holocene C37/OC ratios are only slightly higher than glacial values. 2.4 Discussion 2.4.1 Organic carbon The mass accumulation rate of total organic carbon off Vancouver Island was highest during the last deglacial. While this is primarily a reflection of the very high sedimentation rates at this time, there is also an increase in the absolute concentration of organic carbon during the latter half this period (i.e., the Allerod; Fig. 2.2b). The data also suggest that the accumulation of organic carbon was higher during the late glacial (16.0 to 14.3 kyr B.P.) than during the Holocene. These results are contrary to what is expected based on climate 31 2.5a) Fig. 2.5. C29 H-alkane normalized to organic carbon (C29/OC) versus a) the C-isotopic composition of organic matter (813Corg) and b) the N-isotopic composition (8I5N) of the bulk sediment. 32 modeling and palaeoproductivity data from other regions within the California Current System. Climate modeling suggest that during the last glacial the Aleutian Low pressure system dominated the Northeast Pacific region (COHMAP, 1988). The situation should have been similar to present-day winter conditions when winds are from the south and thus do not drive upwelling. As a result of this seasonal switch in atmospheric forcing, modern productivity off Vancouver Island is indeed lower in winters (Thomas et al., 1994). The same situation should have held true for the last glacial, and this suggestion is supported by geochemical and palaeontological studies that indicate that coastal upwelling and marine productivity were significantly reduced as far south as central California during the last glacial maximum (Dymond et al., 1992; Lyle et al., 1992; Sancetta et al., 1992; Ortiz et al, 1997; Dean and Gardner, 1998; Mix et al., 1999). Radiocarbon data indicate that a significant percentage (22 to 51 %; Table 2.1) of the organic matter deposited during the late glacial and early deglacial is, in radiocarbon terms, of infinite age. Therefore, bulk organic carbon data cannot be used to infer changes in palaeoproductivity. Instead, it is necessary to determine and correct for the amount of old organic matter in each sample. To do this we first determine whether the old organic matter is of marine and/or terrestrial origin. The ratio of organic carbon to total nitrogen (OC/TN) is commonly used to determine whether the dominant source of organic matter in sediments is terrestrial or marine. Low OC/TN ratios (<10) are often cited as evidence that marine organic matter dominates while higher ratios (>20) indicate the presence of abundant terrestrial organic matter. If our interpretation was based solely on OC/TN ratios, which range from 6.9 to 11.2, we would conclude that much of the organic matter in Core JT96-09 is marine, except for a slight relative increase in the terrestrial content during the deglacial (Fig. 2.3). However, this interpretation conflicts with n-alkane results. A^-alkanes are organic compounds that occur in epicuticular waxes of higher terrestrial plants (Eglinton and Hamilton, 1967) and they are commonly used as a terrestrial biomarker. When the C29 n-alkane concentration is 33 normalized to organic carbon, to correct for problems of variable dilution by inorganic materials, it can be seen that a relatively larger fraction of the total organic carbon deposited prior to -13.5 kyr B.P. is terrestrial (Fig. 2.3). It is rather surprising that high C29/OC ratios are not accompanied by OC/TN ratios substantially greater than 11. Contamination by inorganic and microbial nitrogen can lead to anomalously low OC/TN ratios in organic-poor sediments (Meyers, 1997). However, in Core JT96-09 the concentration of inorganic N is quite low (~ 0.0025 %) and does not significantly influence OC/TN values. The unusually low OC/TN numbers are also not the result of post-depositional alteration of organic matter since this tends to attack the N-rich compounds, thus increasing OC/TN values. However, not all types of terrigenous organic matter are characterized by high OC/TN ratios. Fine grained terrigenous organic detritus has OC/TN ratios similar to those of marine organic material (7 to 15; Hedges and Oades, 1997). It is quite possible, given the fine-grained nature of the sediments in Core JT96-09, that this accounts for the relatively low OC/TN values observed. We conclude that OC/TN ratios are of little use in this instance and are even rather misleading. A strong positive correlation between 513Corg and 8I5N values occurs in the late glacial and deglacial sediments (Fig. 2.4). While many factors can influence both carbon and nitrogen isotope values, their positive correlation here suggests that mixing of terrestrial and marine organic matter may provide the primary control on the isotopic composition of these deposits. This interpretation is supported by the strong correlation between C29/OC ratios 13 15 and both 8 Corg and 8 N (Figs. 2.5a and b). The fraction of terrestrial organic matter can be estimated using Equation 2.2 (Jasper and Gagosian, 1993; Minoura et al., 1997; Fernandes and Sicre, 2000). Terrigenous fraction = (S13Csmpi - 813Cmar) / (813Cterr - 813Cmar) (2.2) 34 where 8 Csmpi, 8 Cmar and 8 Cterr are the isotopic compositions of the sample, the marine endmember and the terrestrial endmember, respectively. Application of this equation requires that we know, or can reasonably estimate, 813Cmar and S13Cterr-13 The 8 C composition of marine organic matter ranges from -35 to -16 %o (Goericke and Fry, 1994). Much of this variability can be explained by changes in the concentration of C02 dissolved in seawater, because the 813C of marine organic matter varies inversely with C02 concentration (e.g., a 3-4 umolar increase in C02 roughly equates to a 1-2 %o decrease in 813C; Popp et al., 1989; Rau et al., 1989; Rau et al., 1991). However, other factors including phytoplankton growth rate (Fry and Wainright, 1991; Law et al, 1995; Rau et al., 1997), cell geometry (Popp et al., 1998; Popp et al., 1999), and cell membrane permeability 13 13 (Francois et al., 1994) also influence 8 C. At present, the 8 C of settling particulate organic matter off Vancouver Island ranges from -23.0 to -20.6 %o, with an annual flux-weighted average of -21.9 %o (Wu et al., 1999a). Surface sediments deposited on the continental slope are characterized by similar values (-21.5 to -21.2 %0; Chapter 4) suggesting the isotopic composition of organic matter is not altered as it settles through the water column. Furthermore, there is no change in the 813Corg composition in the upper 20 cm of slope deposits (Chapter 4) so post-depositional diagenesis is of little concern, even in the slowly accumulating Holocene deposits. These data suggest that a S13Cmar value of-21 %o is most applicable in Equation 2.2. Our choice of -21 %o is confirmed using the linear regression equation obtained when 813C values are plotted versus C29/OC (Fig. 2.5a). The 813C of terrestrial organic matter (813Cterr) can be highly variable due to different carbon uptake pathways. Terrestrial C3 plants (e.g., trees, shrubs and cool-climate grasses), that passively transport carbon dioxide into their cells, have relatively low 813C values (-23 %o to -34 %o; Faure, 1986). In comparison, C4 plants (e.g., tropical grasses, sedges and salt marsh plants) that employ a carbon dioxide concentrating mechanism have relatively high 813C values (-19 to -6 %o; Faure, 1986). The 813C value for terrestrial organic matter transported to the Pacific by the Fraser River in southwestern British Columbia is -27 %o 35 (Wu, 1997). This value is similar to estimates of 8 Cterr for Columbia River sediments (-25.7 %o; Prahl et al., 1994) and terrestrial organic matter found in northeast Pacific sediments (-26 %<>; Peters et al., 1978). A 813Cterr value of -27 %o is also consistent with the modern vegetation, which is dominated by C3 plants (i.e., coniferous forests). Pollen data from the Queen Charlotte Islands suggest that prior to the Holocene, grasses and sedges were relatively more important contributors to the terrestrial vegetative cover in the region (Mathewes et al., 1993). Nevertheless, regression of the carbon and nitrogen isotopic data for the late Pleistocene yields a 813Cterr value of -27 %o, assuming a 815Nterr endmember 13 value of l%o, and argues for a dominantly C3 plant source over the last 16 kyr. A 8 Cterr value of -27 %o is therefore used in Equation 2.2. The resulting calculations show that the terrigenous fraction accounts for 50 to 70 % of the total organic sedimentary organic matter pool during the late glacial and early deglacial (i.e., prior to 13.5 kyr B.P.; Fig. 2.6). The terrigenous fraction is approximately half this proportion (20 to 40 %) between 13.4 and 11.2 kyr B.P., and then progressively decreases to only a few percent in the latter part of the Holocene (Fig. 2.6). We have assumed that the endmember values assigned to both 813Cmar and 813Cterr have remained constant for the last 16 kyr. However, during the last glacial a lower atmospheric C02 concentration and thus lower concentration of dissolved C02 in the surface ocean might have led to a higher 813Cmar value (Jasper and Hayes, 1990; Rau et al., 1991). Substituting a value of -19 %o in Equation 2.2 increases the estimate of terrigenous organic matter content by 10 to 20% (late glacial and Holocene, respectively), but the temporal variations remain unchanged. An increase in the 8 Cterr value, due to a higher C4 plant input during the glacial, would have a similar effect. 13 In general, the terrigenous fraction estimates determined using 8 C data are similar to those obtained from the radiocarbon dating of bulk organic matter, assuming an infinite age for the old organic carbon, and suggest that most of the old organic carbon is terrestrial in origin. However, in the late glacial sediments (i.e., older than 15.6 kyr B.P.) the estimates of 36 Terrigenous Fraction Fig. 2.6. The percentage of terrigenous organic carbon (i.e., the terrigenous fraction) in Core JT96-09 estimated using both the 513Corg and S^N data. The following endmember isotopic compositions were employed: 5i3Cterr - -27 %o, 8i3Cmar = -21 %o, 5i5Nterr = +1 %0 and 5'5Nmar= +9 %o. If the endmember isotopic compositions are assumed to have an error of ± 1 %o then the final terrigenous fraction estimates could vary by as much as 20 % but the trends remain unchanged. The values estimated using both the C and N isotope data compare very well for the Late Pleistocene and Early Holocene but diverge significantly for the Mid- to Late Holocene. The values determined using the 815N data are though to be too high in the Holocene because the 815N composition of marine organic matter has not remained constant over time. See the text for further discussion. 37 % terrestrial organic matter are noticeably higher than the estimates of % old organic carbon (Table 2.1). This discrepancy suggests that some fraction of the terrigenous organic matter may not be of infinite age and/or that our estimates of terrigenous input determined using Equation 2.2 are high. Lowering the estimated terrigenous input would require decreasing one (or both) of the 513C endmember values. During the glacial, lower atmospheric C02 concentrations and a switch from C3 to C4 plants would increase, not decrease, 813Qerr values and thus cannot explain the discrepancy in % old organic carbon and % terrigenous estimates. However, if the local concentration of C02 dissolved in seawater was higher during the last glacial as a result of colder sea-surface temperatures and/or lower productivity then it is quite possible that 813Cmar values were lower. A final explanation for the discrepancy could lie in a switch from a diatom-dominated ecosystem to one based on nanoplankton, similar to what occurs from summer to winter off Vancouver Island at present (Wu et al., 1999a and 1999b). Such a switch, for which micropalaeontological evidence is required, would have resulted in a more negative 813Cmar value. The estimates of terrigenous input for Holocene sediments off Vancouver Island (< 20 %) are lower than expected for a continental margin setting. The organic matter in shelf sediments off Washington State, for example, is more than 50 % terrestrial organic matter while slope sediments contain 20 to 30 % terrigenous material (Prahl et al., 1994). The source of this allochthonous detritus in this region is the Columbia River (Prahl et al., 1994). In contrast, the low terrigenous fraction in Holocene sediments off Vancouver Island can be explained by the high trapping efficiency of river-borne detritus within the fjords that are ubiquitous along the British Columbia coast. Such trapping in conjunction with high wave and current energies also explains the very low sedimentation rate characteristic of the present-day shelf and upper slope west of Vancouver Island (Bornhold and Yorath, 1984). Holocene sedimentation rates are generally <2 cm/kyr, except in local depressions on the shelf where sediment focusing results in much higher values (e.g., 40 cm/kyr; Chapter 4). 38 Concentration profiles for the terrestrial and marine organic carbon fractions, calculated using the 5BCorg data, are shown in Figure 2.7. The validity of the marine carbon estimates is supported by a strong positive correlation with the measured concentrations of alkenones (Fig. 2.8a). A positive, though weaker, correlation is also observed between the terrigenous organic carbon estimates and the concentration of the terrigenous biomarker C29 n-alkane (Fig. 2.8b). To try to circumvent the problems associated with varying degrees of dilution, and to get a better understanding of the flux of carbon to the sediment, mass accumulation rates (MARs) have also been calculated (Fig. 2.7). During the late glacial (16.0 to 14.3 kyr B.P.) the MAR of terrestrial organic matter was relatively high, while that for marine organic matter was low. The accumulation of both fractions increased at -14.3 kyr B.P., coincident with a substantial increase in sedimentation rate (i.e., from - 47 up to 169 cm/kyr). During the initial stage of the deglacial between 14.3 and 13.5 kyr B.P. (i.e., the Boiling) terrestrial organic matter dominated the bulk organic fraction. This situation reversed at 13.5 kyr and marine organic matter accumulation was dominant between 13.5 and 12.7 kyr B.P. (i.e., the Allerod). The MAR of both fractions decreased at the end of the Allerod and remained low throughout the Younger Dryas. Marine organic carbon accumulation subsequently increased slightly between 11.2 and 10.0 kyr B.P., while terrestrial organic carbon accumulation continued to decrease. By -15.5 kyr B.P. deglaciation of the shelf off Vancouver Island had commenced (i.e., 15 to 14 14C kyr; Barrie and Conway, 1999; Clague and James, 2002). At this time the linear sedimentation rate at Site JT96-09 was quite high (e.g., 169 cm/kyr during the Boiling), in response to the influx of glacial rock flour generated by the rapid retreat of glaciers. At the beginning of the Allerod, deglaciation of the shelf was essentially complete and the sedimentation rate began to decrease. The relatively low sedimentation rates which characterize Site JT96-09 at present (-5 cm/kyr) were established in the Early Holocene as the shelf was flooded. 39 2.7a) Terrigenous Organic Carbon (wt.%) 0 0.1 0.2 0.3 0.4 I • • • i I 0 0.2 0.4 0.6 0.8 1 MAR of Terrestrial Organic Carbon (g/cm2/kyr) 2.7b) Marine Organic Carbon (wt.%) 0 1 2 3 4 5 0 I • • •  ' -J ' • • MAR | |wt. %| I 1 1 1 i 1 1 1 i I 0 0.2 0.4 0.6 0.8 1 MAR of Marine Organic Carbon (g/cm2/kyr) Fig. 2.7. a) Terrestrial organic matter concentration (open circles) and mass accumulation rate (MAR, solid circles) over the past 16 kyr B.P. b) Marine organic matter concentration (open squares) and mass accumulation rate (MAR, solid squares) over the past 16 kyr B.P. The concentrations of terrestrial and marine organic matter were obtained by first estimating the terrigenous fraction using the 513Corg data (see results in Fig. 2.6) and correcting the total organic carbon data accordingly. The estimated concentrations of terrestrial and marine organic matter have an estimated error of -20 %. 40 2.8a) I? c a 8.1 §13 get s e 5 1200 1000 800 600 400 200 4 O (O^P o o o y= 50.119+218.23x R2 = 0.79 —I—i—i—i—i—|—i—i—i—i—j—i—i—i—i—j—i—i—i—i—p i i i i 0.5 1.5 2.5 3 2.8b) Marine Organic Carbon (wt.%) 4 I 500-400 J 3004 200 J 1004 o-h y= 122.99+393.08x R2 = 0.12 oo 1 I 1 0.05 o.i oo 0 o 8 o° Ooooo0oo°o OO o ° O on O 8o= o 8° 0.15 i I i i i i I i i i i I i i i i I i i i i 0.2 0.25 0.3 0.35 0.4 Terrigenous Organic Carbon (wt.%) Fig. 2.8. The measured concentration of alkenones, a marine biomarker, versus the estimated marine organic carbon concentration, b) The measured concentration of C29 n-alkanes, a terrestrial biomarker, versus the estimated terrestrial organic carbon concentration. 41 There is no evidence that the downslope slumping and/or transport of material was responsible for the high deposition rates throughout the Boiling-Allerod. Such phenomena tend to increase the grain-size and heavy-mineral content of the material laterally transported and deposited on the mid- and upper continental margins (Ganeshram et al., 1999). Coarser-grained sediments are associated with higher Zr/Al ratios as a reflection of the increased zircon content. Thus, had downslope transport been responsible for the very high sedimentation rates during the deglacial, the deposits would be characterized by relatively high Zr/Al ratios. Instead, Zr/Al ratios in the deglacial interval are relatively low (Fig. 2.9), suggesting that these sediments are generally finer-grained than underlying glacial and overlying Holocene strata. The only exception to this general observation is the thin sandy turbidite deposited at the start of the Allerod (Fig. 2.9). The dramatic changes in sedimentation rates over the past 16 kyr make it difficult to interpret changes in palaeoproductivity directly from MARs. For example, the increase in marine organic carbon MAR in the Boiling could simply be due to the large jump in sedimentation rate, given that there is no corresponding increase in the concentration of marine organic carbon. Despite such constraints, it is not possible to rule-out an increase in palaeoproductivity during the Boiling. During the Allerod both the MAR and concentration of marine organic carbon increase. There is no evidence that this is due to enhanced organic matter preservation due to the higher sedimentation rate and/or lower bottom water oxygen concentration at the time (Chapter 3). We therefore infer that export production (i.e., palaeoproductivity) was higher during the Allerod. Increased export production during the deglacial (i.e., the Boiling/Allerod) has been documented for other regions of the CSS to the south of Vancouver Island (Lyle et al, 1992; Gardner et al., 1997; Dean and Gardner, 1998; Mix et al., 1999), and has been attributed to the initiation of upwelling as atmospheric circulation returned to its interglacial state. Sabin and Pisias (1992) used a radiolarian-based transfer function to infer that upwelling began at -15 kyr B.P. and suggested that by 13.0 kyr B.P. present-day upwelling conditions had been 42 Zr/Al Fig. 2.9. The downcore Zr/Al profile for Core JT96-09. The precision (as RSD) of these ratios is ~16 % (i.e., ±0.0001 to 0.0002). 43 attained. Evidence of a return to seasonal upwelling off Vancouver Island during the deglacial and stronger upwelling than at present during the Allerod is provided by radiocarbon dating. During the Allerod the reservoir age of surface waters, recorded in an uplifted deposit found in the Fraser River lowlands, increased from 800 to 1100 years (Kovanen and Easterbrook, 2002) suggesting enhanced upwelling of old, 14C-depleted water from depth. Furthermore, the age difference between contemporaneous benthic and planktonic foraminifera decreased during the Allerod (Chapter 4). The simplest explanation for this decrease is the upwelling of old, nutrient-rich water that would have stimulated primary production. The relatively low marine organic carbon accumulation rates observed for the Holocene are unexpected given that the west coast of Vancouver Island is characterized by 2 2 high primary productivity at present (-250 gC/m /yr on the lower slope to 400 gC/m /yr on the shelf; Antoine etai., 1996). In large part, the decrease in mass accumulation rates reflects the substantial decrease in sedimentation rate, but preservational influences may also have played a role. Numerous studies have shown that organic carbon is better preserved when sedimentation rate is high, possibly because exposure time to various oxidants is limited (Hedges and Keil, 1995; Gelinas et al., 2001). It is also possible that the large supply of fine grained aluminosilicate particles, associated with a higher sedimentation rate, offers more mineral surface area on which organic compounds can be adsorbed (Keil et al., 1994; Mayer, 1994). Once adsorbed this organic material maybe protected from degradation. In contrast, when the sedimentation rate is low there is continued exposure to oxidants via bioturbation and diffusive influx, and less mineral surface area on which organics may be adsorbed. In summary, the very low sedimentation rate (i.e., poor preservation) and apparently lower primary productivity relative to the deglacial lead to a lower rate of organic matter accumulation throughout the Holocene. 2.4.2 Biogenic barium and opal accumulation AA Barium is often enriched in sediments beneath highly productive surface waters (Goldberg and Arrhenius, 1958; Dehairs et al., 1980; Dehairs et al., 1992), presumably because authigenic barite formation is linked to the supply of organic particles. The exact mechanism responsible for the precipitation of barite is still not fully understood, but it appears to be related to the formation of the mineral within decaying organic particles, such as fecal pellets (Dehairs et al., 1980; Bishop, 1988; Bernstein et al. 1992; Ganeshram et al., in press). Based on the observed spatial correlation with export production, the barium content of sediments is commonly used as a palaeoproductivity indicator (e.g., Schmitz, 1987; Dymond et al., 1992; Gingele and Dahmke, 1994; Shimmield et al., 1994; Francois et al., 1995; Ganeshram, 1996; Nurnberg et al., 1997; Thompson and Schmitz, 1997). In Core JT96-09, both the concentration and MAR of bio-barium are high throughout the Allerod. These increases are contemporaneous with the increase in marine organic carbon accumulation, which supports the interpretation that export production was high during this interval. However, the mass accumulation rate of bio-barium is also relatively high during the Boiling. Most if not all, of this apparent increase is a reflection of the dramatically higher sedimentation rate at this time (169 cm/kyr) and the influence this has on the calculation of mass accumulation rates. Likewise, the higher MARs during the late glacial, relative to the Holocene, can be explained by differences in sedimentation rates (47 cm/kyr versus 5 cm/kyr for the late glacial and Holocene, respectively) and do not imply that export production was higher during the glacial. Only when both concentrations and accumulation rates increase contemporaneously, such as during the Allerod, is it possible to infer higher export production. The primary source of biogenic silica (i.e., opal) in these sediments is diatoms. As with organic carbon and bio-barite, opal preservation efficiency increases as sediment accumulation rate increases (Ragueneau et al., 2000) because rapid sedimentation more 45 quickly isolates porewaters from diffusive contact with Si-undersaturated bottom waters. Given this strong preservational artifact, there is some question as to whether opal content is a reliable indicator of surface productivity (Archer et al., 1993). Nevertheless, various studies have shown that if used cautiously the opal content of sediments can reflect export production history (Lyle et al., 1992; Francois et al., 1997; Ganeshram et al., 2000). In Core JT96-09 the slowly accumulating Holocene sediments (5 cm/kyr) have a higher opal content than the more rapidly deposited glacial and deglacial sediments (48 to 169 cm/kyr). This trend is opposite to what is expected if sedimentation rate was the only control on opal concentration. Export of opal from surface waters to the sediment must also have influenced sedimentary concentrations. As already discussed, MAR data for Core JT96-09 may be somewhat misleading, given the large variability in sedimentation rates. However, in the Allerod both the concentration and MAR of opal increase, coincident with the increase in organic carbon and bio-barium concentrations and MARs. Thus, opal data further support the interpretation that marine productivity and subsequent export flux to the sediment was probably relatively high during the Allerod. 2.4.3 Alkenones C37 alkenones are produced by certain prymnesiophyte algae (family Gephyrocapsaceae) and thus can be used as a biomarker for these phytoplankton (Marlowe et al., 1990). The coccolithophorid Emiliania huxleyi is the most important alkenone producer in the northeast Pacific (Okada and Honjo, 1973; Prahl and Wakeham, 1987; Marlowe et al., 1990), although a minor contribution by other prymesiophytes cannot be discounted. Relative alkenone abundance (i.e., C37/OC) in Core JT96-09 reaches a maximum between 13.5 and 11.9 kyr B.P., at approximately the same time that other palaeoproductivity proxies peak (Fig. 2.2). However, the relative abundance remains high when the accumulation rates of the other proxies decrease at -12.5 kyr B.P. (i.e., the start of Younger Dryas). In part, this 46 can be attributed to differences in the factors that control the productivity of coccolithophorids compared to those that control diatoms. When nutrient concentrations are high, such as during upwelling events, diatoms are favoured whereas coccolithophorids can thrive when waters become nutrient-depleted (Werne et al., 2000 and references therein). The low accumulation rates of marine organic carbon, bio-barium and opal during the Younger Dryas suggest that nutrient supply and thus diatom-based productivity was diminished at this time. However, given the ecological contrasts between diatoms and coccolithophorids this would not have had the same impact on alkenone production. 2.4.4 Nitrogen-isotope record During nitrate uptake, marine phytoplankton preferentially incorporate 14N into organic matter, but as nitrate becomes scarce there is less discrimination against 15N (Montoya et al., 1994 and references therein). Thus, the nitrogen isotopic composition of organic matter is commonly used as a proxy for nitrate utilization, with relatively high 815N values reflecting high nitrate utilization relative to the supply (Altabet et al., 1991; Altabet and Francois, 1994a and b; Farrell et al., 1995). However, other factors such as nitrogen fixation, diagenesis, water column denitrification, the presence of inorganic nitrogen, changes in trophic cycling, and the presence of terrestrial organic matter also may influence the 815N composition of sediments (Montoya, 1994). Nitrogen-fixation involves little isotopic fractionation and thus produces organic matter with low 815N values (~ 0%o; Montoya et al., 1994; Carpenter et al., 1997). However, this process is currently believed to be important only under oligotropic conditions and in tropical to subtropical waters where the cyanobacterium Trichodesmium occurs (Carpenter et al., 1997; Capone et al., 1997). This process can therefore be ruled-out as the cause of low 815N values in late glacial and deglacial sediments. 47 The influence that diagenesis has on the 515N composition of organic settling through the water column is dependant on water depth. In the deep ocean a diagenetic offset of as much as +4 to +7 %o may occur between sinking particles collected from deep waters and surface sediments (Saino and Hattori, 1987; Altabet and Francois, 1994b, Altabet, 1996). But when the water depth is relatively shallow and settling is rapid there is little change in the S15N (Altabet et al., 1999; Pride et al., 1999). Such is likely the case for Core JT96-09 that was collected from 920 m water depth. The influence that sedimentary diagenesis has on the 815N composition of organic matter is more ambiguous. Some studies have documented a decrease in 815N of a few permil due the degradation (i.e., loss) of 15N-enriched organic compounds (De Lange et al., 1994; Prahl et al., 1997). In contrast, other studies observe an increase in 815N (Sigman et al., 1999; Freudenthal et al., 2001) or no change at all (Ganeshram et al., 2000). In near-surface sediments collected from the continental margin off Vancouver Island there is no change in bulk sediment 815N values to a depth of 20 cm below the sediment-water interface (Chapter 4). These data suggest that, despite very slow Holocene sedimentation rates, degradation of organic matter in this environment has little effect on the 8I5N composition of organic matter. Low sedimentary 815N values can be the result of adsorption of inorganic nitrogen onto illitic clay minerals (Freudenthal et al., 2001). The presence of inorganic (or excess) nitrogen would be recorded by a positive intercept on the nitrogen-axis of a C-N plot. No such positive N-intercept is observed for samples from Core JT96-09 when total organic carbon and total nitrogen are plotted, but when the concentration of marine organic carbon is plotted versus nitrogen a N-intercept of 0.0025 wt. % is obtained. This suggests that some inorganic N is present, but the concentration is so low as to have a negligible effect on 815N values. We have shown above that the input of terrigenous organic matter is largely 13 responsible for the low 8 Corg values of late glacial and early deglacial sediments off Vancouver Island, and it is reasonable to assume that a similar effect should be seen in the 48 S15N data. Terrestrial organic matter is typified by low 515N values, reflecting the uptake and fixation of atmospheric N2 by land plants. Indeed, there is a clear decrease in 515N values during the Boiling when terrestrial input was highest and an increase in 815N values during the Allerod when terrestrial input decreased (Fig. 2.2c). If terrestrial organic matter input was the only factor controlling the N-isotopic composition it should be possible to estimate the terrigenous proportion using 815N data in the same way that 813Corg data were used, and the results should be similar. This calculation requires reasonable estimates for the 815N composition of terrigenous and marine organic matter (815Nterr and 815Nraar, respectively). The S15Nterr can be calculated using the regression equation of late Pleistocene carbon- and nitrogen-isotopic data (Equation 2.3; see Fig. 2.4). 815Nterr = 40.25 + 1.45(813CterT) (r = 0.94) (2.3) Assuming a terrestrial endmember 813Cterr value of -27%o (Wu, 1997) yields a 815NterT value of+l.l%o. A slightly higher 815Nterr value of +2.6%o is obtained if a 813Cterr value of-26%o is instead used. These 815Nterr estimates are similar to measured values for organic matter from various California rivers and lakes (0.3 to 6.9 %o, avg. 4 %o; Peters et al., 1978). Estimating the nitrogen isotopic composition of the marine end-member (815Nmar) is more complicated because 815Nmar values can vary spatially and seasonally within an upwelling system. Surface sediments from the continental shelf off Vancouver Island have relatively low 815N values (+4.8%o) while slope sediments are characterized by higher 815N values (+6.3 to +6.8%o; Chapter 4). This difference is the result of upwelling of water onto the shelf followed by the preferential uptake of 14N-enriched nitrate by phytoplankton that leaves the residual nitrate 15N-enriched. As waters are advected away from the site of upwelling (i.e., offshore) they become progressively more 15N-enriched due to continued biological activity. The organic matter produced from this nitrate and which is eventually deposited on the slope is thus 15N-enriched relative to that deposited on the shelf (Chapter 4). Data from sediment 49 traps moored on the upper slope off Vancouver Island indicate that the 815N of settling particulate organic matter (SPOM) is relatively heavy, ranging from +7.7 %o during bloom events in summer and fall up to +9.0 %o in winter, with an average annual flux-weighted value of +8.1 %o (Wu, et al., 1999a). However, surface sediments from the slope have lower values (+6.3 to +6.8 %o) than this SPOM possibly due to the input of minor terrestrial organic matter and/or the downslope transport of 14N-enriched organic matter from the shelf. Our 15 13 best estimate of 8 Nmar is obtained using Equation 2.3 and a marine 8 Corg value of -21 %o. This calculation yields a 815Nmar value of +9.8 %o which is significantly higher than the 815N composition of trap and surface sediments. The regression derived from a plot of C29/OC and 8I5N data yields a similar value (+9.5 %o; Fig. 2.5b). The 815Nmar values obtained from the linear regression of the data are more similar to the 815N values of settling particulate organic matter in winter (~ 9 %o) rather than the annual flux-weighted value of Wu et al., (1999a). Thus, the present-day winter ecosystem (i.e., nanoplankton-dominated community) may be a better model for the late glacial ecosystem. Although speculative, this is consistent with the inference that upwelling was greatly diminished or non-existent off Vancouver Island during the last glacial and thus the nitrate supply was reduced. Under such conditions a nanoplankton-dominated ecosystem, which is based on the use of regenerated forms of nitrogen (i.e., urea), would have been favoured. However, it is also possible that the isotopic composition of the source nitrate (i.e., the nitrate that is supplied to the euphotic zone) was different. This possibility is discussed below. Substituting the 815N values of terrestrial and marine organic matter (+1 and +9 %o, respectively) for their 8 C counterparts in Equation 2.2 yields terrigenous values of 40 to 70 % for the late glacial and early deglacial, <10 to 40 % for the late deglacial, and 8 to 31 % for 13 the Holocene (Fig. 2.6). These estimates compare moderately well with 8 C-based estimates for the late glacial, deglacial and early Holocene, but are up to 25 % higher for the mid- to late Holocene. (Fig. 2.6) Deriving mid to late Holocene values that are comparable to estimates based on 813C values requires using either a lower S15Nterr value (< +1 %o) or a 50 lower 815Nmar value (+7 %o). It is unlikely that 515Nten- values would have varied or been less than 1 %o. It is more likely that 815Nmar changed over time. Variations in S^Nmar could have resulted from a change in the ecosystem (e.g., diatom-dominated, shorter foodwebs result in lower 515Nmar values; Montoya et al., 1994; Wu et al., 1999a and b). It is also possible that the isotopic composition of the source nitrate changed. At present, the 515N of source nitrate off Vancouver Island ranges from approximately +5 to +6%o (Wu et al., 1997; Kienast et al., 2002). Waters with nitrate 815N values substantially higher than +5 %o (i.e., up to +19 %o) are found in regions characterized by water column denitrification (Cline and Kaplan, 1975; Brandes et al., 1998; Altabet et al., 1999). While denitrification is not occurring off Vancouver Island at present it may have occurred in the past when the oxygen minimum zone (OMZ) was more intense (i.e., more oxygen depleted). Episodes of OMZ intensification during the Allered and at the Pleistocene-Holocene boundary are characterized by relatively high sedimentary 8I5N values (Chapter 3). While high sedimentary 815N values suggest that denitrification influenced the N-isotopic composition of organic matter off Vancouver Island, they do not provide conclusive evidence that denitrification was occurring in this specific region. It is also possible that 15N-enriched nitrate was imported from the Equatorial East Pacific by the California Undercurrent (Liu and Kaplan, 1989; Altabet et al., 1999; Kienast et al., 2002). If this is the case, then the 815N composition of nitrate off Vancouver Island would have been influenced by secular changes in denitrification in the Equatorial East Pacific (Ganeshram et al., 1995; Pride et al., 1999) and by the strength of California Undercurrent (Kienast et al., 2002). Ongoing work based on new cores collected from the region will hopefully resolve this question. 2.5 Summary At the end of the last glacial (-14.3 kyr B.P.) the nature of organic matter accumulation on the continental margin off Western Canada changed dramatically. Prior to 51 14.3 kyr B.P. the accumulation of terrestrial organic matter was high while the accumulation of marine organic matter was relatively low. This was a reflection of low primary production, and resulting low export production, because glacial-mode atmospheric circulation did not drive upwelling in the region. At the start of the deglacial (i.e., the Boiling) the accumulation of marine organic matter increased, suggesting a return to modern atmospheric and oceanic circulation. However, most if not all of this apparent increase is due to the very high sedimentation rate (169 cm/kyr), although higher marine export production cannot be totally discounted. Marine organic matter deposited during the early deglacial was diluted by abundant terrestrial organic matter, generated as glaciers retreated from the continental shelf. This situation continued until the start of the Allerod when the input of terrestrial organic matter decreased and the accumulation of marine organic matter doubled. The combination of higher marine organic carbon concentrations and mass accumulation rates suggests that marine productivity did increase at this time. There was a brief return to glacial-like conditions during the Younger Dryas and a corresponding decrease in marine organic matter accumulation. Palaeo-proxy data suggest the accumulation of marine organic matter during the Holocene was, and still is, lower than during the late Pleistocene. This may be a reflection of low organic matter preservation due to the relatively low sedimentation rates (i.e., relatively long oxygen exposure times; see Chapter 3) and intense biological utilization. 52 2.6 References Altabet, M.A., 1996. Nitrogen and carbon isotopic tracers of the source and transformation of particles in the deep-sea. In: Ittekkot, V., Schafer, P., Honjo, S., Depetris, P.J. (Eds.), Particle Flux in the Ocean. John Wiley and Sons, London, pp. 155-184. Altabet, M.A., Francois, R., 1994a. Sedimentary nitrogen isotopic ratio as a recorder for surface ocean nitrate utilization. Global Biogeochemical Cycles 8, 103-116. Altabet, M.A., Francois, R., 1994b. The use of nitrogen isotopic ratios for reconstruction of past changes in surface ocean nutrient utilization. In: Zahn, R., Pedersen, T.F., Kaminski, M.A., Labeyrie, L. (Eds.), Carbon Cycling in the Glacial Ocean: Constraints on the Ocean's Role in Global Change. NATO ASI Series, Vol. 117. Springer-Verlag, Berlin, pp. 281-306. Altabet, M.A., Deuser, W.G., Honjo, S., Stienen, C, 1991. Seasonal and depth-related changes in the source of sinking particles in the North Atlantic. Nature 354, 136-138. Altabet, M.A., Pilskaln, C, Thunell, R., Pride, C, Sigman, D., Chavez, F., Francois, R., 1999. The nitrogen isotope biogeochemistry of sinking particles from the margin of the Eastern North Pacific. Deep Sea Research (I) 46, 655-679. Antoine, D., Andre, J.M., Morel, A., 1996. Oceanic primary production. 2. Estimation at global scale from satellite (coastal zone colour scanner) chlorophyll. Global Biogeochemical Cycles 10, 57-69. Archer, D., Lyle, M., Rodgers, K., Froelich, P., 1993. What controls opal preservation in tropical deep-sea sediment? Paleoceanography 8, 7-21. Bakun, A., 1990. Global climate change and intensification of coastal ocean upwelling. Science 247, 198-201. Bard, E., Hamelin, B., Arnold, M., Montaggioni, L., Cabioch, G., Faure, G., Rougerie, F., 1996. Deglacial sea-level record from Tahiti corals and the timing of global meltwater discharge. Nature 382, 241-244. 53 Barry, J.V., Conway, K.W., 1999. Late Quaternary glaciation and postglacial stratigraphy of the Northern Pacific Margin of Canada. Quaternary Research 51, 113-123. Bernstein, R.E., Byrne, R.H., Betzer, P.R., Greco, A.M., 1992. Morphologies and transformations of celestite in seawater: The role of acantharians in strontium and barium geochemistry. Geochimica et Cosmochimica Acta 56, 3273-3279. Bishop, J.K.B., 1988. The barite-opal-organic carbon association in oceanic particulate matter. Nature 332, 341-343. Bornhold, B.D., Yorath, C.J., 1984. Surficial geology of the continental shelf, northwestern Vancouver Island. Marine Geology 57, 89-112. Brandes, J.A., Devol, A.H., Yoshinari, T., Jayakumar, D.A., Naqvi S.W.A., 1998. Isotopic composition of nitrate in the central Arabian Sea and eastern tropical North Pacific: A tracer for mixing and nitrogen cycles. Limnology and Oceanography 43, 1680-1689. Calvert, S.E., 1990. Geochemistry and origin of the Holocene sapropel in the Black Sea. In: Ittekkot, V., Kempe, S., Michaelis, W., Spitzy, A. (Eds.), Facets of Modern Biogeochemistry. Springer-Verlag, Berlin, pp. 326-352. Capone, D.G., Zehr, J.P., Paerl, H.W., Bergman, B., Carpenter, E.J., 1997. Trichodesmium, a globally significant marine cyanobacterium. Science 276, 1221-1229. Carpenter, E.J., Harvey, H.R., Fry, B., Capone, D.G., 1997. Biogeochemical tracers of the marine cyanobacterium Trichodesmium. Deep Sea Research (I) 44, 27-38. Clague, J.J., James, T.S., 2002. History and isostatic effects of the last ice sheet in southern British Columbia. Quaternary Science Reviews 21,71-87, Cline, J.D., Kaplan, I.R., 1975. Isotopic fractionation of dissolved nitrate during denitrification the Eastern Tropical North Pacific Ocean. Marine Chemistry 3, 271-299. COHMAP Members, 1988. Climatic changes of the last 18,000 years: Observations and model simulations. Science 241, 1043-1052. De Lange, G.J., Os, B.V., Pruysers, P.A., Middelburg, J.J., Castradori, D., Van Santvoort, P., Muller, P.J., Eggenkamp, H., Prahl, F.G., 1994. Possible early diagenetic alteration of palaeo proxies. In: Zahn, R., Pedersen, T.F., Kaminski, M.A., Labeyrie, L. (Eds.), Carbon Cycling 54 in the Glacial Ocean: Constraints on the Ocean's Role in Global Change. NATO ASI Series, Vol. 117. Springer-Verlag, Berlin, pp. 225-257. Dean, W.E., Gardner, J.V., 1998. Pleistocene to Holocene contrasts in organic matter production and preservation on the California continental margin. Geological Society of America Bulletin 110, 888-899. Dehairs, F., Chesselet, R., Jedwab, J., 1980. Discrete suspended particles of barite and the barium cycle in the open ocean. Earth and Planetary Science Letters 49, 528-550. Dehairs, F., W. Baeyens, W., Goeyens, L., 1992. Accumulation of suspended barite at mesopelagic depths and export production in the Southern Ocean. Science 258, 1332-1335. Doose, H., Prahl, F.G., Lyle, M.W., 1997. Biomarker temperature estimates for modern and last glacial surface waters of the California Current system between 33° and 42°N. Paleoceanography 12, 615-622. Dymond, J., Suess, E., Lyle, M., 1992. Barium in deep-sea sediment: A geochemical proxy for paleoproductivity. Paleoceanography 7,163-181. Eglinton, G., Hamilton, R.H., 1967. Leaf epicuticular waxes. Science 156, 1322-1335. Falkner, K.K., Klinkhammer, G.P., Bowers, T.S., Todd, J.F., Lewis, B.L., Landing, W.M., Edmond, J.M., 1993. The behavior of barium in anoxic marine waters. Geochimica et Cosmochimica Acta 57, 537-554. Farrell, J.W., Pedersen, T.F., Calvert, S.E., Nielsen, B., 1995. Glacial-interglacial changes in nutrient utilization in the equatorial Pacific Ocean. Nature 377, 514-517. Faure, G., 1986. Principles of Isotope Geology, 2nd edition. John Wiley & Sons Inc., New York, 589pp. Fernandes, M.B., Sicre, M.-A., 2000. The importance of terrestrial organic carbon inputs on Kara Sea shelves as revealed by n-alkanes, OC and 513C values. Organic Geochemistry 31, 363-374. 13 Francois, R., Altabet, M.A., Goericke, R., 1994. Changes in the 8 C of surface water particulate organic matter across the subtropical convergence in the S.W. Indian Ocean. In: 55 Zahn, R., Pedersen, T.F., Kaminski, M.A., Labeyrie, L. (Eds.), Carbon Cycling in the Glacial Ocean: Constraints on the Ocean's Role in Global Change. NATO ASI Series, Vol. 117. Springer-Verlag, Berlin, pp. 307-321. Francois, R., Altabet, M.A., Yu, E.-F., Sigman, D.M., Bacon, M.P., Frank, M., Bohrmann, G., Bareille, G., Labeyrie, L.D., 1997. Contribution of Southern Ocean surface-water stratification to low atmospheric C02 concentrations during the last glacial period. Nature 389, 929-935. Francois, R., Honjo, S., Manganini, S J., Ravizza, G.E., 1995. Biogenic barium fluxes to the deep sea: Implications for paleoproductivity reconstruction. Global Biogeochemical Cycles 9, 289-303. Freudenthal, T., Wagner, T., Wenzhofer, F., Zabel, M., Wefer, G., 2001. Early diagenesis of organic matter from sediments of the eastern subtropical Atlantic: Evidence from stable nitrogen and carbon isotopes. Geochimica et Cosmochimica Acta 65, 1795-1808. Fry, B., Wainright, S.C., 1991. Diatom sources of C-rich carbon in marine food webs. Marine Ecology Progress Series 76,149-157. Ganeshram, R.S., 1996. On the glacial-interglacial variability of upwelling, carbon burial and denitrification on the northwestern Mexican continental margin. Ph.D. dissertation thesis, University of British Columbia. Ganeshram, R.S., Pedersen, T.F., Calvert, S.E., Murray, J.W., 1995. Large changes in oceanic nutrient inventories from glacial to interglacial periods. Nature 376, 755-758. Ganeshram, R.S., Calvert, S.E., Pedersen, T.F., Cowie, G.A., 1999. Factors controlling the burial of organic carbon in laminated and bioturbated sediments off NW Mexico. Implications for hydrocarbon preservation. Geochimica et Cosmochimica Acta 63, 1723-1734. Ganeshram, R.S., Pedersen, T.F., Calvert, S.E., McNeill, G.W., Fontugne, M.R., 2000. Glacial-interglacial variability in denitrification in the world's ocean: Causes and consequences. Paleoceanography 15, 361-376. 56 Ganeshram, R.S., Francois, R., Commeau, J., Brown-Leger, S.L., 2003. An experimental investigation of barite formation in seawater. Geochimica et Cosmochimica Acta 67, 2599-2605. Gardner, J.V., Dean, W.E., Dartnell, P., 1997. Biogenic sedimentation beneath the California Current system for the past 30 kyr and its paleoceanographic significance. Paleoceanography 12, 207-225. Gelinas, Y., Baldock, J.A., Hedges, J.I., 2001. Organic carbon composition of marine sediments: Effect of oxygen exposure on oil generation potential. Science 294, 145-148. Gingele, F., Dahmke, A., 1994. Discrete barite particles and barium as tracers of paleoproductivity in South Atlantic sediments. Paleoceanography 9,151-168. Goericke, R., Fry, B., 1994. Variations of marine plankton 813C with latitude, temperature, and dissolved C02 in the world ocean. Global Biogeochemical Cycles 8, 85-90. Goldberg, E.D., Arrhenius, G., 1958. Chemistry of Pacific pelagic sediments. Geochimica et Cosmochimica Actal3, 153-212. Hedges, J.I., Keil, R.G., 1995. Sedimentary organic matter preservation: An assessment and speculative synthesis. Marine Chemistry 49, 81-115. Hedges, J.I., Oades, J.M., 1997. Comparative organic geochemistry of soils and marine sediments. Organic Geochemistry 27, 319-361. Hickey, B.M., 1998. Coastal oceanography of western North America from the tip of Baja California to Vancouver Island. In: Robinson, A.R., Brink, K.H. (Eds.), The sea. John Wiley & Sons Inc., pp. 345-391. Jasper, F.P., Gagosian, R.B., 1993. The relationship between sedimentary organic carbon isotopic composition and organic biomarker compound concentration. Geochimica et Cosmochimica Acta 57, 167-186. Jasper, J.P., Hayes, J.M., 1990. A carbon isotope record of C02 levels during the late Quaternary. Nature 347, 462-464. 57 Keil, R.G., Tsamakis, E., Fuh, C.B., Giddings, J.C., Hedges, J.I., 1994. Mineralogical and textural controls on the organic composition of coastal marine sediments: Hydrodynamic separation using SPLITT-fractionation. Geochimica et Cosmochimica Acta 58, 879-893. Kienast, S.S, McKay, J.L., 2001. Sea surface temperatures in the subarctic Northeast Pacific reflect millennial-scale climate oscillations during the last 16 kyrs. Geophysical Research Letters 28, 1563-1566. Kienast, S.S., Calvert, S.E., Pedersen, T.F., 2002. Nitrogen isotope and productivity variations along the northeast Pacific margin over the last 120 kyr: Surface and subsurface palaeoceanography. Paleoceanography 17, doi: 10.1029/2001PA000650. Kovanen, D.J., Easterbrook, D.J., 2002. Paleodeviations of radiocarbon marine reservoir values for the northeast Pacific. Geology 30, 243-246. Law, E.A., Popp, B.N., Bidigare, R.R., Kennicutt, M.C., Macko, S.A., 1995. Dependence of phytoplankton carbon isotopic composition on growth rate and [C02]aq.' Theoretical considerations and experimental results. Geochimica et Cosmochimica Acta 59, 1131-1138. Liu, K.-K., Kaplan, I.R., 1989. The eastern tropical Pacific as a source of 15N-enriched nitrate in seawater off southern California. Limnology and Oceanography 35, 820-830. Lyle, M., Zahn, R., Prahl, F., Dymond, J., Colier, R., Pisias, N., Suess, E., 1992. Paleoproductivity and carbon burial across the California Current: The mulitracers transect, 42°N. Paleoceanography 7, 251-272. Mayer, L.M., 1994. Surface area control of organic carbon accumulation in continental shelf sediments. Geochimica et Cosmochimica Acta 58,1271-1284. Marlowe, I.T., Brassell, S.C., Eglinton, G., Green, J.C., 1990. Long-chain alkenones and alkyl alkenoates and the fossil coccolith record of marine sediments. Chemical Geology 88, 349-375. Mathewes, R.W., Heusser, L.E., Patterson, R.T., 1993. Evidence for a Younger Dryas-like cooling event on the British Columbia coast. Geology 21, 101-104. 58 McManus, J., Berelson, W.M., Klinkhamer, G.P., Kilgore, T.E., Hammon, D.E., 1994. Remobilization of barium in continental margin sediments. Geochimica et Cosmochimica Acta 58, 4899-4907. McManus, J., Berelson, W.M., Klinkhammer, G.P., Johnson, K.S., Coale, K.H., Anderson, R.F., Kumar, N., Burdige, D.J., Hammond, D.E., Brumsack, H.J., McCorkle, D.C., Rushdi, A., 1998. Geochemistry of barium in marine sediments: Implications for its use as a paleoproxy. Geochimica et Cosmochimica Acta 62, 3453-3473. Meyers, P.A., 1997. Organic geochemical proxies of paleoceanographic, paleolimnologic, and paleoclimatic processes. Organic Geochemistry 27, 213-250. Minoura, K., Hoshino, K., Nakamura, T., Wada, E., 1997. Late Pleistocene-Holocene paleoproductivity circulation in the Japan Sea: Sea-level control on 813C and 815N records of sediment organic material. Palaeogeography, Palaeoclimatology, Palaeoecology 135, 41-50. Mix, A.C., Lund, D.C., Pisias, N.G., Boden, P., Bornmalm, L., Lyle, M., Pike, J., 1999. Rapid climate oscillations in the northeast Pacific during the last deglaciation reflect northern and southern hemisphere sources. In: Clark, P.U., Webb, R.S., Keigwin, L.D. (Eds.), Mechanisms of Global Climate Change at Millennian Time Scales, Geophysical Monograph 112. American Geophysical Union, Washington, D.C., pp. 127-148. Montoya, J.P., 1994. Nitrogen isotope fractionation in the modern ocean: Implications for the sedimentary record. In: Zahn, R., Pedersen, T.F., Kaminski, M.A., Labeyrie, L. (Eds.), Carbon Cycling in the Glacial Ocean: Constraints on the Ocean's Role in Global Change. NATO ASI Series, Vol. 117. Springer-Verlag, Berlin, pp. 259-279. Mortlock, R.A., Froelich, P.N., 1989. A simple method for the rapid determination of biogenic opal in pelagic marine sediments. Deep Sea Research 36, 1415-1426. Mortyn, P.G., Thunell, R.C., Anderson, D.M., Stott, L.D., Le, J., 1996. Sea surface temperature changes in the Southern California Borderlands during the last glacial-interglacial cycle. Paleoceanography 11, 415-430. Nurnberg, CC, Bohrmann, G., Schluter, M., 1997. Barium accumulation in the Atlantic sector of the Southern Ocean: Results from 190,000-year records. Paleoceanography 12, 594-603. 59 Ohkouchi, N., Kawamura, K., Taira, A., 1997. Molecular paleoclimatology: Reconstruction of climate variabilities in the late Quaternary. Organic Geochemistry 27, 173-183. Okada, H., Honjo, S., 1973. The distribution of oceanic coccolithophorids in the Pacific. Deep Sea Research 20, 355-374. Ortiz, J., Mix, A., Hostetler, S., Kashgarian, M., 1997. The California Current of the last glacial maximum: Reconstruction at 42°N based on multiple proxies. Paleoceanography 12, 191-205. Pedersen, T.F., 1983. Increased productivity in the eastern equatorial Pacific during the last glacial maximum (19,000 to 14,000 yr B.P.). Geology 11, 16-19. Peters, K.E., Sweeney, R.E., Kaplan, I.R., 1978. Correlation of carbon and nitrogen stable isotope ratios in sedimentary organic matter. Limnology and Oceanography 23, 598-604. Popp, B.N., Takigiku, R., Hayes, J.M., Louda, J.W., Baker, E.W., 1989. The post-Paleozoic 13 • chronology and mechanism of C depletion in primary marine organic matter. American Journal of Science 289, 436-454. Popp, B.N., Laws, E.A., Bidigare, R.R., Dore, J.E., Hanson, K.L., Wakeham, S.G., 1998. Effect of phytoplankton cell geometry on carbon isotopic fractionation. Geochimica et Cosmochimica Acta 62, 69-77. Popp, B.N., Trull, T., Kenig, F., Wakeham, S.G., Rust, T.M., Tilbrook, B., Griffiths, F.B., Wright, S.W., Marchant, H.J., Bidigare, R.R., Laws, E.A., 1999. Controls on the carbon isotopic composition of Southern Ocean phytoplankton. Global Biogeochemical Cycles 11, 827-843. Prahl, F.G., Wakeham, S.G., 1987. Calibration of unsaturation patterns in long-chain ketone compositions for palaeotemperature assessment. Nature 330, 367-369. Prahl, F.G., Muehlhausen, L.A., Lyle, M., 1989. An organic geochemical assessment of oceanographic conditions at MANOP Site C over the past 26,000 years. Paleoceanography 4, 495-510. 60 Prahl, F.G., Ertel, J.R., Goni, M.A., Sparrow, M.A., Eversmeyer, B., 1994. Terrestrial organic carbon contributions to sediments on the Washington margin. Geochimica et Cosmochimica Acta 58, 3035-3048. Prahl, F.G., De Lange, G.J., Scholten, S., Cowie, G.L., 1997. A case of post-depositional aerobic degradation of terrestrial organic matter in turbidite deposits from the Madeira Abyssal Plain. Organic Geochemistry 27,141-152. Pride, C, Thunell, T., Sigman, D., Keigwin, L., Altabet, M., Tappa, E., 1999. Nitrogen isotopic variations in the Gulf of California since the last deglaciation: Response to global climate change. Paleoceanography 14, 397-409. Ragueneau, O., Treguer, P., Leynaert, A., Anderson, R.F., Brzezinski, M.A., DeMaster, D.J., Dugdale, R.C., Dymond, J., Fischer, G., Francois, R., Ffeinze, C, Maier-Reimer, E., Martin-Jezequel, V., Nelson, D.M., Queguiner, B., 2000. A review of the Si cycle in the modern ocean: Recent progress and missing gaps in the application of biogenic opal as a paleoproductivity proxy. Global Planetary Change 26, 317-365. Rau, G.H., 1994. Variations in sedimentary organic 813C as a proxy for past changes in ocean and atmospheric [C02]. In: Zahn, R., Pedersen, T.F., Kaminski, M.A., Labeyrie, L. (Eds.), Carbon Cycling in the Glacial Ocean: Constraints on the Ocean's Role in Global Change. NATO ASI Series, Vol. 117. Springer-Verlag, Berlin, pp. 307-322. 13 Rau, G.H., Takahashi, T., Des Marais, D.J., 1989. Latitudinal variations in plankton 8 C: implications for C02 and productivity in past oceans. Nature, 341, 516-518. Rau, G.H., Froelich, P.N., Takahashi, T., Des Marais, D.J., 1991. Does sedimentary organic 813C record variations in Quaternary ocean [C02(aq)]? Paleoceanography 6, 335-347. Rau, G.H., Riebesell, U., Wolf-Gladrow, D., 1997. C02aq-dependent photosynthesis 13C fractionation in the ocean: A model versus measurements. Global Biogeochemical Cycles 11,267-278. Sabin, A.L., Pisias, N.G., 1996. Sea surface temperature changes in the northeastern Pacific Ocean during the past 20,000 years and their relationship to climate change in northwestern North America. Quaternary Research 46, 48-61. 61 Saino, T., Hattori, A., 1987. Geographical variations of the water column distribution of suspended particulate organic nitrogen and its 15N natural abundance in the Pacific and its marginal seas. Deep Sea Research 34, 807-827. Sancetta, C, Lyle, M., Heusser, L., Zahn, R., Bradbury, J., 1992. Late-glacial to Holocene changes in winds and upwelling, and seasonal production of the northern California Current system. Quaternary Research 38, 359-370. Sarnthein, M., Winn, K., Duplessy, J.-C, Fontugne, M.R., 1988. Global variations of surface ocean productivity in low and mid latitudes: Influence on C02 reservoirs of the deep ocean and atmosphere during the last 21,000 years. Paleoceanography 3, 361-399. Schmitz, B., 1987. Barium, equatorial high productivity, and the northward wandering of the Indian continent. Paleoceanography 2, 63-77. Schubert, C.J., Villanueva, J., Calvert, S.E., Cowie, G.L., von Rad, U., Schulz, H., Berner, U. and Erlenkeuser, H., 1998. Stable phytoplankton community structure in the Arabian Sea over the past 200,000 years. Nature 394, 563-566. Shimmield, G., Derrick, S., Mackensen, A., Grobe, H., Pudsey, C, 1994. The history of barium, biogenic silica and organic carbon accumulation in the Weddell Sea and Antarctic Ocean over the last 150,000 years. In: Zahn, R., Pedersen, T.F., Kaminski, M.A., Labeyrie, L. (Eds.), Carbon Cycling in the Glacial Ocean: Constraints on the Ocean's Role in Global Change. NATO ASI Series, Vol. 117. Springer-Verlag, Berlin, pp. 555-574. Sigman, D.M., Altabet, M.A., Francois, R., McCorkle, D.C., Gaillard, J.-F., 1999. The isotopic composition of diatom-bound nitrogen in Southern Ocean sediments. Paleoceanography, 118-134. Stuiver, M., Reimer, P.J., Bard, E., Beck, J.W., Burr, G.S., Hughen, K.A., Dromer, B., McCormac, G., Van der Plicht, J., Spurk, M., 1998. INTCAL98 radiocarbon age calibration, 24,000-0 cal BP. Radiocarbon 40,1041-1083. Taylor, R.S., McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell Scientific, Boston, 312pp. 13 Thompson, E.I., Schmitz, B., 1997. Barium and the late Paleocene 8 C maximum: Evidence of increased marine surface productivity. Paleoceanography 12, 239-254. 62 Thomas, A.C., Huang, F., Strub, P.T., James, C, 1994. Comparison of the seasonal and interannual variability of phytoplankton pigment concentrations in the Peru and California Current systems. Journal of Geophysical Research 99, 7355-7370. Thomson, R.E., 1981. Oceanography of the British Columbia Coast. Canadian Special Publication of Fisheries and Aquatic Sciences 56, 291pp. Thunell, R.C, Mortyn, P.G., 1995. Glacial climate instability in the Northeast Pacific Ocean. Nature 376, 504-506. Toggweiler, J.R., 1999. Variations of atmospheric C02 by ventilation of the ocean's deepest water. Paleoceanography 14, 571-588. Villanueva, J., Grimalt, J.O., Cortijo, E., Vidal, L., Labeyrie, L., 1997a. A biomarker approach to the organic matter deposited in the North Atlantic during the last climatic cycle. Geochimica et Cosmochimica Acta 61, 4633-4646. Villanueva, J., Pelejero, C, Grimalt, J.O., 1997b. Clean-up procedures for the unbiased estimation of C37-C39 alkenones sea surface temperatures and terrigenous n-alkane inputs in paleoceanography. Journal of Chromatography A757, 145-151. von Breymann, M.T., Emeis, K.-C, Suess, E., 1992. Water depth and diagenetic constraints on the use of barium as a palaeoproductivity indicator. In: Summerhayes, CP., Prell, W.L., Emeis, K.-C. (Eds.), Upwelling Systems: Evolution Since the Early Miocene. Geological Society Special Publication 64, pp. 273-284. Werne, J.P., Hollander, D.J., Lyons, T.W., Peterson, L.C., 2000. Climate-induced variations in productivity and planktonic ecosystem structure from the Younger Dryas to Holocene in the Cariaco Basin, Venezuela. Paleoceanography 15, 19-29. Wu, J., 1997. The carbon and nitrogen isotope compositions of particulate organic matter in the subarctic northeast Pacific Ocean. PhD. Thesis, University of British Columbia, 252p. Wu, J., Calvert, S.E., Wong, C.S., 1997. Nitrogen isotope variations in the subarctic northeast Pacific: relationships to nitrate utilization and trophic structure. Deep Sea Research 44, 287-314. 63 Wu, J., Calvert, S.E., Wong, C.S., 1999a. Carbon and nitrogen isotope ratios in sedimenting particulate organic matter at an upwelling site off Vancouver Island. Estuarine, Coastal and Shelf Science 48, 193-203. Wu, J., Calvert, S.E., Wong, C.S., Whitney, F.A., 1999b. Carbon and nitrogen isotopic composition of sedimenting particulate material at Station Papa in the subarctic northeast Pacific. Deep Sea Research (II) 46, 2793-2832. Yorath, C.J., Nasmith, H.W., 1995. The Geology of Southern Vancouver Island. Orca Book Publishers, Victoria, Canada, 172pp. 64 3. Intensification of the Oxygen Minimum Zone in the Northeast Pacific during the Last Deglaciation: Ventilation or Export Production? 3.1 Introduction Extensive evidence indicates that the intensity (i.e., degree of oxygen depletion) of the oxygen minimum zone (OMZ) along the eastern margin of the North Pacific Ocean has fluctuated significantly over time. Laminated sediments have been observed in cores collected from the Gulf of California (Keigwin and Jones, 1990), partially isolated basins within the California Borderlands region (e.g., Santa Barbara Basin - Behl and Kennett, 1996; and Santa Monica Basin - Stott et al., 2000b) and from the open continental margin off California (Gardner and Hemphill-Haley, 1986; Anderson et al., 1987; Dean et al, 1994; van Geen et al., 1996) and northwestern Mexico (Ganeshram, 1996). The light and dark layers that constitute a single lamination are seasonal layers preserved when bottom water containing <0.1 ml/1 oxygen inhibits bioturbation (e.g., Thunell et al., 1995; Pilskaln and Pike, 2001). The laminated deposits are interbedded with partly- to well-homogenized sediments that were bioturbated when the bottom water was more oxygenated. In addition to preserving laminations, a low dissolved oxygen concentration in the bottom water also directly influences the species of benthic foraminifera that occur in the underlying sediments (Kaiho, 1994). In cores collected from the California margin for example, a foraminiferal assemblage tolerant of very low oxygen conditions alternates with an assemblage that requires better oxygenation (Cannariato et al., 1999; Cannariato and Kennett, 1999), and in the Santa Barbara Basin the appearance of a foraminiferal assemblage tolerant of low oxygen conditions is synchronous with the change to laminated sediments (Cannariato et al., 1999). The concentrations of redox-sensitive trace metals (e.g., Mo) in sediments have also varied on glacial to interglacial and shorter time scales (Nammeroff, 1996; Dean et al., 1997; Zheng et al., 2000; Ivanochko, 2001) further indicating that bottom water oxygen levels have 65 fluctuated. Collectively, the accumulated evidence suggests that the OMZ along the eastern margin of the North Pacific was less intense (i.e., relatively oxygen-rich) during colder intervals such as the Younger Dryas, MIS 2, and stadials within MIS 3 and 4, and more intense (i.e., relatively oxygen-depleted) during warm periods (Keigwin and Jones, 1990; Kennett and Ingram, 1995; Behl and Kennett, 1996). Not only has the intensity of the OMZ fluctuated, but its boundaries may have shifted vertically causing the OMZ to expand and contract over time (van Geen et al., 1996; Cannariato and Kennett; 1999; Stott et al., 2000b). The intensity of the OMZ is a function of two primary variables: i) ocean circulation, hence-forth referred to as ventilation; and ii) oxygen consumption (Wyrtki, 1962). Ventilation refers to any advective process that transfers surface conditions, in this instance high oxygen concentrations, to subsurface waters (Van Scoy and Druffel, 1993). It may occur via the cropping out of a particular isopycnal at the ocean surface (and then its' subduction), by the formation of a denser water mass during, for example, winter cooling and sea ice formation, and/or by vertical mixing (Talley, 1991). Oxygen consumption refers to the utilization of oxygen during degradation of organic matter, both as it settles through the water column and after deposition on the sea floor. The higher the rate of export of labile organic matter from surface waters, the greater the oxygen depletion of the underlying intermediate waters. There is considerable debate over which of these mechanisms, ventilation or productivity, or some combination of the two was responsible for changes in intensity of the OMZ along the eastern margin of the North Pacific. At present, ventilation of the OMZ in the northeast Pacific Ocean reflects a balance between the relatively oxygen-rich North Pacific Intermediate Water (NPIW) and the oxygen-poor Subtropical Subsurface Water (SSW). NPIW is identified by a salinity minimum at a density of 26.7 to 26.9 GQ (Reid, 1965). This water mass occurs throughout the North Pacific Subtropical Gyre and extends as far north as the Gulf of Alaska (Talley, 1993). NPIW forms in the region of the Sea of Okhotsk where Western Subarctic Water, which is carried south by the Oyashio Current, mixes with cold and relatively fresh Sea of 66 Okhotsk Water (Talley, 1991; Freeland et al., 1998; Wong et al. 1998). The resulting water is then modified by mixing with Kuroshio Water in the region just to the east of the Sea of Okhotsk and NPIW is produced (Talley, 1993; Kono, 1996). There is also evidence of ventilation of NPIW within the Alaskan gyre due to winter mixing and occasional cropping out at the surface of the 26.8 <JQ density surface (Van Scoy et al., 1991a and b; You et al., 2000). A number of studies have suggested that enhanced formation of NPIW can explain the weakening of the OMZ during cold climatic periods (Duplessy et al., 1988; Keigwin and Jones, 1990; Kennett and Ingram, 1995; van Geen et al., 1996; Behl and Kennett, 1996; Keigwin, 1998; Zheng et al., 2000). Indeed, coupled ocean-atmosphere climate models predict enhanced NPIW formation during the Younger Dryas (Mikolajewicz et al, 1997) and the Last Glacial Maximum (Ganopolski et al., 1998). Better ventilation of NPIW during the last glacial has been inferred from relatively high 5 C values of benthic foraminifera which suggest the presence of a younger, relatively nutrient-poor water mass in the northwest Pacific at intermediate water depths (Duplessy et al., 1988; Keigwin et al., 1998). Radiocarbon data for coeval benthic and planktonic foraminifera (i.e., benthic-planktonic age differences) for a core from northwestern Pacific also suggest increased ventilation during the last glacial, as well as decreased ventilation (i.e., older intermediate water) during the deglacial (Duplessy et al., 1989). On the eastern side of the Pacific, radiocarbon data from the Santa Barbara Basin indicate increased ventilation during the Last Glacial Maximum and Younger Dryas (Ingram and Kennett, 1995; Kennett and Ingram, 1995). Radiocarbon data from cores collected on the open California margin are more ambiguous, indicating decreased ventilation between 11 and 9 kyr B.P. at a water depth of 800 m at 35°N (van Geen et al., 1996). However, 14C measurements made on ODP Hole 1019 collected from the continental slope (980 m) off northern California (41 °N) suggest increased, not decreased, ventilation during the early Holocene and Bolling-Allerod at the same time as intensification of the OMZ occurred (Mix et al., 1999). 67 Most previous studies that have focused on the history of the OMZ in the eastern Pacific have dwelt on changes in NPIW formation and have ignored possible changes in the supply and/or oxygen content of Subtropical Subsurface Water (SSW; Wyrtki, 1967). This water mass is characterized by relatively high temperature, high salinity, high nutrient content and a low oxygen concentration (Lynn and Simpson, 1987). Oxygen depletion of SSW is the result of high productivity in the Equatorial East Pacific and subsequent degradation of the exported organic matter (Wyrtki, 1967). This oxygen-depleted water is then transported northward along the west coast of North America as far north as Vancouver Island, Canada by the California Undercurrent (Reed and Halpern, 1976; Mackas et al., 1987). However, mixing with adjacent water masses modifies the physical and chemical properties of SSW as it moves northward (e.g., decreasing temperature - Halpern et al., 1978; decreasing 815N values - Kienast et al., 2002). Off Vancouver Island the highest percentage of SSW is observed at a depth of 100 to 300 m where the California Undercurrent (CUC) is strongest (Reed and Halpern, 1976). Weaker flow below 300 m allows increased mixing of SSW with other water masses; however, SSW is still recognizable as deep as 1300 m (Reed and Halpern, 1976). Because the percentage of SSW in intermediate waters is controlled by the strength of the CUC it is probable that changes in CUC strength would influence the oxygen concentration of intermediate waters. During the Last Glacial Maximum, the Aleutian Low was situated more to the east, over the Northeast Pacific (COHMAP, 1988) and the Subarctic Gyre expanded southward (Doose et al., 1997). As a result, the CUC would not have extended as far north as Vancouver Island and thus the supply of imported oxygen-poor SSW should have been cut off. In regions that remained within the California Current System it is also possible that the supply of oxygen-poor waters was reduced during the glacial, but for different reasons. Evidence from the continental slope off Mazatlan, Mexico indicates that the OMZ was relatively oxygen-rich during the Last Glacial Maximum (Ganeshram et al., 1995). Thus, SSW carried northward by the CUC would have contained more oxygen than at present. 68 Also, the strength of the California Current appears to have been significantly reduced north of ~35°N (Doose et al., 1997) and presumably so was the strength of the CUC. Weakening of the CUC would have allowed increased mixing between SSW and more oxygen-rich water masses such as Pacific Subarctic Water and NPIW. Thus increased ventilation may have occurred even if the strength of NPIW formation had remained constant. From what is known at present, it is still unclear whether changes in ventilation are responsible for variations in the intensity of the OMZ along the northeastern margin of the Pacific Ocean. Furthermore, all previous research has focused on the OMZ off southern Oregon, California and Mexico but nothing is known about the history of the OMZ north of 42°N. This study was initiated to fill this gap in our knowledge. The study area is located off the west coast of Vancouver Island, Canada (48°54'N, 126°53'W; Fig. 3.1) and sits at the northern end of the California Current System where the eastward-flowing Subarctic Current and North Pacific Current split into the northward flowing Alaska Current and southern flowing California Current (Thomson, 1981). In this region high primary productivity is driven by seasonal (spring-summer), wind-driven upwelling which is related to the northerly position (38°N) of the North Pacific High (Huyer, 1983). In winter, the North Pacific High shifts southward to 28°N and is replaced by the expanded Aleutian Low. This leads to a reversal of the wind direction, cessation of upwelling off Vancouver Island, and stronger winter storms crossing into North America. Over the last 16 calendar kyr B.P. primary productivity and organic carbon flux to the sediment has varied substantially in this region (Chapter 2). Variability in organic carbon production may have influenced the intensity of the OMZ locally. However, the study area is also proximal to regions of NPIW formation and ventilation. Therefore, if variations in the intensity of the OMZ were related to changes in ventilation of intermediate waters by NPIW, we would expect to see evidence of this off Vancouver Island. The primary objectives of this study are thus to investigate changes in the intensity of the OMZ, focusing on the period 69 127"W 126°W 125°W Fig. 3.1. The study area is located off the west coast of Vancouver Island, British Columbia, Canada (Inset). Sediment cores were collected from Station JT96-09 that is located on the continental slope at a water depth of920m. 70 between 16 and 10 calendar kyr B.P., and to determine whether such changes are the result of fluctuations in ventilation and/or export productivity. Analyses were conducted on Piston Core JT96-09, raised from a water depth of 920 m, within the core of the present-day OMZ. Changes in the intensity of the OMZ are inferred from trace metal data and benthic-forminifera species information. Radiocarbon dating of benthic and planktonic foraminiferal pairs is used to determine if variability in the intensity of the OMZ is the result of changes in ventilation, while variations in palaeoproductivity are estimated using a variety of proxies (e.g., organic carbon, biogenic barium, opal and alkenone concentrations and/or mass accumulation rates). These data are discussed in detail in Chapter 2. 3.2 Materials and methods Sediment cores were collected from the continental slope (920 m water depth) west of Vancouver Island, British Columbia, Canada during a 1996 Canadian Joint Global Ocean Flux Study cruise. Data are presented here for a 374 cm long Piston Core (JT96-09pc) and a corresponding 40 cm long multicore (JT96-09mc), both collected at the same site (48°54'N, 126°53'W; Fig. 3.1). The multicore and upper 51 cm of the piston core are composed of a homogeneous, organic carbon-rich (1 to 3 wt.%) and carbonate-poor (<1 wt.%), olive green mud. In the piston core this mud is underlain by 85 cm of grayish green clay and at the base of this clay is a 16 cm thick sandy turbidite. The remainder of the core (222 cm) is a dense, cohesive gray clay. It is important to note that no laminated sediments were observed. The sediment-water interface was captured in the multicore and a comparison of geochemical data from the multi- and piston cores suggests that ~12 cm of sediment, not 20 cm as originally stated in Kienast and McKay (2001), were lost off the top of the piston core during its collection. Piston core depths have been corrected for this loss (+12 cm) and for 71 the presence of the turbidite (-16 cm), and the records for both cores have been merged to yield a composite record for site JT96-09. The age model for JT96-09 is based on nine accelerator mass spectrometry (AMS) radiocarbon dates measured at the Lawrence Livermore National Laboratory on a mixed assemblage of planktonic foraminifera (N. pachyderma right- and left-coiling and G. bulloides). Results were converted from radiocarbon to calendar ages using CALIB 4.3 (Stuiver et al., 1998) assuming a reservoir age of 800 years for radiocarbon ages younger than 12.5 14C kyr and a reservoir age of 1100 years was applied to older ages. Details of the age model are discussed in Kienast and McKay (2001); however, as a result of the new estimate of sediment loss during coring, the age model for the Holocene differs slightly from that previously published. Sedimentation rates, which were calculated by assuming linearity between calendar-age dates, range from 5 cm/kyr during the Holocene up to 169 cm/kyr during the deglacial (Fig. 3.2). Therefore, although Core JT96-09 does not extend back to the Last Glacial Maximum it does yield a high resolution record of the last deglaciation. Alkenone palaeothermometry measurements previously conducted on Core JT96-09 yielded evidence of rapid sea-surface temperature (SST) fluctuations during the deglacial (Kienast and McKay, 2001). The pattern of deglacial SST fluctuations is remarkably similar to temperature fluctuations in the GISP-2 ice core record suggesting that the Boiling-Allerod and Younger Dryas events are recorded in Core JT96-09. More importantly these events appear to be nearly synchronous with those in GISP-2, although the match is not perfect (see Kienast and McKay, 2001). Offsets between the two records, in some instances in the order of hundreds of years, most probably reflect errors in the age model of Core JT96-09 resulting from: i) problems inherent to radiocarbon dating (e.g., 14C plateaus); ii) a poorly known reservoir age (e.g., using a reservoir age of 800 yrs rather than 1100 yrs for samples older than 12.0 14C years increases most ages by -300 yrs); iii) bioturbation, although the effects of this are limited by the high sedimentation rates during the deglacial; and iv) large errors associated with the radiocarbon dating of small samples. 72 Possible changes in intermediate water ventilation are determined by calculating the age difference between benthic and planktonic foraminifera separated from the same sample. This method assumes that the effects of bioturbation can be ignored, a reasonable assumption given the high sedimentation rates that prevailed during the deglacial (»10 cm/kyr). However, the age difference may be biased by another influence. Uvigerina and Bolivina spp. are infaunal organisms and probably grew at depth alongside the shells of previously deposited planktonic foraminifera. This would increase the benthic-planktonic age difference, but given the high sedimentation rates, the quantitative impact of such an effect should be trivial. No attempt has been made to correct for it. A further complication may arise from the degradation of "old" (i.e., I4C-depleted) organic matter and ensuing addition of "old" C02 to the porewaters that could theoretically increase the apparent age of deeper-dwelling infaunal species. This possibility is discussed in more detail below. No foraminifera occur in the upper 30 cm of the composite core (~0 to 5 kyr B.P.), while in sediments deposited between 5 and 10 kyr B.P. foram shells are present but extensive dissolution is observed (i.e., fragments are abundant). The paucity of shells in the Holocene deposits can be attributed to the highly corrosive nature of North Pacific waters (Zahn et al., 1991; Karlin et al., 1992) and the adverse effect that degradation of organic matter has on carbonate preservation (De Lange et al., 1994; Jahnke et al., 1997). In contrast, foraminifera are abundant and well preserved in sediments older than 10 kyr B.P. apparently reflecting less-corrosive conditions in the past. No single species of benthic or planktonic foraminifera is found throughout the entire core and therefore, it was necessary to use a variety of taxa for 14C dating. Radiocarbon data were obtained for 10 benthic-planktonic foraminiferal pairs spanning the period between 10.0 and 15.6 kyr B.P. Planktonic samples comprise a mixture of N. pachyderma (right and left coiling) and G. bulloides. Most benthic samples consist of either Uvigerina spp. or Bolivina argentea. In two instances it was necessary to use a mixture of Bolivina spp. (B. argentea and B. spissa) to obtain enough material; however, these results were not used for reasons that will be discussed below. 73 Sample size ranged from <1 to 4.8 mg carbonate (usually 1.3 to 4.8 mg). Prior to radiocarbon dating, samples were sonicated in methanol and briefly etched with 0.000IN HC1. When conducting stable isotope studies of benthic foraminifera, the preference is to use epifaunal species such as Cibicidoides which obtain their carbon from the overlying water (McCorkle et al., 1990). In Core JT96-09 no such species are present and thus a combination of Uvigerina spp. (3-15 individuals) and Bolivina argentea (>10 individuals) were used. As noted earlier, these are shallow-dwelling infaunal taxa (McCorkle et al., 1990; Stott 2000a). The isotopic composition of Bolivina pacifica, a deeper dwelling infaunal species (Stott et al., 2000a), was also measured. Prior to analysis foraminifera were sonicated in methanol and roasted under vacuum for 30 minutes at 430°C. The samples were then reacted in a common orthophosphoric acid bath at 90°C and analyzed using a PRISM series 1, triple ion collector mass spectrometer. Data are reported in the standard 8-notation relative to Vienna Peedee belemnite (VPDB). The reproducibility of three in-house laboratory standards (Mexical, UQ6 and a foraminiferal standard), which were calibrated using the international standard NBS-19, was ± 0.15%ofor5,oO. Trace metal concentrations (i.e., Re, Ag, Cd, Mo and U) were measured by isotope-dilution inductively-coupled plasma mass spectrometery. Sample preparation involved adding known amounts of isotopically-enriched spike solutions to -20 mg of powdered sediment. Samples were then microwave digested in a mixture of concentrated HN03, HC1 and HF. The digests were evaporated on a hotplate overnight and then re-digested in 5N HC1. Aliquots were taken for Mo and U analysis. The remaining sample was run through an anion exchange column (Dowex 1-X8 resin) to remove Zr and Nb which form compounds (ZrO, ZrOH and NbO) that interfere with the analysis of Ag, and Mo that forms MoO" which interferes with Cd analysis. Further details of the sample preparation can be found in Ivanochko (2001). To check precision a University of British Columbia sediment standard (SNB) was analyzed with each batch of samples. The resulting RSD (la) for Re, U, Mo, Cd 74 and Ag analyses were 11 %, 9 %, 7 %, 10 % and 8 %, respectively. Accuracy was assessed by measuring the concentrations of these metals in the NRC sediment standard MESS-1 and was 8 % or better for Re, Mo and U, and -14 % for Cd. 3.3 Results Radiocarbon data are summarized in Table 3.1. Benthic-planktonic age differences 2 2 range from 610 to 1550 years. The errors for these data (i.e., the square root of a + b ; where a and b are the individual errors) range from ± 92 to 305 years. Errors greater than 200 years are the result of using very small samples (<1 mg carbonate). For sample 1 (Table 3.1 and Fig. 3.2), the benthic-planktonic age difference was calculated using two separate benthic species (i.e., Uvigerina spp. and Bolivina argentea) and the results are identical within the error (950 ± 92 and 890 ±114 years, respectively). Therefore, changes in benthic-planktonic age differences observed in Core JT96-09 are most probably not the result of using two different benthic foraminifera taxa. However, this conclusion is only valid because Bolivina argentea and Uvigerina spp. are shallow infaunal species that obtain most of their carbon from the overlying bottom water. The same assumption does not hold true when a mixture of Bolivina spp. is used because different species live at different depths in the sediment. Deeper-dwelling species very probably obtain a larger amount of their carbon from the degradation of organic matter. In the deglacial sediments of Piston Core JT96-09 bulk organic matter is 2000 to 3000 years older than planktonic foraminifera in the same sample (Chapter 2). Therefore, it is probable that deeper-dwelling infaunal species are anomalously depleted in 14C. As a result, radiocarbon data for samples that contain mixed Bolivina spp. (e.g., samples 3 and 5; Table 3.1 and Fig. 3.2), which include deeper dwelling species, are likely biased toward older ages and benthic-planktonic age differences will be larger. These data are therefore considered unreliable and are not discussed further. 75 14C Age (kyrs) 9 10 11 12 13 14 15 S 50 100 150 200 a <u Q a> h O U -o 5 250 o v ^ 300 350 400 1 1 1 1 -1—I I 1 ii i i1 i r n— . •? .a® '• : '• • '• : \ W \@: -• Mixed Planktonics • Benthics-Uvigerina spp. • Benthics-Bolivina spp. • ••• _1_ . .. 1. . . . 1 . . . . i. . .. 1. , • • 5 cm/kyr 14 cm/kyr 20 cm/kyr 143 cm/kyr 30 cm/kvr 113 cm/kyr 169 cm/kyr 152 cm/kyr 48 cm/kyr Fig. 3.2. Plot of 14C ages of planktonic and benthic foraminifera versus sample depth in Core JT96-09. Sedimentation rates, which are indicated along the right axis, were calculated using ages of planktonic foraminifera and assuming a linear sedimentation rate between age picks. Circled numbers correspond to sample numbers in Table 3.1. 76 Table 3.1. Radiocarbon ages of planktonic and benthic foraminifera in Composite Core JT96-09. Sample Depth Corrected Calendar Age l4C Age l4C Age Benthics Benthic-Planktonic Range (cm) Depth (cm) (kyr B.P.) Mixed Planktonics Uvigerina spp. Bolivina spp. Age Difference (yrs) 1 35-36 47.5 10.03 1 9760 ± 70 10710 ±60 10650 ±90 950 ± 92,890 ± 2 45-46 57.5 10.76 1 9360 ± 240 10910 ±70 1550 ±250 3 65-66 77.5 12.24 1 11210 ±120 12380 ±2803 1170 ±305 4 75-76 87.5 12.73 1 11500 ±110 12110 ± 80 610 ±136 5 90-91 102.5 12.84 1 11600 ±80 12700 ± 7o3 1100±106 6 100-101 112.5 13.172 12460 ± 120 13170 ± 100 710± 156 7 130-131 142.5 13.43 2 12640 ±90 13500 ±80 860±120 8 265-266 261.5 14.142 13410 ±80 14290 ±110 880±136 9 290-291 286.5 14.302 13520 ±70 14350 ± 120 830±139 10 350-351 346.5 15.572 14140 ±70 14830 ±280 680 ± 289 1 Calculated using a reservoir age of 800 years. 2 Calculated using a reservoir age of 1100 years. 3 A mixture of Bolivina spp. were used to make these measurements. All other Bolivina samples are monospecific. 77 The single late glacial sample yields a benthic-planktonic age difference of 680 ± 289 years. Benthic-planktonic age differences for the deglacial range from 830 to 880 years ± 139 to 136 years for the period between 14.3 and 13.4 kyr B.P. and then decrease to 610 ± 136 and 710 ±156 years between 13.2 and 12.7 kyr B.P. Due to a paucity of foraminifera no data are available for the period from 12.7 to 10.8 kyr B.P. The highest benthic-planktonic age difference (1550 ± 250 years; sample 2 in Table 3.1) occurs at the end of the Pleistocene (10.76 cal. kyr B.P.); however, this large value appears to be partly the result of an anomalously young planktonic age (Fig. 3.2). At the Pleistocene-Holocene boundary the benthic-planktonic age difference ranges from 890 ± 114 to 950 ± 92 years. Given the scale of the error estimates, and excluding sample 2, we cannot conclusively state that real differences exist between samples. However, some intriguing changes in benthic-planktonic age differences over time are suggested by the data. Alkenone palaeothermometry of Core JT96-09 has identified a number of rapid SST fluctuations that appear to be coeval with the Bolling-Allerod and Younger Dryas (Kienast and McKay, 2001). When the benthic-planktonic age difference data are placed into this time frame it is observed that values may be slightly higher during the Boiling and at the start of the Allerod and relatively low throughout the Allerod (Fig. 3.3). Due to the lack of foraminifera no data are available for the Younger Dryas. Just prior to the Holocene the benthic-planktonic age difference reaches its highest value (1550 ± 250 years) and then decreases at the Pleistocene-Holocene boundary (Fig. 3.3). Trace metal data for composite Core JT96-09 are presented in Figure 3.4. Rhenium concentrations range from <1 to 65 ng/g, U from 1.1 to 5.8 (xg/g, Cd from <0.1 to 1.4 ug/g, Mo from 0.3 to 3.7 pg/g and Ag from 71 to 395 ng/g. To correct for possible fluctuations in metal content due to changes in dilution by biogenic components (i.e., carbonate and organic matter) over time metal concentrations have been normalized to Al. Concentrations and metal/Al ratios are low during the glacial and early deglacial (16.0 to 13.5 kyr B.P.). At 13.5 kyr B.P. (i.e., the start of the Allerod) these ratios increase substantially and remain high until 78 Age Difference (yrs) 0 500 1000 1500 2000 16 -|—i—i—i—i—|—i—i—i—i—i—i—i—i—i—i—i—i—i—i—I 5 10 15 20 25 Sea Surface Temperature (°C) Fig. 3.3. Benthic-planktonic age differences are plotted against calendar age. The two periods of OMZ intensification are indicated by the shading (i.e., zones 1 and 2). The sea-surface temperature record of Kienast and McKay (2001) is also shown and the Boiling (B), Allerad (A) and Younger Dryas (YD) events are labeled. Circled numbers correspond to sample numbers in Fig. 3.2 and Table 3.1. 79 VC 00 a, 3 BID < DX) "5k 3 ID i o 3 O T3 " S3 ON H O 0) S3 d CO c " * ° o <o Z b U _ J3 ^ 03 - *^ • «> .g <u co OH ."S ^ § 03 .§ ^ <N O +3 .2-2 ™Z $ O • 23 N b <U O O CO CO _g M cs G w c ^ . n O J> 3 ^ .5 CO s a. 1) S o S3 X)T3 «» O l-i cO is3^ ^ S * -§ §: 2 C o w ft tU X S3 .2 g co "33 o in tS S3 -fa -O 3 o"0 * s > g I) s o u 8 .2 N IIS U C 1H CO (U ft O C+H S3 o o CO S3 U CD co b u< .O ^ S 13 g * fa u Si u co ® CSC^ « 5 d> Q -—' g a u a .22 'a „v O JC3 3 O OH ta H 2 ^ av 80 the Younger Dryas at -12.7 kyr B.P. Metal/Al ratios are low throughout the Younger Dryas, and then rise again at 11.0 kyr. The is some evidence that in the early part of the Younger Dryas the low trace metal content might reflect dilution by coarser-grained materials as the Zr/Al ratio is Early to mid Holocene sediments host the highest trace metal enrichments and have correspondingly high metal/Al ratios. The Ag/Al ratio remains high in the late Holocene; however the Re/Al, U/Al, Cd/Al and Mo/Al ratios decrease to near glacial values. No foraminiferal counts were conducted on Core JT96-09 because of the limited number of foraminifera present. However, while picking specimens for stable isotopic analysis and radiocarbon dating it became apparent that certain species of benthic foraminifera occur in specific zones of the core. Two zones are characterized by abundant Bolivina spp., (i.e., B. argentea, B. spissa and B. pacified), but no Uvigerina spp. (zones 1 and 2; Fig. 3.5a). In comparison, Holocene sediments contain a mixture of Uvigerina spp. and Bolivina spp. (only B. argentea), and sediments older than 13.5 kyr B.P. contain Uvigerina spp. and few, if any, Bolivina spp. The 6lsO results are shown in Figure 3.5a. The large decrease in 8180 values of benthic foraminifera at -13.5 kyr B.P. (Fig. 3.5a) most probably represents the mixing in to the ocean of Meltwater Pulse 1A (Fairbanks, 1989). The decrease in benthic 8180 is contemporaneous with an increase in sea-surface temperature (SST) that is inferred to be the Allerod. This timing is similar to that observed in other marine records (e.g., Tahiti corals; Bard, 1996), suggesting that the age model for Core JT96-09 is reasonable and that sea-surface temperature fluctuations in this core are coeval with temperature fluctuations in the GISP-2 ice core, as suggested by Kienast and McKay (2001). 3.4 Discussion 3.4.1 Evidence of OMZ intensification 81 © o 82 The OMZ, defined as that portion of the water column where oxygen is <0.5 ml/1, extends from approximately 750 to 1300 m water depth off the west coast of Vancouver Island (see data in Appendix A5). During the 1996 research cruise the lowest oxygen concentration (0.3 ml/1) was measured at a water depth of 920 m (i.e., the site where Core JT96-09 was collected). Unlike cores from the California and Mexican continental margins, no laminated sediments have been preserved in Core JT96-09 over the past 16 kyr B.P. However, this does not necessarily imply that fully oxygenated conditions existed over this period because laminated sediments are only preserved when bottom water oxygen levels drop below 0.1 ml/1 (Behl and Kennett, 1996). Significant variations in bottom water oxygen concentrations (i.e., OMZ intensity) over the past 16 kyr are inferred from changes in the accumulation of redox-sensitive trace metals and by changes in the assemblage of benthic foraminifera. The reduction and subsequent precipitation of Re and U begins once porewater 02 is depleted and this leads to their enrichment in suboxic and anoxic sediments (Ravizza et al, 1991; Colodner et al., 1993; Crusius et al., 1996). Silver and Cd have a single redox state, but form insoluble sulphides when trace amounts of H2S are available (Koid et al., 1986; Rosenthal et al., 1995). This leads to minor Ag and Cd accumulation in suboxic sediments and large accumulations in anoxic sediments. In comparison, Mo enrichment is only observed in anoxic sediments (Francois et al., 1988; Emerson and Huested, 1991; Crusius et al., 1996; Ivanochko, 2001; Morford et al., 2001). Molybdenum accumulation is related to the conversion of Mo042" to MoS42" (i.e., thiomolybdate) which is readily adsorbed onto Fe-bearing particles such as Fe-sulphides (Helz et al., 1996). The formation of thiomolybdate only occurs when the porewater H2S concentration exceeds ~11 urn (Erickson and Helz, 2000) and thus Mo enrichment is restricted to fully anoxic sediments. At present, sediments deposited within the OMZ off Vancouver Island become suboxic within millimetres of the sediment-water interface (Chapter 4). However, near-surface sediments never become fully anoxic and thus Mo enrichment does not occur in the upper 50 cm (i.e., <1 pg/g Mo; Fig. 83 3.4). This observation is consistent with measured bottom water oxygen concentrations (> 0.3 ml/1) because, in general, Mo accumulation only occurs when oxygen levels drop below ~0.2 ml/1 (Zheng et al., 2000). In comparison, two periods of Mo enrichment are observed in older sediments (zones 1 and 2; Fig. 3.4) and these are accompanied by Re, U, Cd and Ag enrichments. The first episode of marked trace metal enrichment is observed in the deglacial clay deposited during the Allerod (Zone 1; Fig. 3.4). There is no evidence that this enrichment is result of metal remobilization due to oxygen influx (i.e., burndown) and subsequent reprecipitation in underlying reduced sediments because the distribution of all redox-sensitive metals is similar. Burndown commonly re-distributes elements at different depths due to their different chemical behaviours (Colodner et al., 1992; Thomson et al., 1993; Thomson et al., 1995; Crusius and Thomson, 2000). Therefore, metal enrichment in Zone 1 must reflect the development of anoxic conditions within the sediment. Such conditions may have developed as a result of: i) increased sedimentation rate and corresponding decrease in oxygen influx; ii) decreased ventilation of the bottom water; and/or iii) increased carbon flux to the sediment. Increased sedimentation is ruled out as the cause of trace metal enrichment because the sedimentation rate was substantially higher (>150 cm/kyr; Fig. 3.2) during deposition of the trace metal-poor sediments in the Boiling, between 14.3 and 13.5 kyr B.P. (290 and -130 cm sub-bottom depth, respectively). Whether oxygen depletion in near-surface sediments was the result of decreased ventilation of bottom waters and/or increased organic carbon flux to the sediment, which also would have caused water column oxygen content to decline, is discussed in Sections 3.4.2 and 3.4.3. In either case, trace metal data imply that the OMZ was more intense during the Allerod. The second interval of trace metal enrichment occurs in sediments deposited between 11 and 8 kyr B.P. (Zone 2; Fig. 3.4). Several factors, acting collectively, appear to have produced the observed enrichments. The increase in Re at -4 kyr B.P. (i.e., 20 cm below the sediment-water interface) reflects the approximate depth where Re reduction and 84 accumulation are occurring at present (Chapter 4). Relatively low sedimentation in the Holocene (~5 cm/kyr) coupled with sufficient oxidant demand has allowed significant authigenic metal enrichment to occur well below the sediment-water interface. Thus, trace metal enrichments do not necessarily indicate stronger reducing conditions at, or shortly after, sediment deposition. Although this observation constrains detailed interpretation of OMZ history during the Holocene, these data clearly indicated that near-surface sediments at Station JT96-09 have been continuously reducing since the Younger Dryas, and imply low bottom water oxygen concentrations throughout the Holocene. This conclusion is consistent with deductions made to the south along the eastern margin of the North Pacific (Anderson et al., 1987; Ganeshram et al., 1995; Behl and Kennett, 1996; Cannariato and Kennett, 1999), all of which imply that a relatively intense OMZ has been a permanent fixture during the Holocene. Molybdenum enrichment in the Allerod is contemporaneous with a shift from a Uvigerina-dominated benthic foraminifera assemblage to one composed almost entirely of Bolivina spp. (Fig. 3.5a). A similar species shift occurs between -10 and 11 kyr B.P. which is also a period of enhanced Mo accumulation. Bottom water oxygen concentration has a direct influence on the species of benthic foraminifera that occur in sediments (Kaiho, 1994). Tio/j'vma-dominated assemblages are typical of the most intense portions of the OMZ in the eastern Pacific and thrive at dissolved oxygen levels of < 0.3 ml/1 (Mullins et al., 1985; Sen Gupta and Machain-Castillo, 1993). In contrast, Uvigerina spp. prefer a more oxygenated environment. This relationship has been documented at many locations along the northeastern Pacific margin (e.g., Mullins et al., 1985; Quinterno and Gardner, 1987) and it appears to hold true in the past (Behl and Kennett, 1996; Cannariato et al., 1999; Cannariato and Kennett, 1999). The presence of Uvigerina spp. also may be related to a high organic carbon supply to the sediment (Quinterno and Gardner, 1987), but this association is not observed if bottom water oxygen is low (Kaiho, 1994). We conclude that the occurrence of a benthic foraminifera assemblage composed almost exclusively of Bolivina spp. in sediments 85 characterized by Mo enrichment provides further evidence that bottom water oxygen levels were very low (< 0.3 ml/1). The combination of trace element and benthic foraminifera data suggest that intensification of the OMZ in the region off Vancouver Island occurred during the deglacial between 13.5 and 12.7 kyr B.P. (i.e., Allerod) and possibly again between 11.0 and 8.0 kyr B.P. In general, intensification occurred contemporaneously with times of high SST and brackets the Younger Dryas (Fig. 3.5). In this respect the timing of OMZ intensification off Vancouver Island is similar to the timing off California and Mexico, with one notable difference. Intensification of the OMZ in the southern CCS apparently occurs at the beginning of the Boiling (Cannariato and Kennett, 1999; Mix et al., 1999; Zheng et al, 2000, Ivanochko, 2001) while off Vancouver Island it is delayed until the Allerod, a lag of-1.5 kyr. 3.4.2 Ventilation changes? The age of intermediate water at a depth of 920 m is recorded by benthic foraminifera while the age of surface waters is recorded by planktonic foraminifera. The difference in the ages of benthic and planktonic foraminifera obtained from the same sample therefore establishes the age difference between the intermediate and surface waters. If the intensification of the OMZ inferred from trace element and faunal data discussed above was the result of decreased ventilation, the benthic-planktonic (B-P) age difference should be greater, reflecting the reduced influence of relatively young NPIW and the increased influence of older, oxygen-depleted SSW. Lack of foraminifera in surface sediments precludes direct determination of the modern B-P age difference. However, using A14C water column data of Ostlund and Stuiver (1980) for the North Pacific off the coast of California we estimate that benthic foraminifera from a depth of 900 m off Vancouver Island should be -1780 years old, or perhaps slightly 86 younger given that our study area lies closer to regions of NPIW ventilation. The present age of planktonic foraminifera growing within surface waters (i.e., reservoir age of surface water) is ~800 years (Robinson and Thompson, 1981; Southon et al., 1990). Thus, the modern B-P age difference should be -1000 years. Similar values are found in late glacial and early deglacial sediments studied here (Fig. 3.3). These results do not indicate increased ventilation of the OMZ during the late glacial as suggested for other locations by Duplessy et al. (1989) and Ingram and Kennett (1995). During the Allerod (i.e., the first period of OMZ intensification) the B-P age difference appears to decrease slightly from 860 ± 120 to 610 ± 136 years (Fig. 3.3). This decrease, which we note is not statistically robust, implies better ventilation of the OMZ, but it is inconsistent with geochemical and foraminiferal data that indicate OMZ intensification. Thus, if the reduced age difference observed during the Allerod is real, it must have resulted from some other factor. Rapid, short-term fluctuations in the atmospheric 14C concentration (A14C atm) can influence the reservoir age of surface waters, and thus B-P age differences. Radiocarbon calibration data for the Allerod (Ffughen et al., 2000; Kitagawa and van der Plicht, 2000) show no evidence for a rapid and sustained decline in A14C atm that could explain the slightly lower B-P age difference seen in the Allerod in Core JT96-09. The reservoir age will also be influenced by factors which affect exchange between the atmosphere and ocean (e.g., wind mixing), as well as by changes in oceanic circulation (Bard, 1988). The present-day reservoir age for the Northeast Pacific is approximately 800 years (Robinson and Thomson; 1981; Southon et al., 1990) and this has not changed significantly since the Younger Dryas (Southon et al., 1990). Kovanen and Easterbrook (2002) recently documented a reservoir age of 1100 years during the Allerod. These authors speculated that melting of Cordilleran ice might have supplied relative old C02 to surface waters, thereby increasing their apparent age. If this were the case regional B-P age differences should have decreased when the rapid retreat of the Cordilleran ice sheet commenced at —15.5 kyr B.P. (i.e., 15.0 and 14.0 14C kyr; Clague and James, 2002). Thus, substantial meltwater influx must have occurred in the 87 region prior to the Allerod (13.5 to 12.7 cal. kyr B.P.) and it could therefore not have contributed to an increase in surface-water reservoir age at that time. A more likely possibility is that upwelling brought 14C -depleted subsurface waters to the surface during the Allerod, thus increasing planktonic ages. Upwelling was greatly reduced along the northern and central portions of the California Current System during the Last Glacial Maximum (Sabin and Pisias, 1996) because the Aleutian Low dominated, producing southerly winds and inducing downwelling. Atmospheric circulation in the Northeast Pacific returned to its interglacial mode and upwelling was re-established by -13.0 calendar kyr (Sabin and Pisias, 1996). Thus, the apparent lower B-P age difference observed for the Allerod in Core JT96-09 could have resulted from the upwelling of relatively old water. If this hypothesis is correct primary production should have been enhanced. We present evidence in support of this scenario in the following section. 3.4.3 Changes in productivity? During interglacial periods in the sub-tropical northeast Pacific off northwest Mexico, increased export production led to a decrease in oxygen within the OMZ (Ganeshram et al., 1995). Similar observations have been made off central California (Gardner et al, 1997; Dean and Gardner, 1998) and northern California and Oregon (Lyle et al., 1992; Mix et al., 1999; Kienast et al., 2002). It is not unreasonable, therefore, to suggest that changes in primary and export production at the end of the last glacial affected the intensity of the OMZ along the entire northeastern margin of the Pacific Ocean. But did such changes extend as far north as Vancouver Island? Modern primary production off the west coast of Vancouver Island is influenced by large scale atmospheric circulation. In late spring and summer the North Pacific High drives northerly winds, offshore Ekman transport and upwelling of nutrient-rich water (Huyer, 1983). The strength of these winds and thus upwelling intensity is affected by the strength of 88 the pressure gradient between the North Pacific High and the continental thermal low, such that the larger the gradient the more intense the upwelling (Bakun, 1990). In winter when the North Pacific High shifts southward from ~38°N to ~28°N and is replaced by the Aleutian Low, winds switch direction and upwelling ceases north of ~40°N (Huyer, 1983). Climate modeling and palaeo-evidence suggest that during the Last Glacial Maximum the North Pacific High was positioned further south in summer (COHMAP, 1988; Thunell and Mortyn, 1995; Mortyn et al., 1996; Sabin and Pisias, 1996; Doose et al., 1997), a situation analogous to modern winters. As a result, upwelling along the central and northern portions of the California Current System was greatly diminished and primary productivity was reduced during the Last Glacial Maximum (Dymond et al., 1992; Lyle et al., 1992; Sancetta et al., 1992; Ortiz et al., 1997; Dean and Gardner, 1998; Mix et al., 1999). Off Vancouver Island the burial of marine organic matter was also relatively low during the late glacial, but it began to increase at the start of the Boiling (-14.3 kyr B.P.; Fig. 3.6). The most dramatic increase in marine organic carbon burial (i.e., an apparent six-fold increase relative to late glacial) occurred during the Allerod (-13.5 to 12.6kyr B.P.) and was coincident with the first period of OMZ intensification (Fig. 3.6). A minor increase in organic carbon accumulation is also evident during the second period of OMZ intensification. It could be argued that rather than causing intensification of the OMZ, increased organic carbon burial was the result of better preservation due to lower oxygen concentrations in the bottom water (e.g., Dean et al., 1994; Zheng et al, 2000). In the geological record, laminated sediments commonly have high organic carbon contents and this led to the hypothesis that organic matter preservation is enhanced in anoxic environments (Emerson, 1985), the presumption being that anaerobic bacteria are less efficient at degrading complex organic molecules. However, the sediments in Core JT96-09 are bioturbated, even during periods of inferred OMZ intensification, and thus bottom waters never became anoxic. Therefore, simple enhanced preservation of organic matter resulting from anoxic bottom waters cannot explain increased organic carbon burial during the 89 Marine Organic Carbon (wt.%) 0 1 2 3 4 5 0 0.2 0.4 0.6 0.8 1 1.2 MAR of Marine Organic Carbon (g/cm2/kyr) Fig. 3.6. Marine organic carbon concentration (open squares) and mass accumulation rate (MAR, solid squares) at Station JT96-09 from 8 to 16 kyr B.P. Data are from McKay et al. (submitted). 90 Allerod. It has been suggested recently that high sedimentation rate and low oxygen work in combination to enhance organic matter preservation by controlling the length of time that organic compounds are exposed to oxygen (Hedges and Keil, 1995; Gelinas et al., 2001). We have estimated oxygen exposure times (OETs) for the Holocene, Allerod, Boiling and the Late Glacial at site JT96-09 (Table 3.2). If the oxygen penetration depth remained constant then OETs for the Allerod and Boiling are similar (i.e., 1.7 and 1,3 years, respectively). If the oxygen penetration depth decreased in the Allerod, as the trace metal data imply, but remained the same for the Boiling, computed OETs remain similar (i.e., <1 and 1.3 years). Establishment of a substantial difference in Allerod and Boiling OETs would have required a deeper penetration depth in the Boiling. There is however no evidence to suggest that oxygen penetrated more deeply during the Boiling than during the Holocene. Sedimentary Mn/Al ratios, for example, are low during both periods (0.0035 and 0.0045, respectively) indicating that the near-surface sediments have been continuously suboxic. Furthermore, the higher sedimentation rate during the Boiling would have hindered oxygen penetration. We cannot rule out the possibility that a combination of high sedimentation rate and reduced supply of oxidants from the bottom water played a role in enhancing organic matter accumulation during the Allerod. However, the large increase in the mass accumulation of marine organic matter during the Allerod relative to the Boiling, given that OETs were probably similar, suggests that high export productivity was the dominant factor. Finally, changes in OET should not have affected the burial of barium or opal, yet these palaeo-proxies also imply relatively high export production during the Allerod (Chapter 2). Thus, these observations lead us to conclude that high organic carbon accumulation during the Allerod was primarily the result of increased export production and that the increased settling flux of organic matter caused intensification of the OMZ. Increased production was probably caused by the initiation of upwelling off Vancouver Island as atmospheric and oceanic circulation switched from a glacial to interglacial mode. Upwelling of relatively old 91 Table 3.2. Oxygen exposure times for sediments in Composite Core JT96-09. Time Period Sedimentation Rate Oxygen Penetration 1 OET '•1 (cm/kyr) Depth (cm) (years) Holocene 5 0.2 (-)2 40 (-) Allerod 116 0.2(0.1) 1.7 (< 1) Boiling 150 0.2(1.0) 1.3(7) Late Glacial 48 0.2(1.0) 4.2(20) ' A second set of oxygen penetration depths and resulting OET values are given in brackets. 2 The oxygen penetration depth of 0.2 cm is based on the thickness of the brown "fluff layer observed in Multicore JT96-09mc. The switch from brown to greenish sediments is commonly assumed to be the oxic-suboxic boundary. 5 Oxygen exposure Time (OET) = Oxygen penetration depth divided by sedimentation rate (Hedges and Keil, 1995). 92 waters also may explain why benthic-planktonic age difference decreased slightly during the Allerod. If intensification of the OMZ along the eastern margin of the North Pacific Ocean was the result of increased marine export production at the end of the last glacial there should be evidence from other locations. In general, the last deglacial (13 to 8 kyr B.P.) was a period enhanced marine productivity throughout the North Pacific (e.g., the Northwest Pacific off Kamchatka - Keigwin et al., 1992; the Gulf of Alaska - de Vernal and Pedersen, 1996), as well as within the California Current System off Oregon (Lyle et al., 1992) and California (Gardner et al., 1997; Dean and Gardner, 1998; Mix et al., 1999). However, in the Santa Barbara Basin (California Borderlands region) there is no conclusive evidence that productivity increased during periods of OMZ intensification (Behl and Kennett, 1996), except during the Holocene (Ivanochko, 2001). At present, offshore Ekman transport is lower in the California Borderlands region in comparison to areas just north and south. This difference reflects the more offshore position of the California Current and the unfavorable orientation of the coastline relative to winds (Huyer, 1983). As a result, local upwelling in this region is almost absent in summer when upwelling is strongest along the northern and central portion of the CCS, but local upwelling does occur in winter and early spring (Hickey, 1998). The differences in offshore Ekman transport and upwelling are reflected in phytoplankton pigment concentrations which show that primary productivity within the Southern California Bight and off Baja is generally lower than primary productivity in the northern CCS (Thomas et al., 1994). This may explain why palaeoproductivity records in Santa Barbara Basin do not correlate to similar records for northern California and elsewhere within the CSS. 3.5 Summary 93 Trace metal and benthic foraminifera species data indicate that the OMZ in the northeastern Pacific off Vancouver Island, Canada was more intense (i.e., more oxygen-depleted) between 13.5 to 12.7 calendar kyr B.P. (i.e., the Allerod) and again between 11 and 8 kyr B.P.. The timing of OMZ intensification is similar to that off California and Mexico, suggesting it was a regional phenomenon. Radiocarbon dating of benthic-planktonic foraminiferal pairs indicate that by ~16 kyr B.P. ventilation of the intermediate water mass (920 m water depth) off Vancouver Island was similar to that at present. Furthermore, there is no evidence that ventilation decreased during periods of OMZ intensification. There is however, evidence that export production increased during times of OMZ intensification. During the Allerod organic carbon accumulation increased 6-fold relative to glacial values. The concentrations and accumulation rates of other palaeo-productivity proxies (e.g., bio-barium, opal and alkenone concentrations; Chapter 2) also increased during the Allerod, and to a lesser extend at the Pleistocene-Holocene boundary. These benthic-planktonic age data together with the geochemical indices suggest that increased export production, most probably the direct result of increased primary production in surface waters, was the principal cause of OMZ intensification. Primary productivity was probably stimulated by the re initiation of upwelling as atmospheric and oceanic circulation switched from a glacial to an interglacial mode. 94 3.6 References Anderson, R.Y., Hemphill-Haley, E., Gardner, J.V., 1987. Persistent late Pleistocene-Holocene seasonal upwelling and varves off the coast of California. Quaternary Research 28,307-313. Bard, E., 1988. Correction of accelerator mass spectrometry 14C ages measured in planktonic foraminifera: Paleoceanographic implications. Paleoceanography 3, 635-645. Bard, E., Hamelin, B., Arnold, M., Montaggioni, L., Cabioch, G., Faure, G., Rougerie, F., 1996. Deglacial sea-level record from Tahiti corals and the timing of global meltwater discharge. Nature 382, 241-244. Bakun, A., 1990. Global climate change and intensification of coastal ocean upwelling. Science 247, 198-201. Behl, R.J., Kennett, J.P., 1996. Brief interstadial events in the Santa Barbara Basin, NE Pacific, during the past 60 kyr. Nature 379, 243-246. Cannariato, K.G., Kennett, J.P., Behl, R.J., 1999. Biotic response to Late Quaternary rapid climate switches in Santa Barbara Basin: Ecological and evolutionary implications. Geology 27, 63-66. Cannariato, K.G., Kennett, J.P., 1999. Climatically related millennial-scale fluctuations in strength of California margin oxygen-minimum zone during the past 60 k.y.. Geology 27, 975-978. Clague, J.J., James, T.S., 2002. Historoy and isostatic effects of the last ice sheet in southern British Columbia. Quaternary Science Reviews 21, 71-87. COHMAP Members, 1988. Climatic changes of the last 18,000 years: Observations and model simulations. Science 241,1043-1052. Colodner, D., Sachs, J., Ravizza, G., Turekian, K., Edmond, J., Boyle, E., 1993. The geochemical cycle of rhenium: A reconnaissance. Earth and Planetary Science Letters 117, 205-221. Colodner, D,C, Boyle, E.A., Edmond, J.M., Thomson, J., 1992. Post-depositional mobility of platinum, iridium and rhenium in marine sediments. Nature 358, 402-404. Crusius, J., Calvert, S.E., Pedersen, T.F., Sage, D., 1996. Rhenium and molybdenum enrichments in sediments as indicators of oxic, suboxic and sulphidic conditions of deposition. Earth and Planetary Science Letters 145, 65-78. Crusius, J., Thomson, J., 2000. Comparative behavior of authigenic Re, U, and Mo during reoxidation and subsequent long-term burial in marine sediments. Geochimica et Cosmochimica Acta 64, 2233-2242. De Lange, G.J., Van Os, B., Pruysers, P.A., Middelburg, J.J., Castradori, D., Van Santvoort, P., Muller, P.J., Eggenkamp, H., Prahl, F.G., 1994. Possible early diagenetic alteration of palaeo proxies. In: Zahn, R., Pedersen, T.F., Kaminski, M.A., Labeyrie, L. (Eds.), Carbon 95 Cycling in the Glacial Ocean: Constraints on the Ocean's Role in Global Change, NATO ASI Series, Vol. 117, Springer-Verlag, Berlin, pp. 225-257. Dean, W.E., Gardner, J.V., Anderson, R.Y., 1994. Geochemical evidence for enhanced preservation of organic matter in the oxygen minimum zone of the continental margin of northern California during the late Pleistocene. Paleoceanography 9,47-61. Dean, W.E., Gardner, J.V., 1998. Pleistocene to Holocene contrasts in organic matter production and preservation on the California continental margin. Geological Society of America Bulletin 110, 888-899. Dean, W.E., Gardner, J.V., Piper, D.Z., 1997. Inorganic geochemical indicators of glacial-interglacial changes in productivity and anoxia on the California continental margin. Geochimica et Cosmochimica Acta 61, 4507-4518. Doose, H., Prahl, F.G., Lyle, M.W., 1997. Biomarker temperature estimates for modern and last glacial surface waters of the California Current system between 33° and 42°N. Paleoceanography 12, 615-622. Duplessy, J.-C, Shackleton, N.J., Fairbanks, R.G., Labeyrie, L., Oppo, D., Kallel, N., 1988. Deepwater source variations during the last climate cycle and their impact on the global deepwater circulation. Paleoceanography 3, 343-360. Duplessy, J.-C, Arnold, M., Bard, E., Juillet-Leclerc, J., Kallel, N., Labeyrie, L., 1989. AMS 14C study of transient events and of the ventilation rate of the Pacific Intermediate Water during the last deglaciation. Radiocarbon 31, 493-502. Dymond, J., Suess, E., Lyle, M., 1992. Barium in deep-sea sediment: A geochemical proxy for paleoproductivity. Paleoceanography 7, 163-181. Emerson, S.R., 1985. Organic carbon preservation in marine sediments. In: The carbon cycle and atmospheric C02: Natural variations Archean to Present. Geophysical Monograph Series, v. 32, American Geophysical Union, Washington, D.C., pp. 78-87. Emerson, S.R., Huested, S.S., 1991. Ocean anoxia and the concentration of molybdenum and vanadium in seawater. Marine Chemistry 34, 177-196. Erickson, B.E., Helz, G.R., 2000. Molybdenum(VI) speciation in sulfidic waters: Stability and lability of thiomolybdates. Geochimica et Cosmochimica Acta 64, 1149-1158. Fairbanks, R.G., 1989. A 17,000-year glacio-eustatic sea level record: Influence of glacial melting rates on the Younger Dryas event and deep-ocean circulation. Nature 342, 637-642. Francois, R., 1988. A study of the regulation of the concentrations of some trace metals (Rb, Sr, Zn, Pb, Cu, V, Cr, Ni, Mn and Mo) in Saanich Inlet Sediments, British Columbia, Canada. Marine Geology 83, 285-308. Freeland, H.J., Bychkov, A.S., Whitney, F., Taylor, C, Wong, C.S., Yurasov, G.I., 1998. WOCE section P1W in the Sea of Okhotsk 1. Oceanographic data description. Journal of Geophysical Research 103, 15613-15623. 96 Ganeshram, R.S., 1996. On the glacial-interglacial variability of upwelling, carbon burial and denitrification on the northwestern Mexican continental margin. Ph.D. Thesis, University of British Columbia, 233p. Ganeshram, R.S., Pedersen, T.F., Calvert, S.E., Murray, J.W., 1995. Large changes in oceanic nutrient inventories from glacial to interglacial periods. Nature 376, 755-758. Ganopolski, A., Rahmstorf, S., Petoukhov, V., Claussen, M., 1998. Simulation of modern and glacial climates with a coupled global model of intermediate complexity. Nature 391, 351-356. Gardner, J.V., Hemphill-Haley, E., 1986. Evidence for a stronger oxygen-minimum zone off central California during late Pleistocene to early Holocene. Geology14, 691-694. Gardner, J.V., Dean, W.E., Dartnell, P., 1997. Biogenic sedimentation beneath the California Current System for the past 30 kyr and its paleoceanographic significance. Paleoceanography 12, 207-225. Gelinas, Y., Baldock, J.A., Hedges, J.I., 2001. Organic carbon composition of marine sediments: Effect of oxygen exposure on oil generation potential. Science 294, 145-148. Halpern, D., Smith, R.L., Reed, R.K., 1978. On the California Undercurrent over the continental slope off Oregon. Journal of Geophysical Research 83,1366-1372. Hedges, J.I., Keil, R.G., 1995. Sedimentary organic matter preservation: An assessment and speculative synthesis. Marine Chemistry 49, 81-115. Helz, G.R., Miller, C.V., Charnock, J.M., Mosselmans, J.F.W., Pattrick, R.A.D., Garner, CD., Vaughan, D.J., 1996. Mechanisms of molybdenum removal from the sea and its concentration in black shales: EXAFS evidence. Geochimica et Cosmochimica Acta 60, 3631-3642. Hughen, K.W., Southon, J.R., Lehman, S.J., Overpeck, J.T., 2000. Synchronous radiocarbon and climate shifts during the last deglaciation. Science 290, 1951-1954. Huyer, A., 1983. Coastal upwelling in the California Current System. Progress in Oceanography 12, 259-284. Ingram, B.L.. Kennett, J.P., 1995. Radiocarbon chronology and planktonic-benthic foraminifera C age differences in Santa Barbara Basin sediments, Hole 893A. In: Kennett, J.P., Baldauf, J.G., Lyle, M. (Eds.), Proceedings of the Ocean Drilling Program, Scientific Results, Vol. 146 (Pt. 2), pp. 19-27. Ivanochko, T., 2001. Productivity influences on oxygenation of the Santa Barbara Basin, California, during the Late Quaternary. M.Sc. Thesis, University of British Columbia. Jahnke, R.A., Craven, D.B., McCorkle, D.C, Reimers, C.E., 1997. CaC03 dissolution in California continental margin sediments: The influence of organic matter remineralizaiton. Geochimica et Cosmochimica Acta 61, 3587-3604. Kaiho, K., 1994. Benthic foraminifera dissolved-oxygen index and dissolved-oxygen levels in the modern ocean. Geology 22, 719-722. 97 Karlin, R., Lyle, M., Zahn, R., 1992. Carbonate variations in the northeast Pacific during the late Quaternary. Paleoceanography 7,43-61. Keigwin, L.D., 1998. Glacial-age hydrography of the far northwest Pacific Ocean. Paleoceanography 13, 323-339. Keigwin, L.D., Jones, G.A., 1990. Deglacial climatic oscillations in the Gulf of California. Paleoceanography 5, 1009-1023. Keigwin, L.D., Jones, G.A., Froelich, P.N., 1992. A 15,000 year paleoenvironmental record from Meiji Seamount, far northwestern Pacific. Earth and Planetary Science Letters 111, 425-440. Kennett, J.P., Ingram, B.L., 1995. A 20,000-year record of ocean circulation and climate change from the Santa Barbara basin. Nature 377, 510-514. Kienast, S.S., Calvert, S.E., Pedersen, T.F., 2002. Nitrogen isotope and productivity variations along the North East Pacific margin over the last 120 kyr: Surface and subsurface palaeoceanography. Paleoceanography 17, doi: 10.1029/2001PA000650. Kienast, S.S., McKay, J.L., 2001. Sea surface temperatures in the subarctic Northeast Pacific reflect millennial-scale climate oscillations during the last 16 kyrs. Geophysical Research Letters 28, 1563-1566. Kitagawa, H., van der Plicht, J., 2000. Atmospheric radiocarbon calibration beyond 11900 cal BP from Lake Suigetsu laminated sediments. Radiocarbon 42, 369-380. Koide, M., Hodge, V.F., Yang, J.S., Stallard, M., Goldberg, E.G., Calhoun, J., Bertine, K.K., 1986. Some comparative marine chemistries of rhenium, gold, silver and molybdenum. Applied Geochemistry 1, 705-714. Kono, T., 1996. Modification processes of the intermediate subarctic water in the western North Pacific and its relation to formation of the North Pacific Intermediate Water. Bulletin Hokkaido National Fisheries Resources Institute 60, 145-223. Kovanen, D.J., Easterbrook, D.J., 2002. Paleodeviations of radiocarbon marine reservoir values for the northeast Pacific. Geology 30, 243-246. Lyle, M., Zahn, R., Prahl, F. Dymond, J., Colier, R. Pisias, N., Suess, E., 1992. Paleoproductivity and carbon burial across the California Current: The mulitracers transect, 42°N. Paleoceanography 7, 251-272. Lynn, R.J., Simpson, J.J., 1987. The California Current System: The seasonal variability of its physical characteristics. Journal of Geophysical Research 92,12,947-12,966. Mackas, D.L., Denman, K.L., Bennett, A.F., 1987. Least squares multiple tracer analysis of water mass composition. Journal of Geophysical Research 92, 2907-2918. McCorkle, D.C., Keigwin, L.D., Corliss, B.H., Emerson, S.R., 1990. The influence of microhabitats on the carbon isotopic composition of deep-sea benthic foraminifera. Paleoceanography 5, 161-185. 98 Mikolajewicz, U., Crowley, T.J., Schiller, A., Voss, R., 1997. Modelling teleconnections between the North Atlantic and North Pacific during the Younger Dryas. Nature 387, 384-387. Mix, A.C., Lund, D.C., Pisias, N.G., Boden, P., Bornmalm, L., Lyle, M., Pike, J., 1999. Rapid climate oscillations in the Northeast Pacific during the last deglaciation reflect northern and southern hemisphere sources. In: Clark, P.U., Webb, R.S., Keigwin, L.D. (Eds.), Mechanisms of Global Climate Change at Millennial Time Scales. Geophysical Monograph 112, American Geophysical Union, Washington, D.C., pp. 127-148. Morford, J.L., Russell, A.D., Emerson, S., 2001. Trace metal evidence for changes in the redox environment associated with the transition from terrigenous clay to diatomaceous sediment, Saanich Inlet, BC. Marine Geology 174, 355-369. Mortyn, P.G., Thunell, R.C., Anderson, D.M., Stott, L.D., Le, J., 1996. Sea surface temperature changes in the Southern California Borderlands during the last glacial-interglacial cycle. Paleoceanography 11, 415-430. Mullins, H.T., Thompson, J.B., McDougall, K., Vercoutere, T.L., 1985. Oxygen-minimum zone edge effects: Evidence from the central California coastal upwelling system. Geology 13,481-494. Nameroff, T.J., 1996. Suboxic trace metal geochemistry and paleo-record in continental margin sediments of the Eastern Tropical North Pacific. Ph.D. Thesis, University of Washington. Ortiz, J., Mix, A., Hostetler, S., Kashgarian, M., 1997. The California Current of the last glacial maximum: Reconstruction at 42°N based on multiple proxies. Paleoceanography 12, 191-205. Ostlund, H.G., Stuiver, M., 1980. GEOSECS Pacific radiocarbon. Radiocarbon 22, 25-53. Pilskaln, C.H., Pike, J., 2001. Formation of Holocene sedimentary laminae in the Black Sea and the role of the benthic flocculent layer. Paleoceanography 16, 1-19. Quinterno, P.J., Gardner, J.V., 1987. Benthic foraminifers on the continental shelf and upper slope, Russian River area, northern California. Journal of Foraminiferal Research 17, 132-152. Ravizza, G., Turekian, K.K., Hay, B.J., 1991. The geochemistry of rhenium and osmium in recent sediments from the Black Sea. Geochimica et Cosmochimica Acta 55, 3741-3752. Reed, R.K., Halpern, D., 1976. Observations of the California Undercurrent off Washington and Vancouver Island. Limnology and Oceanography 21, 389-398. Reid, J.L., Jr., 1965. Intermediate waters of the Pacific Ocean. Johns Hopkins Oceanography Study, No. 2, 85pp. Robinson, S.W., Thompson, G., 1981. Radiocarbon corrections for marine shell dates with applications to southern Pacific Northwest Coast prehistory. Syesis 14, 45-57. 99 Rosenthal, Y., Lam, P., Boyle, E.A., Thomson, J., 1995. Authigenic cadmium enrichments in suboxic sediments: precipitation and postdepositional mobility. Earth and Planetary Science Letters 132, 99-111. Sabin, A.L., Pisias, N.G., 1996. Sea surface temperature changes in the northeastern Pacific Ocean during the past 20,000 years and their relationship to climate change in northwestern North America. Quaternary Research 46, 48-61. Sancetta, C, Lyle, M., Heusser, L., Zahn, R., Bradbury, J., 1992. Late-glacial to Holocene changes in winds and upwelling, and seasonal production of the northern California Current system. Quaternary Research 38, 359-370. Sen Gupta, B.K., Machain-Castillo, M.L., 1993. Benthic foraminifera in oxygen-poor habitats. Marine Micropalentology 20,183-201. Southon, J.R., Nelson, D.E., Vogel, J.S., 1990. A record of past ocean-atmosphere radiocarbon differences from the norhteast Pacific. Paleoceanography 5,197-206. Stott, L.D., Berelson, W., Douglas, R., Gorsline, D., 2000a. Increased dissolved oxygen in Pacific intermediate waters due to lower rates of carbon oxidation in sediments. Nature 407, 367-370. Stott, L.D., Neumann, M., Hammond, D., 2000b. Intermediate water ventilation on the northeastern Pacific margin during the late Pleistocene inferred from benthic foraminiferal 513C. Paleoceanography 15, 161-169. Stuiver, M., Reimer, P.J., Bard, E., Beck, J.W., Burr, G.S., Hughen, K.A., Dromer, B., McCormac, G., Van der Plicht, J., Spurk, M., 1998. INTCAL98 radiocarbon age calibration, 24,000-0 cal BP. Radiocarbon 40, 1041-1083. Talley, L.D., 1991. An Okhotsk Sea water anomaly: Implications for ventilation in the North Pacific. Deep Sea Research 38, S171-S190. Talley, L.D., 1993. Distribution and formation of North Pacific Intermediate Water. Journal of Physical Oceanography 23, 517-537. Thomson, J., Higgs, N.C, Croudace, I.W., Colley, S., Hydes, D.J., 1993. Redox zonation of elements as an oxic-post-oxic boundary in deep-sea sediments. Geochimica et Cosmochimica Acta 57, 579-595. Thomson, J., Higgs, N.C, Wilson, T.R.S., Croudace, I.W., de Lange, G.J., van Santvoort, P.J.M., 1995. Redistribution and geochemical behaviour of redox-sensitive elements around SI, the most recent eastern Mediterranean sapropel. Geochimica et Cosmochimica Acta 59, 3487-3501. Thomson, R.E., 1981. Oceanography of the British Columbia Coast. Canadian Special Publication of Fisheries and Aquatic Sciences 56, 291 p. Thunell, R.C., Mortyn, P.G., 1995. Glacial climate instability in the Northeast Pacific Ocean. Nature 376, 504-506. 100 Thunell, R.C., Tappa, E., Anderson, D.M., 1995. Sediment fluxes and varve formation in Santa Barbara Basin, offshore California. Geology 23,1083-1086. van Geen, A., Fairbanks, R.G., Dartnell, P., McGann, M., Gardner, J.V., Kashgarian, M., 1996. Ventilation changes in the northeast Pacific during the last deglaciation. Paleoceanography 11, 519-528. Van Scoy, K.A., Fine, R.A., Ostlund, H.G., 1991a. Two decades of mixing tritium into the North Pacific Ocean. Deep Sea Research 38, SI91-S219. Van Scoy, K.A., Olson, D.B., Fine, R.A., 1991b. Ventilation of North Pacific intermediate waters: The role of the Alaska Gyre. Journal of Geophysical Research 96, 16801-16810. Van Scoy, K.A., Druffel, E.R.M., 1993. Ventilation and transport of thermocline and intermediate watrs in the Northeast Pacific during recent El Ninos. Journal of Geophysical Research 98, 18083-18088. Wong, C.S., Matear, R.J., Freeland, H.J., Whitney, F.A., Bychkov, A.S., 1998. WOCE line P1W in the Sea of Okhotsk 2. CFCs and the formation rate of intermediate water. Journal of Geophysical Research 103, 15,625-15642. Wyrtki, K., 1962. The oxygen minima in relation to ocean circulation. Deep Sea Research 9,11-23. Wyrtki, K., 1967. Circulation and water masses in the eastern equatorial Pacific Ocean. International Journal of Oceanology and Limnology 1, 117-147. You, Y., Suginohara, N., Fukasawa, M., Yasuda, I., Kaneko, L, Yoritaka, FL, Kawamiya, M., 2000. Roles of the Okhotsk Sea and Gulf of Alaska in forming the North Pacific Intermediate Water. Journal of Geophysical Research 105, 3253-3280. Zahn, R., Pedersen, T.F., Bornhold, B.D., Mix, A.C., 1991. Water mass conversion in the glacial Subarctic Pacific (54°N, 148°W): Physical constraints and the benthic-planktonic stable isotope record. Paleoceanography 6, 543-560. Zheng, Y., van Geen, A., Anderson, R.F., Gardner, J.V., Dean, W.E., 2000. Intensification of the northeast Pacific oxygen minimum zone during the Bolling-Allerod warm period. Paleoceanography 15, 528-536. 101 4. Accumulation of Redox-sensitive Trace Metals in Continental Margin Sediments off Western Canada 4.1 Introduction The concentrations of certain metals in sediments are directly or indirectly controlled by redox conditions through either a change in redox state (e.g., Mn, Re and U) and/or speciation (e.g., Mo) which results in their accumulation or loss. Other redox-sensitive metals (e.g., Cd and Ag) have a single redox state but readily react with the reduced forms of other elements such as sulphur. The observation that certain metals are enriched under specific redox conditions, such as Mo in anoxic sediments, has led to their use as palaeoenvironmental proxies (e.g., Piper and Isaacs, 1995; Rosenthal et al., 1995a; Yang et al., 1995; Crusius et al., 1999; Yarincik et al., 2000; Gobeil et al., 2001; Adelson et al., 2001, Pailler et al., 2002). Changes in the intensity of the oxygen minimum zone on the northeastern margin of the Pacific, for example, have been inferred from variations in the sedimentary concentrations of redox-sensitive trace metals (Dean et al., 1997; Zheng, et al., 2000a; Ivanochko, 2001). Trace metals are supplied to sediments via the settling fluxes of lithogenic and biogenic detritus, both of which may fluctuate over time due to changes in sediment provenance and grain-size, and surface water productivity. However, it is the authigenic flux (i.e., the diffusion of metals into the sediment from the overlying water column) that is of most interest because it is controlled by redox conditions in the sediment and overlying bottom water. The degree to which redox-sensitive trace elements accumulate is largely determined by their porewater concentration gradients. These drive diffusion and are influenced by the depth of the redox boundaries. In general, when the suboxic redox boundary (i.e., where oxygen content falls to zero) and anoxic redox boundary (i.e., where sulphate reduction commences) are shallow, diffusion gradients are steeper and the flux from the overlying water column into the sediment is enhanced. Such conditions commonly occur 102 where organic matter flux to the seafloor is high and/or oxygen concentration in the bottom water is low. However, as will be shown in this paper, it is also possible to build-up high metal concentrations even if the suboxic and anoxic redox boundaries are quite deep, as long as they remain relatively stationary for a long period of time. This occurs when sedimentation rate is relatively low, but oxidant demand is high (e.g., Pedersen et al., 1989). A high sedimentation rate can also lead to enhanced trace metal accumulation, but only for those metals that are very rapidly fixed within the sediment (e.g., Mo; Chapter 5). Application of redox-sensitive trace metals as palaeo-proxies requires an understanding of their geochemistry in a wide variety of modern environments. Anoxic environments such as the Black Sea, Cariaco Basin and Saanich Inlet have been well studied (e.g., Francois, 1988; Calvert et al., 1990; Dean et al., 1999; Yarincik et al., 2000; Morford et al., 2001). However, detailed studies of continental margin environments are relatively limited (e.g., Crusius et al., 1996; Dean et al., 1997; Morford and Emerson, 1999; Nameroff et al., 2002). It is therefore a primary goal of this study to provide further insight on the sedimentary geochemistry of redox-sensitive trace metals (i.e., Re, U, Cd and Mo) in modern continental margin settings. The geochemical behaviour of naturally occurring (i.e., non-anthropogenic) Ag in marine sediments is also discussed in this paper. It has been hypothesized that the concentration of sedimentary Ag might be useful as a proxy for the settling flux of diatoms since Ag is incorporated into, and delivered to the sediment with, opal (Friedl et al., 1997). We show here, however, that Ag accumulation in sediments is influenced by more than just the delivery of opal to the seafloor. The study area is located at the northern end of the California Current System (CCS) off Vancouver Island, British Columbia, Canada (48° 54' N, 126° 53' W; Fig. 4.1). This region sits within a transition zone where the eastward-flowing Subarctic and North Pacific currents split into the northward flowing Alaska Current and southward flowing California Current (Thomson, 1981). With the exception of a local and relatively nearshore muddy 103 facies, much of the shelf along the western margin of Vancouver Island is covered by relict deposits (Bornhold and Yorath, 1984). This reflects three factors: i) intense turbulence and scouring associated with winter storm waves, ii) the lack of a large, proximal source of riverine detritus, and iii) efficient particle trapping by the fjords that are common along the western margin of Vancouver Island. The last two factors also limit sedimentation rates on the adjacent slope. The area immediately west of Vancouver Island is typified by high primary productivity (up to 400 gC/mVyr; Antoine et al., 1996) related to seasonal (spring -summer) wind-induced upwelling of nutrients (Hickey, 1998). The region is also characterized by a relatively intense oxygen minimum zone (OMZ) which is most pronounced between 750 and 1300 m with 02 concentrations ranging from 0.3 to 0.5 ml/1. During periods of upwelling this oxygen-poor water is transported onto the shelf (Mackas et al., 1987). High primary productivity and resulting high organic carbon flux to the sediment in combination with low bottom water oxygen concentrations should normally establish strong reducing conditions within the sediment and foster high trace metal concentrations. However, we have found the near-surface sediments on the continental margin off Vancouver Island to be relatively metal-poor. The question is, why? 4.2 Methods Boxcores (be) and multicores (mc) were collected from six sites on the continental margin off the west coast of Vancouver Island (Fig. 4.1) along a transect that extended from the shelf down the slope to a depth of 1750 m. Cores Olmc, 04mc and 06bc were collected from above the OMZ (<750 m), 09mc from within the OMZ, and 02mc and 05bc from below the OMZ (>1300 m). The exact locations and water depths from which the cores were taken are provided in Table 4.1. These cores were sampled every 1 to 2 cm, except in the lower portions of Olmc and 06bc which were sampled every 5 cm and 2 cm, respectively. Samples were freeze-dried and 104 127°W 126°W 125°W Fig. 4.1. The study area off the west coast of Vancouver Island, British Columbia, Canada (Inset) showing the locations where multicores (mc) and box cores (be) were collected. Exact water depths and core descriptions are provided in Table 4.1. 105 Table 4.1. General data for sampling locations and core descriptions. Core Latitude Longitude Water Depth (m) Bottom Water 02 (ml/1) Core Description Olmc 48° 45.95'N 125° 29.57' W 120 2.4 38 cm olive green mud 04mc 49° 00.71'N 126° 49.82'W 407 1.0 9 cm olive green muddy sand underlain by 10 cm gray clay 06bc 48° 58.73' N 126° 52.68'W 720 0.4 18 cm olive green sandy mud underlain by 17 cm gray clay 09mc 48° 54.76' N 126° 53.44'W 920 0.3 40 cm olive green mud 02mc 49° 12.81'N 127° 18.57" W 1340 0.4 1 18 cm olive green mud 05bc 49° 07.91'N 127° 33.12'W 1750 1.2 48 cm olive green mud 1 dissolved O, concentration was measured at 1240 m. 106 in most instances hand ground, except for sandy sediments which were pulverized in a Tema WC disc mill. A sedimentation rate for each core was estimated by assigning an age of 0 years to the core top and obtaining an age estimate from near the core base; linear sedimentation was assumed between the two assigned ages. Where possible 14C dating was performed by accelerator mass spectrometry (AMS) on a mixed assemblage of planktonic foraminifera (i.e., Cores 09mc and 05bc). However, this was not always feasible given the scarcity of carbonate micro fossils in these deposits. In Core Olmc no foraminifera were available so bulk organic carbon was dated by AMS. Such materials typically yield older than expected ages due to the presence of "old" organic matter. To validate the use of bulk organic carbon ages, 14C measurements of coexisting benthic foraminifera and organic carbon were obtained for one sample taken from 09mc. For Cores 06bc and 04mc no radiocarbon measurements were made and sedimentation rates were calculated by assuming that the contact between the Holocene green mud and deglacial gray clay, which is observed in both cores, is the same age as the mud-clay contact observed in a nearby piston core. The sedimentation rate for Core 02mc was also estimated by dating sediments in a nearby core. The depth of bioturbation was estimated using excess 210Pb data. The method of Eakin and Morrison (1978), which determines the 210Pb concentration by measuring the content of its 210Po grand-daughter via alpha counting, was employed. Total carbon (Ctot), nitrogen (Ntot) and sulphur (Stot) were measured using a Carlo-Erba NA-1500 elemental analyzer. Precision and accuracy were determined for two National Research Council of Canada (NRC) standards (PACS-1 and MESS-1). The relative standard deviation (RSD, la) was 3%, 5% and 6% for C, N and S, respectively and the accuracy was within 7% of the values recommended for all three elements. Carbonate carbon (Ccarb) was determined by coulometry. Repeat analyses (n=141) of a CaC03 standard yielded a mean value of 11.93% and a RSD of 1%. The percent organic carbon (Corg) was calculated by difference (C = Ctot - Ccarb) and has an aggregate error (as RSD) of -4%. 107 Biogenic silica was measured using the NajCOj dissolution method of Mortlock and Froelich (1989). The RSD determined for two in-house standards SBS and JV5, which contain approximately 28% and 11% opal respectively, was 4%. However, the precision of the method is much poorer at opal concentrations <10% (R. Ganeshram, pers. comm.), possibly due to the dissolution of volcanic glass and clay minerals. Major and minor element contents were determined by X-ray fluorescence spectrometry (XRF) following Calvert (1990) and have RSDs of 5% and 15%, respectively. Total Ba (Batot) and Al data were used to calculate the percent biogenic barium (Babi0) via the Dymond et al. (1992) equation: Babl0 = Batot-(AlxBa/AlHJ (4.1) An average Ba/Al lithogenic ratio (Ba/Allith) of 0.0027 was employed. This value is significantly lower than the average Ba/Aliith ratio of crustal rocks (0.0075; Dymond et al., 1992); however, this ratio can vary greatly depending on the composition of the sediment source (e.g., 0.005 to 0.010; Taylor and McLennan, 1985). The geology of Vancouver Island is complex and includes a mixture of metamorphic, igneous and sedimentary rock types (Yorath and Nasmith, 1995); thus it is not possible to chose a Ba/Aliith ratio representative of any particular rock type. Rather, the BaZAl^ ratio used here (i.e., 0.0027) was determined from an exponential regression of the Ba/Al ratios of surface sediments versus water depth following the method Klump et al. (2000). This method assumes that the fraction of biogenic Ba increases seaward and that close to land (i.e., ~0 m water depth) all of the Ba is terrigenous in origin. The low Ba/Alijth estimated using this regression method was confirmed by chemically extracting the bio-barium from six samples using a 2M solution of NH4CI (Schenau et al., 2001) and measuring the barium content of the residue (i.e., the lithogenic Ba) by XRF. This chemical dissolution method yields a slightly higher average Ba/Allith ratio of 0.0033 (range 0.0026 to 0.0038; Table 4.2). 108 Table 4.2. Results of Biogenic Barium Dissolution Tests. Core / Sample Age (Cal. kyr) Total Ba 1 (ppm) Ba,ieh2 (ppm) Total Al (ppm) Ba/Al,lth Ba^1 (ppm) Ha*,4 (ppm) JT96-01mc 0-1 cm 0.01 411 527 140000 0.0038 bdl 61 JT96-09mc 5-6 cm 1.16 624 529 147000 0.0036 96 226 JT96-09pc 241-242 cm 14.00 464 497 159200 0.0031 bdl 34 JT96-09pc 341-342 cm 15.38 573 622 162500 0.0038 bdl 135 JT96-02mc 5-6 cm 0.23 759 505 167000 0.0030 255 341 JT96-05bc 5-6 cm 1.20 1057 402 152000 0.0026 655 677 Average Ba/Allith ratio 0.0033 1 Total Ba concentration in untreated samples. 2 Lithogenic Ba measured in samples that were treated with NH4C1 to remove biogenic Ba. 3 Biogenic Ba calculated by difference (Batot - Balith). 4 Biogenic Ba calculated using Equation 4.1 and a Ba/Aliith ratio of 0.0027. 109 Stable isotope data were obtained by continuous-flow mass spectrometry using a Fisons NA-1500 elemental analyzer in line with a VG Prism mass spectrometer. Samples for carbon isotopic analysis of organic matter (513Corg) were pretreated with 10% HC1 to remove carbonate and then dried at 50°C overnight. These samples were not washed with distilled water prior to drying. Nitrogen isotope results (815N) were obtained for untreated bulk sediment samples. Data are reported in the standard 5-notation relative to VPDB for carbon and atmospheric N2 for nitrogen. The reproducibility for an in-house isotopic standard was ± 0.1 %o for carbon and ± 0.2%o for nitrogen. Trace metal concentrations (i.e., Re, Ag, Cd, Mo and U) were measured by isotope-dilution inductively-coupled plasma mass spectrometry. Sample preparation involved adding known amounts of isotopically-enriched spike solutions to ~20 mg of powdered sediment. Samples were then microwave digested in a mixture of concentrated HN03, HC1 and HF. The digests were evaporated on a hotplate overnight and then re-digested in 5N HC1. An aliquot was taken for Mo and U analysis. The remaining sample was run through an anion exchange column (Dowex 1-X8 resin) to remove Zr and Nb, as these form compounds (ZrO, ZrOH and NbO) that interfere with the analysis of Ag, and Mo as it forms MoO" that interferes with Cd analysis. A detailed description of the sample preparation and analysis can be found Ivanochko (2001). To check precision, a University of British Columbia sediment standard (SNB) was analyzed with each batch of samples. The resulting RSD (la) for Re, U, Mo, Cd and Ag analyses were 11%, 9%, 7%, 10% and 8%, respectively. Accuracy, assessed by measuring the concentrations of these metals in the NRC sediment standard MESS-1, was 8% or better for Re, Mo and U, and -14% for Cd. The accuracy for Ag could not be evaluated as there is no accepted Ag value for MESS-1 at present. Trace metal data are present as concentrations rather than metal/Al ratios because the concentration data yield a finer resolution and because the sediments are relatively homogeneous; thus downcore metal/Al and concentration profiles do not differ significantly. Also, no attempt 110 has been made to calculate the authigenic fraction because the lithogenic concentrations of these metals are poorly constrained. 4.3 Results Core descriptions are provided in Table 4.1. The olive green muds, sandy muds and muddy sands observed in these near-surface sediment cores are typical of Holocene deposits off Vancouver Island (Bornhold and Barrie, 1991). The gray clay found at the base of Cores 04mc and 06bc is most probably Pleistocene in age and of glaciomarine origin (Bornhold and Yorath, 1984; Bornhold and Barrie, 1991). Radiocarbon data and estimated sedimentation rates are given in Table 4.3. Holocene sedimentation rate for slope Cores 09mc and 05bc is -5 cm/kyr. The radiocarbon date obtained for bulk organic carbon from a depth of 39 cm in 09mc (9920 ± 40 14C yrs) is essentially identical to the age determined by AMS dating of benthic foraminifera (9830 ± 110 14C yrs). This result suggests that bulk organic matter can be used for dating these near-surface sediments. However, the use of bulk organic carbon dates is only valid for Holocene sediments that contain little terrestrial organic matter and not for the Late Pleistocene sediments that contain up to 70 % old terrestrial organic matter (Chapter 2). For the inner-shelf Core Olmc radiocarbon dating bulk organic carbon yields an age of 923 calendar yrs and suggests a sedimentation rate of -40 cm/kyr. No radiocarbon data are available for Multicore 02mc; however, 14C data for nearby Piston core Tul96-05pc (see Table A6 in the Appendix) suggest that at a water depth of -1300 m the Holocene sedimentation rate is -23 cm/kyr. This is admittedly only a rough estimate as sedimentation rates could vary substantially between the two locations. No radiocarbon dates were obtained for Cores 04mc or 06bc, but the contact between the olive green mud and gray clay which is observed in both cores has been dated as -11 calendar kyr B.P. in Piston Core JT96-09 (Chapter 2). Using 111 Table 4.3. Radiocarbon data and estimated sedimentation rates. Core Water Depth Material Sample (m) Dated Depth (cm) ,4CAge Calendar Age3 Sedimentation (yrs) (yrs) Rate (cm/kyr) Olmc 04mc 06bc 09mc 02mc 05bc 120 407 720 920 1340 1750 bulk organic matter 36.5 1780 ±30 923 planktonic forams 1 47.5 benthic forams 39.0 bulk organic matter 39.0 planktonic forams 2 234.5 planktonic forams 47.0 9760 ± 70 9830± 110 9920 ± 40 9800 ± 90 9960 ± 50 10026 10127 10290 39.5 0.8 4 1.64 4.7 23.2 4.6 1 planktonic forams were obtained from piston core JT96-09 that was collected from the same location. 2 determined using radiocarbon data obtained from nearby piston core Tul96-05. 314C data were converted to calendar years using Calib 4.3 (Stuiver et al., 1998) and assuming a reservoir age of 800 yrs. 4 estimated assuming that the contact between the Holocene muds and gray clay is at 11.0 calendar kyrs. 112 this date and assuming that the core top is 0 years yields sedimentation rates of 0.8 and 1.6 cm/kyr for Cores 04mc and 06bc, respectively. Total 210Pb concentrations in surface sediments range from 24.5 dpm/g in Core Olmc to 58.8 dpm/g in Core 02mc. Supported 210Pb produced in situ by the decay of 226Ra ranges from 1.8 to 2.8 dpm/g, similar to that in sediments from the Washington continental margin to the south (Carpenter et al., 1981). Unsupported or excess 210Pb (i.e., 210Pb scavenged from the water column) occurs to depths as deep as 10 cm in 02mc, 19 cm in 09mc, and 30 cm in Olmc (Fig. 4.2). Such depths of penetration are well below those expected given the estimated sedimentation rates (5 to 40 cm/kyr) and short half-life of 210Pb (22.3 yrs) implying that these sediments are deeply bioturbated. By mixing younger and more radiocarbon-rich sediment downwards, such bioturbation will decrease the "true age" of subsurface deposits, resulting in overestimates ofthe "true" rates of sedimentation. No corrections for this effect have been applied to the sedimentation rates listed in Table 4.3. Organic carbon, carbonate and opal concentrations in these near-surface cores are provided in Table 4.4 and their downcore profiles are shown in Figure 4.3. Organic carbon (Corg) contents are highest in the surface sediments of 02mc and 09mc (3.4 and 3.1 wt.%, respectively) and decrease with depth (Fig. 4.3a). Lower values (1.6 to 2.1 wt.%) are seen in Olmc, 05bc and 06bc and, with the exception of 06bc, decrease only slightly with depth (Fig. 4.3a). The abrupt decrease in Corg seen at -20 cm depth in 06bc occurs at the contact between the Holocene mud and Pleistocene gray clay. The lowest Corg content (0.5 to 0.6 wt.%) is found in the sandy sediments of 04mc. The mass accumulation rate of Corg, calculated as the product of the linear sedimentation rate, the dry bulk density and the organic carbon concentration, is lowest in 09mc and 05bc (< 0.1 g/cm2/kyr) and up to an order of magnitude higher in Olmc and 02mc (0.6 to 0.3 g/cm2/kyr, respectively). The CaC03 content in all cores is low, ranging from <1.2 wt.% in surface sediments to 3 % in the lower portions of 09mc, 05bc and 06bc (Fig. 4.3b). Opal concentration ranges from < 3 to 11 wt.% in surface sediments and exhibits little change downcore (Fig. 4.3c). 113 Fig. 4.2. Concentration of "excess" 210Pb in near-surface sediment cores Olmc, 02bc and 09mc. 114 Organic Carbon (wt.%) Carbonate (wt.%) 0 1 2 3 0 1 2 3 Opal (wt. %) 50 Fig. 4.3. Concentrations of a) organic carbon, b) carbonate, and c) opal in near-surface sediment cores. The precision (as RSD) on these data are 4 %, 1 % and 4 % respectively. 115 The carbon isotopic composition of organic matter (813Corg) in surface sediments ranges from -21.8 %o in Olmc to -20.5 %o in 04mc (Table 4.4), and changes little downcore except in 04mc and 06bc (Fig. 4.4a). The Pleistocene gray clay in the lower portion of these cores is characterized by relatively low 513Cor8 values (-23.7 to -23.3 %o; Table 4.4). The nitrogen isotopic composition of bulk surface sediments (515N) ranges from +4.8 %o in shelf Core Olmc to +6.8 %o in 06bc (Table 4.4), and is invariant in the upper 20 cm of all of the cores (Fig. 4.4b). Below this depth there is a minor increase in 515N in 05bc and 09mc (up to +7.9 %o), and a decrease in 06bc (down to +5.8 %o) at the boundary between the Holocene mud and underlying Pleistocene gray clay. Manganese concentrations ranges from 293 to 791 ug/g (Table 4.5). In surface sediments Mn/Al ratios range from 0.0026 to 0.0050, substantially less than the ratio of average shale (0.0106; Turekian and Wedepohl, 1961), and decrease with depth in the upper 2 to 5 cm of all cores (Fig. 4.5a). The Pleistocene gray clay observed in the lower portion of Cores 04mc and 06bc exhibits relatively high Mn/Al ratios, similar to surface sediments in Cores 02mc and 05bc. Iodine concentrations in surface sediments range from 104 to 968 ug/g. The I/Corg ratios of surface sediments range from 120 x 10"4 in Olmc to 500 x 10"4 in 05bc (Table 4.4), and decreases gradually with depth in all cores (Fig. 4.5b). Total Ba content ranges from 291 pg/g in 04mc to 1057 ug/g in 05bc (Table 4.5). The ratio of Bat0/Al ranges from 0.0025 to 0.0070 and tends to increase with increasing water depth (Fig. 4.5c). The exception is Core 04mc that has relatively low Batot/Al ratios, most probably reflecting the sandy nature of these sediments. Biogenic Ba concentrations, calculated using Equation 4.1 and assuming a Ba/Allith ratio of 0.0027, range from ~ 9 ug/g in 04mc to 661 ug/g in 05bc (Table 4.4) and increases with water depth. Rhenium concentrations range from 2 to 65 ng/g (Table 4.5). Concentrations are relatively low (2 to 4 ng/g) in the upper 10 cm of all cores (Fig. 4.6a). There is no change in Re concentration with depth in Cores Olmc, 04mc and 06bc, but values rise to 65 ng/g in the 116 Fig. 4.4. a) The C-isotopic composition of organic matter (S^Corg) and b) the N-isotopic composition of the bulk sediment (S^N) in near-surface sediment cores. The errors on these data are ±0.1 %o and ±0.2 %o for C-and N-isotopic ratios respectively. 117 p Mn/Al 0.002 0.003 0.004 0.005 0 I/Core(xl0-4) 10 20 30 40 50 02mc cr* i pv/3 \ °04mc 4 06bc Olmc \j09mc 05bc ^ -org 100 200 300 ; i M - i . 1 H yy 1 02mc ' 04mc \ • 1 fl • 1 i A • i [ A • 11 i / • 06bc lr ' Olmc 6 1 09mc 1 1 1 05bc 1 \ anoxic \ OX1C Ba/Al 0.004 0.006 i i i j. i i 0.008 04mc (407m) Fig. 4.5. a) Mn/Al, b) I/Corg, and c) Ba/Al ratios in near-surface sediment cores. Note, Ba/Al ratios have not been corrected for the lithogenic contribution which is estimated to be -0.0027. The precision (as RSD) is -16 % for all of the ratios shown. 118 lower portion of 09mc which was collected from the OMZ (Fig. 4.6a). Cores 02mc and 05bc, both of which were collected from below the OMZ, also exhibit Re enrichment at depth (Fig. 4.6a). Uranium content ranges from <2 to 5.8 ug/g (Table 4.5). Concentrations are low (< 2 ug/g) in the upper 10 to 15 cm of all cores, and are slightly enriched at depth except in 04mc (Fig. 4.6b). The enrichment is greatest in OMZ Core 09mc (up to 5.8 ug/g). Comparison of the Re and U profiles (Figs. 4.6a and b, respectively) suggest that the enrichment of U occurs slightly deeper than that of Re. Cadmium concentrations are generally low throughout these cores (< 1.0 ug/g; Table 4.5). Surface sediments contain <0.1 to 0.2 pg/g Cd and there is only a slight increase with depth in Cores 02mc, 05bc and 09mc (0.4 to 1.0 ug/g; Fig. 4.7a). These are the same cores that exhibit Re enrichments at depth. The highest Cd concentration occurs in OMZ Core 09mc. Molybdenum contents range from 0.4 to 1.5 ug/g (Table 4.5). Surface sediments contain from 0.4 to 0.9 pg/g Mo (Fig. 4.7b). The concentration increases slightly with depth, particularly in shelf Core Olmc, but there is no evidence of substantial enrichments above the typical lithogenic background value (Fig. 4.7b). Silver concentrations range from <100 to 357 ng/g in surface sediments (Table 4.5). Concentrations vary little downcore in Olmc, 04mc, 06bc and 09mc, but increase by over 200 ng/g in Cores 02mc and 05bc (up to 584 pg/g; Fig. 4.7c). In surface sediments there is a trend of increasing Ag content with increasing water depth, except in the sandy sediments of Cores 04mc and 06bc. 4.4 Discussion 4.4.1 C and N isotope data 119 Re (ng/g) U (ug/g) Fig. 4.6. a) Re and b) U concentrations in the near-surface sediment cores. Typical lithogenic concentrations are shown by the dashed lines. The precison (as RSD) of Re and U measurements is 11 % and 9 % respectively 120 Ag (ng/g) 04mc (407m) 600 05bc (1750m) Fig. 4.7. The concentrations of a) Cd, b) Mo, and c) Ag in the near-surface sediment cores. Typical lithogenic concentrations are shown by the dashed lines. Water depths from which the cores were collected are shown in Figure 7c. Note that, in general, surface Ag concentrations increase as water depth increases. Thus, Core 05bc, from a depth of 1750 m, is characterized by the highest Ag concentrations. The precision (as RSD) for Mo, Cd and Ag measurements is 7 %, 10 % and 8 % respectively. 121 Table 4.4. General geochemical data for Vancouver Island Margin multicores (mc) and box cores (be). Core Olmc Depth Age 8I5N S13C c Opal Ba^1 I/O* (cm) (Cal. kyrs) (%« vs air) (%o vs VPDB) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (ug/g) (x 10"4) 0-1 0.5 0.01 4.8 -21.8 2.2 0.2 0.3 0.7 2.1 8.6 10.9 61 124 1-2 1.5 0.04 2.0 0.2 0.2 0.6 1.9 8.7 2-3 2.5 0.06 2.0 0.2 0.2 0.6 1.9 8.7 3-4 3.5 0.09 1.9 0.2 0.3 0.7 1.8 8.6 4-5 4.5 0.11 1.9 0.2 0.3 0.6 1.9 8.7 5-6 5.5 0.14 4.8 -22.2 2.0 0.2 0.3 0.6 1.9 9.0 9.2 80 89 6-8 7.0 0.18 1.9 0.2 0.3 0.6 1.8 8.7 8-10 9.0 0.23 1.9 0.2 0.3 0.7 1.8 8.7 10-12 11.0 0.28 4.8 -22.3 1.9 0.2 0.3 0.7 1.8 8.9 9.7 85 89 12-14 13.0 0.33 1.9 0.2 0.3 0.7 1.8 8.8 14-16 15.0 0.38 5.3 -22.3 1.9 0.2 0.3 0.7 1.8 9.0 10.2 65 72 16-18 17.0 0.43 1.8 0.2 0.3 0.6 1.8 9.1 18-20 19.0 0.48 4.8 -22.3 1.9 0.2 0.3 0.7 1.8 9.2 40 78 20-25 22.5 0.57 5.2 -22.3 1.8 0.2 0.4 0.7 1.7 9.2 10.4 73 73 25-30 27.5 0.70 4.9 -22.3 1.7 0.2 0.5 0.8 1.6 9.3 29 62 30-35 32.5 0.82 5.5 -22.1 1.6 0.2 0.5 0.8 1.5 9.2 8.2 72 56 35-38 36.5 0.92 5.2 -22.1 1.7 0.2 0.6 0.8 1.6 10.0 61 54 Core 04mc Depth Age 5"N 513C0, c„ N„ s,„, C0rg C„/N Opal Bab„' I/O, (cm) (Cal. kyrs) (% o vs air) (%o vs VPDB) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (ug/g) (x 10^) 0-1 0.5 -20.5 0.5 0.1 <0.1 0.3 0.5 8.2 2.5 9 203 1-2 1.5 0.5 0.1 <0.1 0.2 0.5 8.4 2-3 2.5 0.5 0.1 <0.1 0.2 0.5 7.6 3-4 3.5 0.5 0.1 <0.1 0.2 0.5 7.6 4-5 4.5 0.5 0.1 <0.1 0.2 0.5 7.9 5-6 5.5 -20.5 0.5 0.1 <0.1 0.2 0.5 7.9 3.2 9 168 6-7 6.5 0.6 0.1 <0.1 0.2 0.5 8.2 7-8 7.5 0.6 0.1 <0.1 0.2 0.5 7.7 8-9 8.5 0.6 0.1 <0.1 0.3 0.6 8.0 9-10 9.5 0.6 0.1 0.1 0.3 0.6 8.5 10-11 10.5 -22.4 0.6 0.1 0.2 0.4 0.5 8.8 3.2 59 88 11-12 11.5 0.6 <0.1 0.4 1.2 0.5 10.2 12-13 12.5 0.5 0.1 0.3 0.4 0.5 9.5 13-14 13.5 0.6 0.1 0.3 0.5 0.5 9.3 14-16 15.0 -23.6 0.6 <0.1 0.3 0.9 0.5 9.5 2.9 57 16-18 17.0 0.6 0.1 0.1 0.4 0.5 8.9 Core 06bc Depth Age 8,SN 8,3C„ N„, s„ Ccarb C0rg Corg/N Opal vc„ (cm) (Cal. kyrs) (% o vs air) (%o vs VPDB) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (ug/g) (x 10J) 0-0.5 0.3 6.8 -21.2 2.1 0.2 0.2 1.2 2.0 8.3 6.3 164 242 0.5-1 0.8 2.1 0.2 0.1 1.1 1.9 8.2 1-2 1.5 1.9 0.2 0.1 0.7 1.8 8.4 2-3 2.5 2.0 0.2 0.1 0.6 1.9 8.3 3-4 3.5 2.0 0.2 0.1 0.5 1.9 8.3 4-5 4.5 6.9 -21.0 1.8 0.2 0.1 0.5 1.8 8.4 5.1 150 176 5-7 6.0 1.8 0.2 0.1 0.4 1.7 8.4 7-9 8.0 1.7 0.2 0.1 0.4 1.6 8.4 9-11 10.0 6.6 -21.2 1.5 0.2 0.1 0.4 1.5 8.5 4.6 125 173 11-13 12.0 1.6 0.2 0.1 0.4 1.5 8.4 13-15 14.0 6.5 -20.9 1.5 0.2 0.1 0.4 1.5 8.5 4.5 146 146 17-19 18.0 1.4 0.2 0.1 0.5 1.3 8.6 21-23 22.0 6.6 -23.7 0.8 0.1 0.3 1.7 0.6 9.5 4.3 205 42 25-27 26.0 6.3 -23.3 0.9 0.1 0.3 3.0 0.6 10.6 3.9 192 37 29-31 30.0 5.8 -23.5 0.9 0.1 0.3 2.7 0.5 9.8 33-35 34.0 5.9 -23.5 0.8 0.1 0.4 2.2 0.6 9.7 5.2 142 30 continued on next page 122 Table 4.4 (continued) Core 09mc Depth Age 5'*N 5,3C0„ c„ N,„, Slot Ccarb Corg CJN Opal Ba^1 I/Co, (cm) (Cal. kyrs) (%o vs air) (%o vs VPDB) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (ug/g) (x 104) 0-1 0.5 0.09 6.7 -21.3 3.2 0.4 0.4 1.0 3.1 8.7 7.0 284 301 1-2 1.5 0.27 3.2 0.4 0.4 0.9 3.0 8.4 2-3 2.5 0.45 3.2 0.4 0.3 0.8 3.1 8.3 3-4 3.5 0.63 3.1 0.4 0.3 0.6 3.0 8.4 4-5 4.5 0.81 3.1 0.4 0.3 0.5 3.0 8.4 5-6 5.5 0.99 7.0 -21.2 3.0 0.3 0.2 0.6 3.0 8.5 6.7 256 244 6-8 7.0 1.26 3.0 0.3 0.2 0.4 3.0 8.5 8-10 9.0 1.63 3.0 0.3 0.2 0.3 3.0 8.5 10-12 11.0 1.99 6.7 -21.2 2.8 0.3 0.3 0.2 2.8 8.9 6.8 233 220 12-14 13.0 2.35 2.8 0.3 0.2 0.3 2.8 8.6 14-16 15.0 2.71 6.8 -21.2 2.7 0.3 0.2 0.3 2.7 8.7 5.8 260 200 16-18 17.0 3.07 2.8 0.3 0.2 0.3 2.7 8.6 18-20 19.0 3.43 2.1 0.2 0.2 0.3 2.1 8.9 20-22 21.0 3.79 6.5 -21.4 2.0 0.2 0.2 0.4 1.9 9.4 4.3 247 129 22-24 23.0 4.15 1.8 0.2 0.1 0.7 1.7 9.1 24-26 25.0 4.52 7.7 -21.7 1.8 0.2 0.2 1.2 1.6 9.3 4.4 205 69 26-28 27.0 4.88 1.8 0.2 0.2 1.9 1.6 9.0 28-30 29.0 5.24 1.9 0.2 0.2 2.1 1.6 9.0 30-32 31.0 5.60 7.4 -21.6 1.9 0.2 0.2 2.3 1.6 9.2 4.5 213 70 32-34 33.0 5.96 2.0 0.2 0.2 2.4 1.7 9.1 34-36 35.0 6.32 7.9 -21.6 2.0 0.2 0.2 2.4 1.7 9.4 4.9 198 65 36-38 37.0 6.68 2.1 0.2 0.2 2.4 1.8 9.1 38-40 39.0 7.05 7.5 -21.9 2.1 0.2 0.3 2.6 1.8 9.2 4.8 214 63 Core 02mc Depth Age 5>5N 813Q, c„ N„ s» Ccarb r ^org C0I£/N Opal Ba^1 i/c„ (cm) (Cal. kyrs) (%o vs air) (%» vs VPDB) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (ng/g) (x \0*) 0-1 0.5 0.02 6.6 -21.5 3.5 0.4 0.4 0.8 3.4 8.5 8.8 378 285 1-2 1.5 0.06 3.3 0.4 0.4 0.8 3.2 8.4 2-3 2.5 0.10 3.3 0.4 0.3 0.6 3.3 8.4 3-4 3.5 0.14 3.3 0.4 0.3 0.5 3.2 8.3 4-5 4.5 0.18 3.2 0.4 0.3 0.4 3.2 8.2 5-6 5.5 0.23 6.2 -21.1 3.3 0.4 0.3 0.3 3.2 8.5 8.7 341 208 6-8 7.0 0.29 3.1 0.4 0.3 0.3 3.1 8.5 8-10 9.0 0.37 3.0 0.4 0.3 0.3 3.0 8.4 10-12 11.0 0.45 6.8 -21.1 2.8 0.3 0.2 0.2 2.8 8.8 8.6 368 172 12-14 13.0 0.53 2.9 0.3 0.3 0.3 2.8 8.6 14-16 15.0 0.62 6.7 -21.1 2.6 0.3 0.3 0.3 2.5 8.8 9.9 371 138 16-18 17.0 0.70 2.5 0.3 0.3 0.4 2.5 8.7 18-20 19.0 0.78 6.7 -21.1 2.5 0.3 0.3 0.4 2.4 8.8 8.1 345 147 Core 05bc Depth Age 5"N 813C N„ Ccarb r Opal Ba^1 i/c„, (cm) (Cal. kyrs) (%» vs air) (%o vs VPDB) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (^g/g) (x 10-") 0-1 0.5 0.11 6.3 -21.3 1.7 0.2 0.2 0.6 1.6 8.3 6.5 560 329 1-2 1.5 0.33 2.0 0.2 0.2 0.6 1.9 8.2 556 259 2-3 2.5 0.55 1.9 0.2 0.2 0.5 1.9 8.2 559 250 3-4 3.5 0.77 1.8 0.2 0.2 0.5 1.7 8.2 4-5 4.5 0.99 1.7 0.2 0.2 0.3 1.7 8.3 5-6 5.5 1.20 6.6 -21.0 1.8 0.2 0.2 0.3 1.8 8.4 7.1 677 250 6-7 6.5 1.42 1.9 0.2 0.2 0.3 1.8 8.3 7-8 7.5 1.64 1.8 0.2 0.2 0.3 1.8 8.3 8-9 8.5 1.86 1.9 0.2 0.2 0.3 1.8 8.3 9-10 9.5 2.08 1.8 0.2 0.2 0.3 1.7 8.4 10-12 11.0 2.41 6.7 -21.1 1.8 0.2 0.2 0.4 1.8 8.4 6.7 664 213 12-14 13.0 2.85 1.8 0.2 0.2 0.3 1.7 8.4 14-16 15.0 3.28 6.5 -21.1 1.7 0.2 0.2 0.2 1.7 8.4 6.5 665 178 continued on next page 123 Table 4.4 (continued) Core05bc Depth Age 815N 8I3C0, CM N„ S„ C, Cor, CJN Opal Baj VCn (cm) (Cal. kyrs) (%o vs air) (%, vs VPDB) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (ng/g) (x if/) 16-18 17.0 3.72 1.7 0.2 0.1 0.2 1.7 8.6 18-20 19.0 4.16 1.7 0.2 0.1 0.3 1.7 8.5 20-22 21.0 4.60 6.9 -21.2 1.6 0.2 0.1 0.4 1.6 8.7 5.9 584 150 22-24 23.0 5.04 1.7 0.2 0.1 0.3 1.7 8.5 24-26 25.0 5.47 6.8 -21.3 1.5 0.2 0.1 0.3 1.4 8.7 5.5 552 142 26-28 27.0 5.91 1.6 0.2 0.1 0.3 1.6 8.7 28-30 29.0 6.35 1.3 0.1 0.1 0.4 1.3 8.8 30-32 31.0 6.79 7.0 -21.5 1.4 0.2 0.1 0.6 1.3 8.8 5.4 513 130 32-34 33.0 7.22 1.5 0.2 0.1 0.8 1.4 8.9 34-36 35.0 7.66 7.2 -21.6 1.5 0.2 0.1 1.1 1.4 8.7 4.6 496 112 36-38 37.0 8.10 1.6 0.2 0.1 1.3 1.4 8.9 38-40 39.0 8.54 1.6 0.2 0.1 2.4 1.4 8.9 40-42 41.0 8.98 7.9 -21.7 1.7 0.2 0.1 2.7 1.4 8.8 5.1 456 94 42-44 43.0 9.41 1.8 0.2 0.1 2.9 1.4 8.9 44-46 45.0 9.85 7.8 -21.9 1.7 0.1 0.2 3.1 1.3 9.0 4.8 385 88 46-48 47.0 10.29 1.6 0.1 0.2 2.6 1.3 9.6 1 Biogenic barium (Ba,*,) values were determined using the Dymond et al. (1992) equation Ba^ = Ba,„ - (Al x Ba/Al,ilh). 124 Table 4.5. Major, minor and trace element data for Vancouver Island Margin multicores (mc) and box cores (be). Core Olmc Al Mn Ba,,,, Zr Ag Cd Re Mo U (wt.%) (ug/g) (ng/g) (ug/g) (ng/g) (ng/g) (ng/g) (ng/g) (ng/g) 0-1 14.0 532 411 158 84 0.2 2 0.6 I.I 1-2 81 0.2 3 0.7 1.3 2-3 94 0.2 3 0.7 1.3 3-4 89 0.2 3 0.7 1.4 4-5 97 0.2 4 0.7 1.6 5-6 14.0 449 430 169 87 0.2 3 0.7 1.3 6-8 94 0.3 2 0.8 1.3 8-10 86 0.2 3 1.0 1.4 10-12 13.0 454 410 166 85 0.2 3 1.0 1.4 12-14 87 0.2 3 1.0 1.3 14-16 14.9 454 438 164 79 0.2 4 1.2 1.3 16-18 96 0.2 4 1.0 1.6 18-20 15.0 418 415 168 86 0.2 3 1.0 1.5 20-25 14.8 438 442 167 92 0.2 4 1.2 1.6 25-30 15.2 430 408 174 85 0.2 5 1.3 1.8 30-35 15.2 444 452 169 82 0.2 5 1.4 2.0 35-38 15.2 442 440 173 75 0.3 6 1.4 2.1 Core 04mc Al Mn Ba„ Zr Ag Cd Re Mo U (wt.%) (ng/g) (ug/g) (ug/g) (ng/g) (ug/g) (ng/g) (ng/g) (ng/g) 0-1 11.7 404 291 129 0.1 2 0.4 1-2 32 0.1 1 0.3 0.9 2-3 - 0.1 2 0.4 0.7 3-4 27 0.1 1 0.4 0.7 4-5 27 0.1 2 0.3 0.7 5-6 11.9 397 293 146 32 0.1 2 0.3 0.9 6-7 27 0.1 2 0.3 0.8 7-8 26 0.1 3 0.3 0.9 8-9 33 0.2 4 0.4 1.2 9-10 47 0.3 5 0.4 1.1 10-11 14.3 550 397 144 70 0.3 5 0.6 1.4 11-12 79 0.2 2 0.8 1.9 12-13 77 0.3 3 0.8 1.5 13-14 76 0.2 4 0.9 1.4 14-16 15.9 672 451 144 81 0.2 3 0.7 1.6 16-18 62 0.3 5 0.5 1.0 Core 06bc Al Mn Bam Zr Ag Cd Re Mo U (wt.%) (ug/g) (ug/g) (ng/g) (ng/g) (ng/g) (ng/g) (ng/g) (ng/g) 0-0.5 13.3 495 498 157 97 0.2 4 0.6 1.5 0.5-1 94 0.2 3 0.5 1.0 1-2 84 0.2 3 0.5 1.3 2-3 100 0.3 4 0.6 1.4 3-4 103 0.3 4 0.7 1.4 4-5 13.8 473 495 168 117 0.2 4 0.6 1.5 5-7 98 0.2 4 0.5 1.5 7-9 94 0.2 4 0.5 1.4 9-11 13.6 472 465 176 106 0.2 4 0.5 1.6 11-13 85 0.2 4 0.4 1.0 13-15 13.6 496 487 169 106 0.2 5 0.6 1.6 17-19 91 0.3 6 0.5 1.2 21-23 16.7 718 622 131 106 0.4 9 0.8 2.2 25-27 16.0 710 591 135 96 0.2 5 1.1 2.4 29-31 106 0.3 4 1.4 2.2 33-35 16.5 791 555 129 119 0.2 3 1.2 2.0 continued on next page 125 Table 4.5 (continued) Dre 09mc Al Mn Ba„ Zr Ag Cd Re Mo U (wt.%) (ng/g) (ug/g) (ug/g) (ng/g) (ug/g) (ng/g) (ug/g) (ug/g 0-1 13.6 492 624 146 223 0.2 4 0.9 1.6 1-2 213 0.2 4 0.6 1.2 2-3 215 0.3 4 0.7 1.4 3-4 219 0.3 5 0.6 1.4 4-5 247 0.3 5 0.7 1.4 5-6 14.7 456 624 139 248 0.3 6 0.6 1.4 6-8 231 0.2 6 0.6 1.3 8-10 235 0.3 7 0.7 1.6 10-12 15.4 492 617 151 187 0.3 8 0.6 1.7 12-14 215 0.3 8 0.7 1.6 14-16 14.9 475 633 153 236 0.3 10 0.6 1.7 16-18 236 0.5 12 0.7 1.9 18-20 295 0.9 22 0.7 2.3 20-22 14.0 542 598 155 251 0.7 25 0.6 2.4 22-24 271 1.0 30 0.8 3.7 24-26 15.6 534 596 156 275 1.0 53 0.8 4.3 26-28 254 0.7 65 0.8 5.4 28-30 245 0.6 50 1.0 5.8 30-32 15.8 570 608 155 221 0.4 41 1.0 5.0 32-34 212 0.5 22 1.0 4.9 34-36 15.7 533 591 147 236 0.5 17 1.0 4.4 36-38 207 0.7 15 0.9 4.2 38-40 16.2 574 618 151 3 0.7 17 1.1 3.8 >re 02mc Al Mn Ba„ Zr Ag Cd Re Mo U (wt.%) (ug/g) (ug/g) (ug/g) (ng/g) (ug/g) (ng/g) (ug/g) (ug/g 0-1 12.4 521 687 135 330 0.2 4 0.7 1.2 1-2 335 0.2 4 0.7 1.3 2-3 420 0.2 4 0.7 1.4 3-4 462 0.3 4 0.7 1.5 4-5 458 0.2 4 0.7 1.5 5-6 16.7 430 759 145 475 0.4 5 0.7 1.6 6-8 476 0.5 5 0.7 1.8 8-10 473 0.5 9 0.7 1.7 10-12 14.7 446 736 146 430 0.4 11 0.6 1.6 12-14 428 0.6 15 0.9 1.8 14-16 13.0 451 697 154 434 0.6 25 0.9 2.1 16-18 422 0.6 30 0.8 2.8 18-20 14.4 455 705 156 388 0.5 17 0.9 2.8 are 05bc Al Mn Ba» Zr Ag Cd Re Mo U (wt.%) (ug/g) (ug/g) (ug/g) (ng/g) (ng/g) (ng/g) (ng/g) (ng/g 0-1 14.4 723 920 132 357 0.1 1 0.5 1.0 1-2 14.2 543 911 129 375 0.1 2 0.4 1.1 2-3 13.7 473 902 124 466 0.1 2 0.4 1.0 3-4 400 0.2 3 0.4 1.3 4-5 431 0.1 2 0.4 1.1 5-6 15.2 475 1057 135 478 0.1 3 0.3 1.2 6-7 501 0.1 3 0.4 1.1 7-8 553 0.2 3 0.4 1.1 8-9 506 0.2 3 0.3 1.1 9-10 474 0.2 3 0.4 1.1 10-12 15.3 472 1046 136 573 0.3 4 0.5 1.5 12-14 538 0.2 5 0.4 1.4 14-16 15.2 478 1045 143 555 0.2 10 0.4 1.7 continued on next page 126 Table 4.5 (continued) Core05bc Al Mn Ba,„ Zr Ag Cd Re Mo U (wt.%) (ng/g) (ug/g) (ug/g) (ng/g) (ng/g) (ng/g) (ug/g) (ug/g) 16-18 584 0.4 14 0.5 1.9 18-20 527 0.3 35 0.6 2.6 20-22 15.5 503 970 145 513 0.7 43 0.5 2.6 22-24 519 0.6 18 0.6 2.5 24-26 15.3 503 935 144 453 0.6 42 0.8 2.9 26-28 410 0.6 27 0.8 2.9 28-30 459 0.5 30 0.6 1.9 30-32 15.7 525 904 142 423 0.7 32 0.6 2.8 32-34 477 0.7 33 0.8 2.9 34-36 15.9 539 892 147 441 0.5 31 0.8 2.7 36-38 413 0.4 16 0.6 2.7 38-40 442 0.4 27 0.7 2.6 40-42 15.7 552 848 140 445 0.4 15 0.7 2.9 42-44 405 0.4 12 0.9 3.1 44-46 15.5 586 774 135 385 0.4 19 0.9 3.2 46-48 402 0.5 26 1.2 2.9 127 Previous work has shown that much of the variation in 813Corg values of sediments from the continental margin off Vancouver Island is due to changes in the relative proportions of terrestrial and marine organic matter (Chapter 2). It is possible to estimate the fraction of terrestrial organic carbon present in near-surface sediments using the following equation: Terrigenous fraction = (5,3Csmpl - 513Cmar)/(513Cterr - 513Cmar) (4.2) where 813Csmpl, 813Cmar and 513Cterr are the isotopic compositions of the sample, the marine end member and the terrestrial end member, respectively. Assuming a S13Cmar value of -21 %o and a 513Cterr value of -27 %o, as suggested in Chapter 2, it is estimated that Holocene sediments contain on average 8 % terrestrial organic matter with the highest concentration (22 %) being observed in shelf core Olmc. In comparison, the Pleistocene gray clay, which is present in the lower portion of Cores 04mc and 06bc, contains significantly more terrestrial organic matter (> 39 %). This result is consistent with a range of other data (i.e., n-alkane concentrations, 14C data, 813Corg and 815N values) from the Pleistocene deposits in Piston Core JT96-09 (Chapter 2). The relatively low 815N values measured for Pleistocene sediments in 06bc most probably reflect the presence of abundant terrestrial organic matter (>40 %). However, the significantly lower 815N values observed in Olmc, which contains <22 % terrestrial organic matter must have another explanation. Station 01 is located on the shelf and is thus directly influenced by newly upwelled, nitrate-rich water. When nitrate is abundant phytoplankton preferentially utilize 14N-enriched nitrate, producing isotopically light organic matter and leaving isotopically heavier residual nitrate (Montoya, 1994 and references therein). As waters are advected away from the site of upwelling they become further 15N-enriched due to continued biological activity and removal of 14N in sinking organic detritus. Phytoplankton growing at sites increasingly more remote from the source of upwelled nutrients therefore 128 exhibit higher 815N signatures and the organic matter which eventually reaches the sediment is similarly enriched (e.g., Farrell et al., 1995; Martinez et al., 2000). This process most probably explains the low 815N values of shelf sediments (i.e., Core Olmc) and the higher 815N values exhibited by continental slope sediments off Vancouver Island. However, data from sediment traps moored on the upper slope off Vancouver Island indicate that 815N values of settling particulate organic matter (SPOM) ranges from +7.7 %o during bloom events in summer and fall up to +9.0 %o in winter, with an average annual flux-weighted value of +8.1 %o (Wu, et al., 1999). Surface sediments from the slope have lower values (+6.3 to +6.8 %o) than this SPOM, possibly due to the input of minor terrestrial organic matter and/or the downslope transport of ,4N-enriched organic matter from the shelf. 4.4.2 Iodine and manganese The Mn/Al ratios of all surface sediments are substantially lower than the average shale value (0.0106; Turekian and Wedepohl, 1961), even in 05bc which exhibits the thickest oxic surface layer. Low Mn concentrations indicate that within the upper 1 cm these sediments become sufficiently reducing to cause Mn(IV), present in oxyhydroxides, to be reduced to soluble Mn(II). Unfortunately, the sampling resolution does not permit the exact depth where Mn reduction begins to be determined. It is well known that iodine species are adsorbed by marine organic matter at the sediment-water interface (Price and Calvert, 1977). Such adsorption occurs only under oxic conditions, where iodate (I03~) is the principal dissolved species of iodine. Under suboxic conditions, where I03" is reduced to I", no adsorption occurs, as I" is not taken up by organic detritus. Therefore, sediments that accumulate in suboxic and anoxic settings are characterized by low I/Corg values (typically <20 x 104) while sediments in well oxygenated environments typically yield I/Corg values of >200 x 10"4 (Price and Calvert, 1977; Francois, 1987). Surface sediments in all cores from the Vancouver Island Margin, with the exception 129 of Olmc, are characterized by I/Corg ratios that exceed 200 x 10"4 (Fig. 4.5b). These data suggest that in general surface sediments in the area are sufficiently oxygenated to allow adsorption of iodine species onto organic matter. Iodine is then lost relative to carbon during the degradation of organic matter (Price and Calvert, 1973 and 1977; Francois, 1987) resulting in a commonly-observed decrease in I/Corg ratios with depth. However, in shelf Core Olmc I/Corg ratios are consistently low, even in surface sediments. Carbon-isotope data indicate that Olmc contains only slightly more terrestrial organic matter than the other cores. Therefore, the low I/Corg values of surface sediments in Olmc cannot be attributed to an abundance of terrigenous organic matter, which does not adsorb dissolved iodine species (Malcolm and Price, 1984). Rather, the low I/Corg ratios imply that, despite the presence of well-oxygenated bottom waters (> 2 ml/1 02), surface sediments at station 01 are more reducing in comparison to the other stations. This may related to the relatively high organic carbon mass accumulation rate (-0.6 g/cm2/kyr) at this site. The combination of high I/Corg and low Mn/Al ratios in most cores indicates that the near-surface redox environment is perched within the suboxic zone, between Mn and I reduction. This is consistent with the lack of Re enrichment in the upper 10 to 15 cm (see Section 4.3). Reducing conditions appear to be more intense in Core Olmc; hence both Mn/Al and I/Corg ratios are low in the surface sediments. 4.4.3 Rhenium Rhenium concentrations are low in oxic marine sediments (< 0.5 ng/g; Boyko et al., 1986; Koide et al., 1986) but may be as much as 300 times higher in suboxic and anoxic deposits (Koide et al., 1986; Ravizza et al., 1991; Colodner et al., 1993; Crusius et al., 1996; Morford and Emerson, 1999). Enrichments of this magnitude are the result of Re diffusion from the overlying water into the sediment followed by its reduction from Re(VII) to Re(IV) under suboxic conditions and precipitation, possibly as Re02 (e.g., Crusius et al., 1996). In 130 general, significant enrichment occurs when porewater oxygen is consumed within 1 cm of the sediment-water interface (Morford and Emerson, 1999), and when sedimentation rates are low enough to allow Re to build-up in the deposits through ongoing diffusion and precipitation. The concentration of Re in surface sediments from the Vancouver continental margin is only slightly elevated (1.6 to 4.2 ng/g) when compared to the typical concentrations in oxic sediments (<0.5 ng/g). Thus, despite the development of suboxic conditions with 1 cm of the sediment-water interface, as indicated by the low Mn/Al ratios, there is no evidence of a large Re enrichment in the upper 10 to 15 cm of any of these cores. This implies that sedimentary redox conditions are not sufficiently reducing to cause Re reduction and precipitation at depths shallower than about 10 cm, which is consistent with I/Corgdata. With the exception of Core Olmc, I/Corgratios are relatively high, indicating that iodate persists in the uppermost porewater. According to Thomson et al. (1993) Re reduction should occur subsequent to I reduction. Thus, Re reduction and accumulation would not be anticipated in these near-surface sediments. However, below -15 cm in cores 09mc, 02mc and 05bc Re is clearly enriched (Fig. 4.6). We hypothesize that this reflects ongoing Re diffusion to depths on the order of 20 cm and subsequent reduction and precipitation. Unfortunately, no porewater data are available to permit a direct assessment of this hypothesis. Instead, a simple steady-state linear-diffusion model is applied to determine whether or not the enrichments could reflect ongoing authigenesis. We use Fick's First Law, J = (0/F)Db(dc/dz) (4.3) to compute the downward diffusion flux (J). Simple calculations imply that Re diffusion and sub-surface precipitation is a dominant control on sedimentary Re profiles. The bulk in situ diffusivity (Db) of the perrhenate ion is estimated to be similar to that of other oxyanions (i.e. - 6 x 10"6cm2 sec"1; Li 131 and Gregory, 1974). The porosity (O) is assigned a value of 0.8, while the formation factor (F) is taken to be 1.3 (Manheim, 1970), a value typical of silty clays. The concentration gradient (dc/dz) is assumed to be linear and is computed as the quotient of the difference in concentration between bottom water (7.8 ng kg"1; Anbar et al., 1992) and a precipitation point 20 cm below the sediment-water interface where the porewater Re content is assumed to be zero. This yields an estimated flux of 45 ng cm/kyr"1. The mass accumulation rate (MAR) of the sediments in Core 09mc, computed as the product of the linear sedimentation rate (~5 cm/kyr) and the dry bulk density [(1- O) times the grain density of 2.5 g cm"3], is 2.5 g cm" 2kyr"'. The ratio of J/MAR represents the steady-state concentration of Re that could be added to the sediments via downward diffusion to a fixation depth of 20 cm, and in this example that equals -18 ng/g. This is of the same order of magnitude as the observed Re concentration at 20 cm depth in 09mc (i.e., 22 ng/g). Application of the same calculation to Core 02mc which has a higher sedimentation rate (-23 cm/kyr) and a shallower precipitation point (10 cm) yields a steady-state Re enrichment of 9 ng/g which is also close to what is observed in the core. These admittedly crude and assumption-dependent results suggest that downward diffusion and sub-surface precipitation is indeed a credible mechanism for enriching Re in these deposits. However, the degree of enrichment is very sensitive to sedimentation rate. When redox boundaries are relatively deep, as they are in the Vancouver margin cores, large enrichments are only possible when sedimentation rates are low. 4.4.4. Uranium In oxygenated seawater U(VI) occurs as the uranyl-carbonate complex U02(C03) 3"4 (Langmuir, 1978). Under suboxic conditions U(VI) is reduced to U(IV) at approximately the same redox potential as Fe reduction (Langmuir, 1978; Cochran et al., 1986; Zheng, et al., 2002) or at a slightly lower potential (Klinkhammer and Palmer, 1991). Diffusion of uranium into sediments followed by its reduction and precipitation as U02 under suboxic and 132 anoxic conditions can lead to U enrichments in excess of 5 pg/g where sedimentation rates are relatively low (Thomson et al., 1990; Klinkhammer and Palmer, 1991; Legeleux et al., 1994; Crusius et al., 1996; Nath et al., 1997). Uranium, like Re, is not significantly enriched in the upper 10 to 15 cm of the Vancouver Island margin sediments (i.e., <2 ug/g). An obvious authigenic accumulation (up to 5.8 pg/g) is only observed below a depth of 20 cm in 09mc. There is also a suggestion of U enrichment in Cores 02mc and 05bc, but because U has a much higher lithogenic background relative to the concentration of dissolved U in seawater the authigenic signal is more difficult to detect (Chapter 5). Site 09 is overlain by the most severely oxygen-depleted water, which should lead to more intense reducing conditions than at the other sites, all other factors being equal. In turn, this should lead to greater authigenic enrichment of redox-sensitive metals like Re and U, which is exactly what is observed. 4.4.5. Cadmium Cadmium may be enriched by a few pg/g in suboxic sediments (Rosenthal et al., 1995b; van Geen et al., 1995; McCorkle and Klinkhammer, 1991) and is often highly enriched (>8 pg/g) in anoxic sediments (Pedersen et al., 1989; van Geen et al., 1995; Morford et al., 2001; Ivanochko, 2001). The most probable mechanism for such enrichment is the formation of CdS (Rosenthal et al., 1995a). Thus, although Cd is not directly affected by redox conditions, its enrichment does require the presence of at least trace amounts of dissolved sulphide (Pedersen et al., 1989; Rosenthal et al., 1995b; van Geen et al., 1995), which is related to the sedimentary redox state. Cadmium is supplied to reducing sediments by diffusion from the overlying water column and in sinking organic detritus. The organically-bound Cd is released into the porewater as organic matter degrades (McCorkle and Klinkhammer, 1991; Rosenthal et al., 1995b) and will be incorporated into the solid phase if dissolved sulphide is present. Unlike many other trace metals, the Cd concentration 133 in the upper water column can fluctuate rapidly when upwelling of nutrient-rich (i.e., Cd-rich) waters changes (van Geen and Husby, 1996; Sanudo-Wilhelmy and Flegal, 1996; Segovia-Zavala et al., 1998). It is therefore possible that sedimentary Cd concentrations could be coupled to changes in upwelling via the incorporation of Cd into organic matter and its subsequently exported to the sediment. Upwelling would also increase bottom water Cd concentrations and thus the diffusive flux into the sediments would also increase. The concentration of Cd is low (< 0.3 pg/g) in the weakly suboxic, near-surface sediments off Vancouver Island. This result is consistent with porewater data that indicate no discernible sulphate reduction in these sediments, with the possible exception of Core Olmc (A. Mucci, unpubl. data). However, there is a slight Cd enrichment in the lower portion of Cores 09mc, 02mc and 05bc which suggests that minor sulphate reduction may be occurring at depth. The enrichment is greatest in Core 09mc (up to 1 pg/g) which is consistent with Re and U data that imply stronger reducing conditions at this OMZ site. 4.4.6. Molybdenum In oxic sediments Mo concentrations can be high (>60 pg/g) as a result of Mo adsorption onto Mn oxyhydroxides (Bertine and Turekian, 1973; Calvert and Price, 1977; Shimmield and Price, 1986), but under suboxic conditions oxyhydroxides dissolve and adsorbed Mo is released into the porewater. As a result, suboxic sediments contain only lithogenic concentrations of Mo (~0.6 pg/g for terrigenous sediments derived from Vancouver Island; Morford et al., 2001). In contrast, anoxic sediments contain as much as 130 pg/g Mo (Bertine and Turekian, 1973; Koide et al., 1986; Francois, 1988; Emerson and Huested, 1991; Crusius et al., 1996; Adelson et al., 2001; Zheng et al., 2000a). Molybdenum enrichment is typically associated with the precipitation of Fe-sulphides (Bertine, 1972; Huerta-Diaz and Morse, 1992), but at low dissolved sulphide concentrations (<100 pM; Zheng et al., 2000b), it is not related to Mo reduction and precipitation of MoS2. Rather, 134 molybdate diffuses into the sediment from the overlying water column and is converted to thiomolybdate (Bertine, 1972; Helz et al., 1996) which is rapidly adsorbed by Fe-bearing particles such as Fe-sulphides (Helz et al., 1996). The conversion of molybdate (Mo042) to thiomolybdate (MoS42) appears to occur at a threshold H2S concentration of-11 pM (Erickson and Helz, 2000), although Mo accumulation has been documented at sulphide concentrations as low as 0.1 pM (Zheng et al., 2000b). The low Mo concentrations (<1.4 pg/g) observed in all box- and multicores suggest that, in those sediments sampled, fully anoxic conditions never develop (i.e., active sulphate reduction is limited) and the formation of Fe-sulphides is negligible. This interpretation is consistent with the lack of a large Cd enrichment and with porewater data that show little, if any, depletion of porewater sulphate (A. Mucci, unpubl. data). The only possible exception is Core Olmc where porewater data indicate the occurrence of minor sulphate reduction (A. Mucci, unpubl. data). This is also the core with the highest Mo content. The fact that anoxic conditions are not developed within the upper 50 cm, even in sediments underlying the OMZ, is unexpected given the relatively high organic carbon content of the deposits and low bottom water oxygen concentration, both of which should favour the early (i.e., shallow) onset of sulphate reduction. The dominance of suboxic, rather than anoxic, conditions in these sediments presumably reflects the moderately low rates of hemipelagic sedimentation and the relatively deep bioturbation as indicated by 210Pb data. 4.4.7. Barium Barium is found in a number of mineral phases within sediments. Lithogenic barium occurs primarily within aluminosilicates and in particular potassium feldspar where Ba substitutes for K. Barium also occurs as discrete barite particles that form within decaying organic material (Dehairs et al., 1980; Bishop, 1988; Ganeshram et al., 2003). This biogenic Ba is commonly enriched in sediments that underlie highly productive surface waters 135 (Dehairs et al., 1980; Dehairs et al., 1992), where the supply of organic particles is high (Dymond and Collier, 1996). In general, sedimentary Ba concentrations and Ba/Al ratios increase with increasing water depth off Vancouver Island (Fig. 4.5c). A similar Ba-depth relationship has been observed in sediments off the coasts of Chile (Klump et al., 2000), Peru (von Breymann et al., 1992), Namibia (Calvert and Price, 1983) and California (McManus et al., 1999), and in the Arabian Sea (McManus et al., 1998; Schenau et al., 2001). Three possible explanations have been put forth to explain increasing Ba content with increasing water depth: i) better preservation, ii) decreasing dilution and/or sediment focusing, and iii) an increase in the amount of biogenic barite delivered to the sediment. Preservation of barite may be influenced by a number of factors. Sulphate reduction in anoxic sediments leads to undersaturation and dissolution of barite (von Breymann et al., 1992; Falkner et al., 1993; McManus et al., 1994), but this cannot explain the pattern of barite accumulation in Holocene sediments off Vancouver Island because there is no evidence that significant sulphate reduction is occurring in the upper 50 cm of these deposits. Dymond et al. (1992) suggested that barite preservation increases with increasing sedimentation rate because barite is more rapidly isolated from undersaturated bottom waters. Off Vancouver Island however, the core with the highest sedimentation rate (i.e., Olmc, 40 cm/kyr) has the lowest Ba/Al ratios while cores with relatively low sedimentation rates (i.e., 09mc and 05bc, 5 cm/kyr) exhibit much higher Ba/Al ratios. Thus, sedimentation rate does not appear to control barite preservation in the study area. If the increase in Ba with depth is caused by decreasing dilution or sediment focusing we would also expect other elements to be similarly affected. The concentration of Zr and the Zr/Al ratio both decrease with increasing water depth suggesting sediments become finer-grained offshore, and there is an inverse correlation between Zr and total Ba concentrations (Fig. 4.8). Given this result it is tempting to attribute the changes in Ba concentration solely to changes in sediment grain size; however, this would be wrong. The concentration of 136 1200 1000 -1 'So ^ 800 -1 *3 600 400 200 A A A AA A 05bc A A A • 120 o 04mc 130 ^,02mc • D 09mc O O 06bc / 140 150 160 \ Olmc 170 180 Zr (jig/g) Fig. 4.8. Concentration of Zr versus the concentration of total Ba near-surface sediments from the Vancouver Island margin. 137 lithogenic Ba (Balith) was measured in a number of samples from which the biogenic fraction (Babi0) had been chemically extracted. The Babio concentration was then estimated by difference (Babio = Batot - Balith). The results show that the concentration of Babio increases with water depth from ~96 ppm in 09mc up to 655 ppm in 05bc, while the Balith/Al ratio remains relatively constant (0.0026 to 0.0038; Table 4.2). A large proportion of biogenic barite is formed in the upper water column where labile organic matter rapidly decays (Chan et al., 1977; Dehairs et al., 1980; Bishop, 1988; Dehairs et al., 1990). However, sediment trap data suggest that some biogenic barite also forms deeper in the water column as organic material settles (Dymond and Collier, 1996; Klump et al., 2000). Thus, we attribute the observed increase in Babio (and Ba/Al ratios) with water depth to the continual formation of barite as organic material settles through the water column. 4.4.8. Silver The concentration of Ag in oxic, marine sediments is low (70 to 100 ng/g; Bowen, 1966), and largely reflects the lattice-held fraction (i.e., lithogenic Ag). In comparison, Ag content in natural anoxic, marine sediments can be quite high (> 400 ng/g; Koide et al., 1986). This authigenic accumulation does not involve a change in the redox state of the metal. Rather, Ag readily precipitates as Ag2S in the presence of trace amounts of dissolved sulphide because Ag2S is highly insoluble (pK = 36; Dyrssen and Kremling, 1990). In this respect the sedimentary geochemistry of Ag is similar to that of Cd. Dissolved Ag in seawater exhibits a "nutrient-type" depth profile (Martin et al., 1983; Flegal et al., 1995; Zhang et al., 2001) which is most similar to Si suggesting it is taken up by marine organisms and incorporated into their "hard parts" (Flegal et al., 1995; Zhang et al., 2001). Laboratory culture experiments show that Ag is actively incorporated into diatom frustules (Fisher and Wente, 1993; Lee and Fisher, 1994) and it is probable that some Ag is thus supplied to the sediments with the diatom flux (Friedl et al., 1997). 138 Cores Olmc, 04mc and 06bc contain background (i.e., lithogenic) concentrations of Ag (< 100 ng/g) while cores 09mc, 02mc and 05bc are enriched in Ag (220 to 360 ng/g). The enrichments are not the result of in situ diagenesis under anoxic conditions. Iodine, Mn, Re and Mo data indicate that the upper 10 to 15 cm in these cores is only weakly suboxic, not reducing enough for Re and U precipitation, and certainly not anoxic. Furthermore, there is no correlation between Ag content and either opal or organic carbon contents (Figs. 4.9a and b, respectively) implying that the sedimentary Ag concentration is not directly related to these biogenic fluxes. Grain-size does seem to have some influence, as shown by the negative correlation between Ag concentrations and Zr/Al ratios (Fig. 4.9c). However, this would imply that the lithogenic fraction of the fine-grained sediments contains almost 600 ng/g of Ag which is 6x higher than the concentration typically observed (i.e., -100 ng/g; Bowen, 1966), so this correlation is probably deceptive. A more probable explanation for the very high Ag concentrations and the increase in Ag content with water depth is suggested by the very strong positive correlation between sedimentary Ag and biogenic Ba (r2 = 0.89; Fig. 4.9d). To our knowledge, this is the first report of a strong correlation between Ag and Ba concentrations in marine sediments and it is geochemically puzzling. The positive association suggests that Ag and Ba may be transported to the sediment in a similar manner. Barium enrichment of sinking organic detritus results from the formation of barite during settling. Recognizing that both dissolved Ba and Ag have similar overall distributions in the marine water column, the close correspondence between the two elements reported here implies that Ag is also scavenged from the water column by settling organic particles. Ag enrichment in these particles cannot be related to the substitution of Ag for Ba in barite since their ionic radii are dissimilar (1.29 and 1.49 A, respectively) as is their valence. Nor can the correlation be the result of the precipitation of Ag sulphate, which is highly soluble. The leading candidate for Ag sequestration by sinking particles is the precipitation of a sulphide phase (e.g., Ag2S), but this presents a paradox. The formation of barite and Ag2S should be 139 05bc A A/ 02mc O 09mc O 06bc a o 04mc Olmc 10 % Opal 12 600 500 ^ 400 I 3 300 < 200 100 0 OSbc A /fe* A 06bc \ ,o o 4? °° • • 02mc o o 49 04mc 0 0.5 1 1.5 2 2.5 3 % Organic Carbon 3.5 A A c A OSbc " 1 *V A ° • A -A a 02mc • O ^> 09mc O O Olmc . 06bc / * • • o04mc 0.0008 0.0010 Zr/Al 0.0012 600 500 400 60 & 300 200 100 0 4 d A • A * ° 05bc A a A (1750m) A ° 02mc (1340m) Olmc (120m) / o o 0 09mc (920m) 1 **x 0<_ 04mc (407m) 06bc (720m) 200 400 600 Biogenic Ba (ug/g) 800 Fig. 4.9. Plots of Ag concentrations versus a) % opal, b) % organic carbon, c) Zr/Al ratios, and d) biogenic Ba content, in near-surface sediment cores. Note that there is a positive correlation between Ag and biogenic barium. A linear regression of these data yield the equation y = -22.22 + 1.1622x (r2 = 0.89). 140 mutually exclusive since barite forms under oxic conditions and dissolves under anoxic conditions, while the precipitation of Ag2S requires anoxia. We therefore tentatively hypothesize that within settling organic particles both oxic and anoxic microenvironments exist. Within the interior of sinking particles the degradation of organic matter could lead to the development of an anoxic microenvironment where sulphate reduction occurs and Ag2S precipitates. Ag would be supplied from the water column and by opal dissolution. Surrounding the "anoxic core" the particle would remain oxic due to contact with the surrounding oxygenated seawater, and in this outer zone barite formation would occur. Although we have no direct evidence in support of this model, the higher concentration of Ag in sediments from deeper locations does imply that there is continual addition of Ag to particles as they settle. Confirmation or refutation of this hypothesis will require further work. 4.5 Summary Despite the development of suboxic conditions within millimetres of the sediment-water interface there is little enrichment of redox-sensitive trace elements in the upper 10 to 15 cm of continental margin sediments off Vancouver Island. These results are consistent with minor element data that indicate near-surface sediments are poised within the suboxic zone. Below the depth of oxygen penetration (-10 to 15 cm) sediments do become sufficiently reducing to allow Re and U enrichment. This enrichment is greatest in Core 09mc, the only core collected from the OMZ, and reflects a lower sedimentation rate (5 cm/kyr) and low bottom water oxygen concentration at this location. There is no evidence that these sediments become fully anoxic (i.e., no Mo enrichment) despite productive surface waters and relatively low bottom water oxygen concentrations in the region. Suboxic conditions appear to be maintained because slow sedimentation and extensive bioturbation 141 allow a small, but continual, influx of oxygen and other oxidants from the overlying water column. Silver concentrations in these continental margin deposits correlate closely with Ba/Al ratios and water depth. This is the first report of such an association and it presents a geochemical puzzle. The data imply that sedimentary Ag content is in large part controlled by the settling flux of particulate Ag and not exclusively by redox conditions within the sediment. We hypothesize that Ag, like Ba, may accumulate within organic particles as they settle through the water column. Thus, at a given location where water depth has been relatively invariant, sedimentary Ag concentrations may provide another palaeo-productivity proxy. However, at least two major questions need to be answered before such this occurs. First, the hypothesis that Ag precipitates as Ag2S within anoxic microenvironments in setting organic particles must be explored, and second, the conditions under which this "biogenic" Ag is preserved need to be clarified. 142 4.6 References Adelson, J.M., Helz, G.R., Miller, C.V., 2001. Reconstructing the rise of recent coastal anoxia; molybdenum in Chesapeake Bay sediments. Geochimica et Cosmochimica Acta 65, 237-252. Anabar, A.D., Creaser, R.A., Papanastassiou, D.A., Wasserburg, G.J., 1992. Rhenium in seawater: Confirmation of generally conservative behavior. Geochimica et Cosmochimica Acta 56, 4099-4103. Antoine, D., Andre, J.M., Morel, A., 1996. Oceanic primary production. 2. Estimation at global scale from satellite (coastal zone colour scanner) chlorophyll. Global Biogeochemical Cycles 10,57-69. Bertine, K.K., 1972. The deposition of molybdenum in anoxic waters. Marine Chemistry 1, 43-53. Bertine, K.K., Turekian, K.K., 1973. Molybdenum in marine deposits. Geochimica et Cosmochimica Acta 37, 1415-1434. Bishop, J.K.B., 1988. The barite-opal-organic carbon association in oceanic particulate matter. Nature 332, 341-343. Bornhold, B.D., Yorath, C.J.,1984. Surficial geology of the continental shelf, northwestern Vancouver Island. Marine Geology 57, 89-112. Bornhold, B.D., Barrie, J.V., 1991. Surficial sediments on the Western Canadian Continental Shelf. Continental Shelf Research 11, 685-699. Bowen, H.J.M., 1966. Trace Metals in Biogeochemistry. Academic Press, New York. Boyko, T.F., Baturin, G.N., Miller, A.D., 1986. Rhenium in recent ocean sediments. Geochemistry International 23, 38-47. Calvert, S.E., 1976. The mineralogy and geochemistry of near-shore sediments. In: J.P. Riley, J.P., Chester, R. (Eds.), Chemical Oceanography, Vol. 6. Academic Press, London, 187-280pp. Calvert, S.E., 1990. Geochemistry and origin of the Holocene sapropel in the Black Sea. In: Ittekkot, V., Kempe, S., Michaelis, W., Spitzy, A. (Eds.), Facets of Modern Biogeochemistry. Springer-Verlag, Berlin, 326-352pp. Calvert, S.E., Price, N.B., 1977. Geochemical variations in ferromanganese nodules and associated sediments from the Pacific Ocean. Marine Chemistry 5,43-74. Calvert, S.E., Price, N.B., 1983. Geochemistry of Namibian shelf sediments. In: Suess, A., Thiede E. (Eds.), Coastal Upwelling. J. Plenum Publishing Corporation, 337-375pp. Carpenter, R., Bennett, J.T-, Peterson, M.L., 1981.210Pb activities in and fluxes to sediments of the Washington continental slope and shelf. Geochimica et Cosmochimica Acta 45, 1155-1172. 143 Chan, L.H., Drummond, D., Edmond, J.M., Grant, B., 1977. On the barium data from the Atlantic GEOSECS expedition. Deep Sea Research 24, 613-649. Cochran, J.K., Carey, A.E., Sholkovitz, E.R., Surprenant, L.D., 1986. The geochemistry of uranium and thorium in coastal marine sediments and sediment pore water. Geochimica et Cosmochimica Acta 50, 663-680. Colodner, D., Sachs, J., Ravizza, G., Turekian, K., Edmond, J., Boyle, E., 1993. The geochemical cycle of rhenium: A reconnaissance. Earth and Planetary Science Letters 117, 205-221. Crusius, J., Calvert, S., Pedersen, T., Sage, D., 1996. Rhenium and molybdenum enrichments in sediments as indicators of oxic, suboxic and sulfidic conditions of deposition. Earth and Planetary Science Letters 145, 65-78. Crusius, J., Pedersen, T.F., Calvert, S.E., Cowie, G.L., Oba, T., 1999. A 36kyr geochemical record from the Sea of Japan of organic matter flux variations and changes in intermediate water oxygen concentrations. Paleoceanography 14, 248-259. Dean, W.E., Gardner, J.V., Piper, D.Z., 1997. Inorganic geochemical indicators of glacial-interglacial changes in productivity and anoxia on the California continental margin. Geochimica et Cosmochimica Acta 61, 4507-4518. Dean, W.E., Piper, D.Z., Peterson, LC, 1999. Molybdenum acccumulation in Cariaco basin sediments over the past 24 k.y.: A record of water-column anoxia and climate. Geology 27, 507-210. Dehairs, F., Chesselet, R., Jedwab, J., 1980. Discrete suspended particles of barite and the barium cycle in the open ocean. Earth and Planetary Science Letters 49, 528-550. Dehairs, F., Goeyens, L., Stroobants, N., Bernard, P., Goyet, C, Poisson, A., Chesselet, R., 1990. On suspended barite and the oxygen minimum in the Southern Ocean. Global Biogeochemical Cycles 4, 85-102. Dehairs, F., Baeyens, W., Goeyens, L., 1992. Accumulation of suspended barite at mesopelagic depths and export production in the Southern Ocean. Science 258, 1332-1335. Dryssen, D. Kremling, K., 1990. Increasing hydrogen sulfide concentration and trace metal behaviour in the anoxic Baltic waters. Marine Chemistry 30, 193-204. Dymond, J., Collier, R., 1996. Particulate barium fluxes and their relationships to biological productivity. Deep-Sea Research 43, 1283-1308. Dymond, J., Suess, E., Lyle, M., 1992. Barium in deep-sea sediments: A geochemical proxy for paleoproductivity. Paleoceanography 7, 163 -181. Eakins, J.D., Morrison, R.T., 1978. A new procedure for the determination of lead-210 in lake and marine sediments. International Journal of Applied Radiation and Isotopes 29, 531-536. Emerson, S.R., Huested, S.S., 1991. Ocean anoxia and the concentration of molybdenum and vanadium in seawater. Marine Chemistry 34, 177-196. 144 Erickson, B.E., Helz, G.R., 2000. Molybdenum(VI) speciation in sulfidic waters: Stability and lability of thiomolybdates. Geochimica et Cosmochimica Acta 64,1149-1158. Falkner, K.K., Klinkhammer, G.P., Bowers, T.S., Todd, J.F., Lewis, B.L., Landing, W.M., Edmond, J.M., 1993. The behavior of barium in anoxic marine waters. Geochimica et Cosmochimica Acta 57, 537-554. Farrel, J.W., Pedersen, T.F., Calvert, S.E., Nielsen, B., 1995. Glacial-interglacial changes in nutrient utilization in the equatorial Pacific Ocean. Nature 377, 514-517. Fisher, N.S., Wente, M., 1993. The release of trace elements by dying marine phytoplankton. Deep-Sea Research 40, 671-694. Flegal, A.R., Sanudo-Wilhelmy, S.A., Scelfo, G.M., 1995. Silver in the eastern Atlantic Ocean. Marine Chemistry 49, 315-320. Francois, R., 1987. The influence of humic substances on the geochemistry of iodine in nearshore and hemipelagic marine sediments. Geochimica et Cosmochimica Acta 51, 2417-2427. Francois, R., 1988. A study on the regulation of the concentrations of some trace metals (Rb, Sr, Zn, Pb, Cu, V, Cr, Ni, Mn and Mo) in Saanich Inlet sediments, British Columbia, Canada. Marine Geology 83, 285-308. Friedl, G., Pedersen, T.F., 1997. Relative enrichment of silver in marine sediments as a reflection of high paleoproductivity. Abstract, American Geophysical Union Fall Meeting, San Francisco. Ganeshram, R.S., 1996. On the glacial-interglacial variability of upwelling, carbon burial and denitrification on the northwestern Mexican continental margin. Ph.D. dissertation thesis, University of British Columbia. Ganeshram, R.S., Francois, R., Commeau, J., Brown-Leger, S.L.. 2003. An experimental investigation of barite formation in seawater. Geochimica et Cosmochimica Acta 67; 2599-2605. Gobeil, C, Sundby, B., Macdonald, R.W., Smith, J.N., 2001. Recent change in organic carbon flux to Arctic Ocean deep basins: Evidence from acid volatile sulfide, manganese and rhenium discord in sediments. Geophysical Research Letters 28 ,1743-1746. Helz, G.R., Miller, C.V., Charnock, J.M., Mosselmans, J.F.W., Pattrick, R.A.D., Garner, CD., Vaughan, D.J., 1996. Mechanism of molybdenum removal from the sea and its concentration in black shales: EXAFS evidence. Geochimica et Cosmochimica Acta 60, 3631-3642. Hickey, B.M., 1998. Coastal oceanography of Western North America from the tip of Baja California to Vancouver Island. In: Robinson, A.R., Brink, K.H. (Eds.), The Sea, Vol. 11, pp.345-393. Huerta-Diaz, M.A., Morse, J.W., 1992. Pyritization of trace metals in anoxic sediments. Geochimica et Cosmochimica Acta 56, 2681-2702. 145 Ivanochko, T.S., 2001. Productivity influences on oxygenation of the Santa Barbara Basin, California during the Late Quaternary. M.Sc. Thesis, The University of British Columbia, 133pp. Klinkhammer, G.P., Palmer, M.R., 1991. Uranium in the oceans: Where it goes and why. Geochimica et Cosmochimica Acta 55,1799-1806. Klump, J., Hebbeln, D., Wefer, G., 2000. The impact of sediment provenance on barium-based productivity estimates. Marine Geology 169, 259-271. Koide, M., Hodge, V.F., Yang, J.S., Stallard, M., Goldberg, E.G., Calhoun, J., Bertine, K.K., 1986. Some comparative marine chemistries of rhenium, gold, silver and molybdenum. Applied Geochemistry 1, 705-714. Langmuir, D., 1978. Uranium solution-mineral equilibria at low temperatures with applications to sedimentary ore deposits. Geochimica et Cosmochimica Acta 42, 547-569. Lee, B., Fisher, N., 1994. Effects of sinking and zooplankton grazing on the release of elements from planktonic debris. Marine Ecology Progress Series 110, 271-281. Legeleux, F., Reyss, J.-L., Bonte, P., Organo, C, 1994. Concomitant enrichments of uranium, molybdenum and arsenic in suboxic continental margin sediments. Oceanologica Acta 17,417-430. Li, Y.-H., Gregory, S., 1974. Diffusion of ions in sea water and deep-sea sediments. Geochimica et Cosmochimica Acta 38, 703-714. Mackas, D.L., Denman, K.L., Bennett, A.F., 1987. Least squares multiple tracer analysis of water mass coomposition. Journal of Geophysical Research 31, 2907-2918. Malcolm, S.J., Price, N.B., 1984. The behaviour of iodine and bromine in estuarine surface sediments. Marine Chemisty 15, 263-271. Manheim, F.T., 1970. The diffusion of ions in unconsolidated sediments. Earth and Planetary Science Letters 9, 307-309. Martin, J.H., Knauer, G.A., Gordon, R.M., 1983. Silver distributions and fluxes in north-east Pacific waters. Nature 305, 306-309. Martinez, P., Bertrand, P., Calvert, S.E., Pedersen, T.F., Shimmield, G.B., Lallier-Verges, E., Fontugne, M.R., 2000. Spatial variations in nutrient utilization, production and diagenesis in the sediments of a coastal upwelling regime (NW Africa): Implications for the paleoceanographic record. Journal of Marine Research 58, 809-835. McCorkle, D.C., Klinkhammer, G.P., 1991. Porewater cadmium geochemistry and the porewater cadmium:813C relationship. Geochimica et Cosmochimica Acta 55, 161-168. McManus, J., Berelson, W.M., Klinkhammer, G.P., Kilgore, T.E., Hammond, D.E., 1994. Remobilization of barium in continental margin sediments. Geochimica et Cosmochimica Acta 58, 4899-4907. McManus, J., Berelson, W.M., Klinkhammer, G.P., Johnson, K.S., Coale, K.H., Anderson, R.F., Kumar, N., Burdige, D.J., Hammond, D.E., Brumsack, H.J., McCorkle, D.C., Rushdi, 146 A., 1998. Geochemistry of barium in marine sediments: Implications for its use as a paleoproxy. Geochimica et Cosmochimica Acta 62, 3453-3473. McManus, J., Berelson, W.M., Hammond, D.E., Klinkhammer, G.P., 1999. Barium cycling in the North Pacific: Implications for the utility of Ba as a paleoproductivity and paleoalkalinity proxy. Paleoceanography 14, 53-61 Morford, J.L., Emerson, S., 1999. The geochemistry of redox sensitive trace metals in sediments. Geochimica et Cosmochimica Acta 63, 1735-1750. Morford, J.L., Russell, A.D., Emerson, S., 2001. Trace metal evidence for changes in the redox environment associated with the transition from terrigenous clay to diatomaceous sediment, Saanich Inlet, B.C.. Marine Geology 174, 355-369. Mortlock, R.A., Froelich, P.N., 1989. A simple method for the rapid determination of biogenic opal in pelagic marine sediments. Deep Sea Research 36, 1415-1426. Montoya, J.P., 1994. Nitrogen-isotope fractionation in the modern ocean: Implications for the sedimentary record. In: Zahn, R., Pedersen, T.F., Kaminski, M.A., Labeyrie, L. (Eds.), Carbon Cycling in the Glacial Ocean: Constraints on the Ocean's Role in Global Change, NATO ASI Series, Vol. 117, Springer-Verlag, Berlin, 259-279pp. Nameroff, T.J., Balistrieri, L.S., Murray, J.W., 2002. Suboxic trace metal geochemistry in the eastern tropical North Pacific. Geochimica et Cosmochimica Acta 66, 1139-1158. Nath, B.N., Bau, M., Rao, B.R., Rao, CM., 1997. Trace and rare earth elemental variation in Arabian Sea sediments through a transect across the oxygen minimum zone. Geochimica et Cosmochimica Acta 61, 2375-2388. Pailler, D., Bard, E., Rostek, F., Zheng, Y., Mortlock, R., van Geen, A., 2002. Burial of redox-sensitive metals and organic matter in the equatorial Indian Ocean linked to precession. Geochimica et Cosmochimica Acta 66, 849-865. Pedersen, T.F., Waters, R.D., MacDonald, R.W., 1989. On the natural enrichment of cadmium and molybdenum in the sediments of Ucluelet Inlet, British Columbia. The Science of the Total Environment 79, 125-139. Piper, D.Z., Isaacs, CM., 1995. Minor elements in Quaternary sediment from the Sea of Japan: A record of surface-water productivity and interemediate-water redox conditions. Geological Society of America Bulletin 107, 54-67. Price, N.B., Calvert, S.E., 1973. The geochemistry of iodine in oxidized and reduced recent marine sediments. Geochimica et Cosmochimica Acta 37, 2149-2158. Price, N.B., Calvert, S.E., 1977. The contrasting geochemical behaviours of iodine and bromine in recent sediments from the Namibian shelf. Geochimica et Cosmochimica Acta 41, 1769-1775. Ravizza, G., Turekian, K.K., Hay, B.J., 1991. The geochemistry of rhenium and osmium in recent sediments from the Black Sea. Geochimica et Cosmochimica Acta 55, 3741-3752. Rosenthal, Y., Boyle, E.A., Labeyrie, L., Oppo, D., 1995a. Glacial enrichments of authigenic Cd and U in subantarctic sediments: A climatic control on the elements oceanic budget? Paleoceanography 10, 395-413. 147 Rosenthal, Y., Lam, P., Boyle, E.A., Thomson, J., 1995b. Authigenic cadmium enrichments in suboxic sediments: Precipitation and postdepositional mobility. Earth and Planetary Science Letters 132, 99-111. Sanudo-Wilhelmy, S.A., Flegal, A.R., 1996. Trace metal concentrations in the surf zone and coastal waters off Baja California, Mexico. Environmental Science and Technology 30, 1575-1580. Schenau, S.J., Prins, M.A., De Lange, G.J., Monnin, C, 2001. Barium accumulation in the Arabian Sea: Controls on barite preservation in marine sediments. Geochimica et Cosmochimica Acta 65, 1545-1556. Segovia-Zavala, J.A., Delgadillo-Hinojosa, F., Alvarez-Borrego, S., 1998. Cadmium in the coastal upwelling area adjacent to the California-Mexico Border. Estuarine, Coastal and Shelf Science 46, 475-481. Shimmield, G.B., Price, N.B., 1986. The behaviour of molybdenum and manganese during early sediment diagenesis - offshore Baja California, Mexico. Marine Chemistry 19, 261-280. Stuiver, M., Reimer, P.J., Bard, E., Beck, J.W., Burr, G.S., Hughen, K.A., Kromer, B., McCormac, G., Van Der Plicht, J., Spurk, M., 1998. INTCAL98 radiocarbon age calibration, 24,000-0 cal BP. Radiocarbon 40, 1041-1083. Taylor, R.S., McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell Scientific, Boston, 312pp. Thomson, J., Wallace, H.E., Colley, S., Toole, J., 1990. Authigenic uranium in Atlantic sediments of the last glacial stage - a diagenetic phenomenon. Earth and Planetary Science Letters 98,222-232. Thomson, J., Higgs, N.C, Croudace, I.W., Colley, S., Hydes, D.J., 1993. Redox zonation of elements at an oxic-post-oxic boundary in deep-sea sediments. Geochimica et Cosmochimica Acta 57, 579-595. Thomson, R.E., 1981. Oceanography of the British Columbia Coast. Canadian Special Publication of Fisheries and Aquatic Sciences 56, 291p. Turekian, K.K., Wedepohl, K.H., 1961. Distribution of the elements in some major units of the earth's crust. Geological Society of America Bulletin 72, 175-192. van Geen, A., McCorkle, D.C, Klinkhammer, G.P., 1995. Sensitivity of the phosphate-cadmium-carbon isotope relation in the ocean to cadmium removal by suboxic sediments. Paleoceanography 10,159-169. van Geen, A., Husby, D.M., 1996. Cadmium in the California Current system: Tracer of past and present upwelling. Journal of Geophysical Research 101, 3489-3507. von Breymann, M.T., Emeis, K.-C, Suess, E., 1992. Water depth and diagenetic constraints on the use of barium as a palaeoproductivity indicator. In: Summerhayes, CP., Prell, W.L., Emeis, K.C. (Eds.), Upwelling Systems: Evolution Since the Early Miocene. Geological Society Special Publication 64, 273-284pp. 148 Wu, J., Calvert, S.E., Wong, C.S., 1999. Carbon and nitrogen isotope ratios in sedimenting particulate organic matter at an upwelling site off Vancouver Island. Estuarine, Coastal and Shelf Science 48, 193-203. Yang, Y.-L., Elderfield, H., Pedersen, T.F., Ivanovich, M., 1995. Geochemical record of the Panama Basin during the last glacial maximum carbon event shows that the glacial ocean was not suboxic. Geology 23, 1115-1118. Yarincik, K.M., Murray, R.W., Lyons, T.W., Peterson, L.C., Haug, G.H., 2000. Oxygenation history of bottom waters in the Cariaco Basin, Venezuela, over the past 578,000 years: Results from redox-sensitive metals (Mo, V, Mn, and Fe). Paleoceanography 15, 593-604. Yorath, C.J., Nasmith, H.W., 1995. The Geology of Southern Vancouver Island. Orca Book Publishers, Victoria, Canada, 172pp. Zhang, Y., Amakawa, H., Nozaki, Y., 2001. Oceanic profiles of dissolved silver: Precise measurements in the basins of western North Pacific, Sea of Okhotsk, and the Japan Sea. Marine Chemistry 75, 151-163. Zheng, Y., Anderson, R.F., van Geen, A., Kuwabara, J., 2000b. Authigenic molybdenum formation in marine sediments: A link to pore water sulfide in the Santa Barbara Basin. Geochimica et Cosmochimica Acta 64, 4165-4178. Zheng, Y., van Geen, A., Anderson, R.F., 2000a. Intensification of the northeast Pacific oxygen minimum zone during the Boiling-Allerod warm period. Paleoceanography 15, 528-536. Zheng, Y., Anderson, R.F., van Geen, A., Fleisher, M.Q., 2002. Remobilization of authigenic uranium in marine sediments by bioturbation. Geochimica et Cosmochimica Acta 66, 1759-1772. 149 5. Geochemical Response to Pulsed Sedimentation on the Western Canadian Continental Margin: Implications for the Use of Mo as a Palaeo-oxygenation Proxy 5.1 Introduction It has become common practice to use the concentrations of certain metals (e.g., Mo and Re) as palaeo-proxies for sedimentary redox conditions. From such data temporal changes in organic carbon flux to the sediment and bottom water oxygen concentrations, both of which play a large role in sedimentary redox conditions, are inferred (e.g., Dean et al., 1997 and 1999; Zheng et al., 2000; Adelson et al., 2001). However, other factors such as sedimentation rate, sediment composition, metal source and post-depositional remobilization can also influence metal accumulation. Reducing sediments are typically characterized by high concentrations of authigenic Re, U, Mo, Cd and Ag (e.g., Koide et al., 1986; Francois, 1988; Calvert, 1990; Calvert and Pedersen, 1993; Crusius et al., 1996; Morford et al., 2001; Nameroff, 2002). However, the redox conditions under which these metals accumulate and the manner in which they are fixed within the sediment differ. Rhenium and U typically are enriched in suboxic sediments (up to 150 ng/g and >5 pg/g, respectively) as a direct result of reduction and precipitation (Crusius et al., 1996). In contrast, Ag and Cd each have a single redox state but form insoluble sulphides when trace amounts of H2S are available (Koide et al., 1986; Rosenthal et al., 1995). This can lead to minor Ag and Cd accumulation in suboxic sediments and large accumulations in anoxic sediments (often >400 ng/g Ag and >8 pg/g Cd). Recent experimental results suggest that Mo enrichment is due to the formation of thiomolybdate that is rapidly adsorbed by Fe-bearing particles such as Fe-sulphides (Helz et al., 1996). Thiomolybdate formation requires a threshold H2S concentration of ~11 pM (Erickson and Helz, 2000) and thus Mo enrichment, often in excess of 10 pg/g, is restricted to anoxic sediments in which sulphate reduction has occurred. 150 In this paper we examine the distribution of redox-sensitive trace metals in Piston Core JT96-02, collected from a water depth of 1340 m off the west coast of Vancouver Island, Canada (49°12.8'N, 127°18.6'W; Fig. 5.1). The upper 16 cm of the core is composed of an olive green mud typical of Holocene deposits on the mid to lower continental slope in this area. This mud is underlain by 18 cm of greenish gray clay and a 56 cm thick zone of interbedded clays and sands with erosional contacts and loading features typical of turbidite deposits. The remainder of the core (90 to 211 cm) is a gray clay that upon opening of the core contained numerous thin (1 to 2 cm) black layers spaced ~5 to 10 cm apart and characterized by sharp upper surfaces and gradational lower boundaries. These black layers faded within hours, indicating that they were composed of FeS that was rapidly oxidized on exposure to air. Two AMS radiocarbon dates of a mixed assemblage of planktonic foraminifera (TV. pachyderma and G. bulloides) were obtained for the gray clay (13,280 ± 60 and 14,240 ± 60 14C yrs from 102 and 202 cm, respectively). At this time the Cordilleran ice sheet, which had reached it maximum extent in southern British Columbia between 15,000 and 14,000 14C yrs, was rapidly retreating (Clague and James, 2002). The continental shelf west of Vancouver Island was deglaciated by ~ 13,000 14C yrs (Blaise et al., 1990; Josenhans etal, 1995). The objectives of this study are: i) to explore the character of deglacial sedimentation on the Vancouver Island Margin; ii) to determine why the sulphide-rich layers formed; and iii) to understand better the controls on the accumulation of Mo and other redox-sensitive metals in continental margin sediments. To this end, a 16 cm interval containing three sulphide layers was sampled every 3 mm for high resolution geochemical analysis. Total carbon (Ctot), nitrogen (Ntot) and sulphur (Stot) were measured by high temperature combustion using a Carlo Erba NA-1500 CNS elemental analyzer. The relative standard deviation (RSD, la) is 3 %, 5 % and 6 % for C, N and S, respectively and the accuracy is within 7 % of the recommended values for all three elements. Carbonate carbon (Ccarb) was determined by coulometry (RSD = 1 %). Percent organic carbon (C ) was calculated by 151 127°W 126°W 125°W Fig. 5.1. The study area is located off the west coast of Vancouver Island, British Columbia, Canada (Inset). Piston Core JT96-02 was collected from the continental slope at a water depth of 1340 m. 152 difference (Corg = Ctot - Ccarb) and has a RSD of ~4 %. The isotopic composition of organic matter (8I3Corg and 815N) was determined by continuous-flow mass spectrometry. Samples for 813Corg analysis were first decarbonated with 10% HC1 and dried at 50°C overnight. Samples for 815N analysis were not pretreated. The reproducibility for isotopic data is ±0.1 %o for carbon and ±0.2 %o for nitrogen. Major and minor element concentrations were determined by X-ray fluorescence and have relative standard deviations of 5 % and 15 %, respectively. Following the standard practice these data are presented as element/Al ratios rather than concentrations, thus minimizing variability due to changes in dilution by non-lithogenic components such as organic matter and biogenic carbonate. The concentrations of Mo, Re, Ag, Cd and U were measured by isotope-dilution, inductively-coupled plasma mass spectrometry following the method described in Ivanochko (2001). The RSD is generally better than 10 % for all trace metals and the accuracy, measured for the National Research Council of Canada sediment standard MESS-1, is 8 % or better for Re, Mo and U, and -14 % for Cd. The accuracy for Ag could not be evaluated as there is no accepted Ag value for MESS-1 at present. 5.2 Geochemical variations The sulphide layers contain higher total sulphur (Stot) and higher organic carbon (Corg) than intervening sediments. The concentration of Stot within the sulphide layers ranges from 0.11 to 0.59 wt.% (avg. 0.25 wt.%) while the intervening sediment contains < 0.1 wt.% (Fig. 5.2a). The Corg content of the sulphide layers averages 0.49 wt.%, almost double the concentration in the clay (Fig. 5.2a). There is also a positive correlation between Stot and Corg (r2 = 0.64). The carbonate content is slightly higher within the sulphide layers (3.97 versus 3.26 wt.%) and progressively decreases upward before rising sharply at the base of the next sulphide layer (Fig. 5.2b). The 813Corg values are up to 1.7 %o higher within the sulphide layers (-23.4 vs -25.1 %0; Fig. 5.2c) as are 815N values (+5.8 vs +4.3 %0; Fig. 5.2c). The 153 combination of higher 513Corg and higher 815N values suggests the presence of slightly more marine organic matter within the sulphide-rich layers. The ratios of Si/Al, Zr/Al and Ti/Al are highest immediately overlying each sulphide layer and then decrease upwards (Fig. 5.3). The lowest ratios are observed within and just below the sulphide layers (Fig. 5.3). Relatively high Si/Al, Zr/Al and Ti/Al ratios suggest higher concentrations of quartz, zircon and titanium-bearing minerals which typically occur in coarser-grained sediments (Calvert, 1976). These data therefore indicate that an abrupt increase in grain-size occurs immediately above each sulphide layer and the deposits then fine upward, characteristics that suggest the presence of thin turbidites overlying the sulphide layers. Trace metal data, both as concentrations and metal/Al ratios, are plotted in Figure 5.4. Sulphide layers are significantly enriched in Mo (>2 pg/g) in comparison to the intervening sediments which contain only lithogenic amounts (Fig. 5.4a), and there is a strong positive correlation between Mo and Stot concentrations (r2 = 0.77; Fig. 5.5). Rhenium, and to a lesser extent Ag, are also enriched within the sulphide layers (Figs. 5.4b and c), but there is no correlation with Stot content (r2 = 0.26 and 0.41, Re and Ag respectively). Cadmium and U are present in lithogenic concentrations only, and there are no consistent downcore variations (Figs. 5.4d and e). 5.3 Origin of the sulphide layers The occurrence of sharp peaks in metal concentrations, such as those for Mo in Core JT96-02pc, are often interpreted as evidence of metal remobilization due to the influx of oxygen (i.e., burndown) and subsequent reprecipitation in underlying reduced sediments (Colley et al., 1984; Colodner et al., 1992; Thomson et al, 1993; Rosenthal et al., 1995; Thomson et al., 1995; Crusius and Thomson, 2000a and b). These studies have shown that, given time, certain solid-phase metals can be remobilized when exposed to oxidants (i.e., 154 5.2a) 5.2b) 5.2c) Fig. 5.2. Downcore profiles of a) total sulphur (Stot) and organic carbon (Corg) contents, b) carbonate content, and c) 8i3Corg and 8i5N values. Shaded zones indicate the location of sulphide layers. The errors (as RSD) are 6 %, 4 % and 1 % for Stot, Corg and carbonate respectively. The C- and N- isotopic data have errors of ±0.1 %o and ±0.2 %o respectively. 155 5.3a) 5.3b) Si/Al Zr/Al ( x 10"4 ) 3.4 3.5 3.6 3.7 3.8 3.9 8.5 95 10.5 11.5 12.5 0.055 0.06 0.065 0.07 Ti/Al Fig. 5.3. Downcore profiles ofthe a) Si/Al ratio (RSD = 6 %) and b) Zr/Al and Ti/Al ratios (RSD = 17 % for both ratios). Shaded areas indicate the location of sulphide layers. 156 Fig. 5.4. a to e) Downcore profiles of trace metal concentrations (symbols) and metal/Al ratios (thick black lines). Shaded areas indicate the location of sulphide layers. Background (i.e., lithogenic) concentrations of the various metals are marked by the dashed lines (see Table 5.1 for references), f) Downcore profile of the sedimentary Re/Mo ratio. The dashed line represents both the Re/Mo ratio of seawater and the estimated lithogenic Re/Mo ratio. The precision (as RSD) for all trace metal data is 10 % or better (i.e., at most 2x the symbol size). The detection limits are as follows: 0.5 ppm for Mo, <1 ppb for Re, 57 ppb for Ag, 0.1 ppm for Cd and 0.04 ppm for U. 157 158 Total Sulphur (wt.%) Fig. 5.5. Plot of Mo concentration versus total sulphur content in the sulphide layers (solid circles) and the intervening sediments (open circles). The regression equation shown was determined using data from both the sulphide layers and the intervening sediment. 159 oxygen and nitrate). Subsequent diffusion and "recapture" by adsorption and/or reprecipitation in underlying reduced sediments can create sharp peaks at, or just below, the deepened redox boundary, although for some metals the zone of enrichment may extend for decimetres below the redox boundary because the trapping mechanism is relatively slow (e.g., Re and U; Colodner et al., 1992; Thomson et al, 1993; Crusius and Thomson, 2000a). Post-depositional burndown can be rejected as a controlling influence on the trace metal distributions observed in Core JT96-02 for four major reasons. First, the average sedimentation rate in this section of the core is -60 cm/kyr. Such rapid deposition would work to inhibit oxygen penetration into the sediments. Furthermore, the presence of multiple organic-rich layers, each of which would have had a significant associated oxidant demand, would have limited 02 diffusion below any given layer. Second, the compositional characteristics (i.e., organic carbon and trace metal contents) of each sulphide layer are remarkably similar, yet it is highly unlikely that the degree of burndown would have been so consistent. Third, there is excellent correspondence between the major element compositional variations, inferred to reflect textural changes, and the distribution of trace metals. Burndown, if it had occurred, would have affected the latter without influencing the former, thus destroying or at least muting any relationship. Finally, all of the redox-sensitive trace metals should display similar trends (i.e., depletion within the clay and enrichment within the sulphide layers) if the result of burndown, but they do not. In fact, the highest concentrations of Cd and U occur within the clay and not in the sulphide layers. These four reasons strongly imply that the observed trace metal enrichments represent an early diagenetic signature and are not the result of subsequent oxidative burndown. The sulphide layers in Core JT96-02pc are geochemically similar to the Mediterranean sapropels described in Calvert and Fontugne (2001), although by definition these layers are not sapropels as they do not contain >2 wt.% organic carbon. Both sapropels and sulphide layers are finer-grained and organic-rich relative to intervening sediments, and both exhibit trace metal enrichments. In the case of the sapropels increased productivity in 160 surface waters resulted in higher organic matter flux to the sediment and degradation of this material led to anoxic conditions and thus trace metal enrichment (Calvert and Fontugne, 2001). The sulphide layers described here are characterized by higher organic carbon and carbonate contents which could be interpreted as representing episodic increases in the settling flux of biogenic detritus, set against a backdrop of glacial rock-flour deposition. However, evidence provided by the Si/Al, Zr/Al and Ti/Al ratios (Fig. 5.3) argues against this hypothesis. These data strongly imply that hemipelagic sedimentation represented by the sulphide-rich layers was episodically interrupted by turbidite deposition, most probably related to rapid deglaciation of the nearby continental shelf. This "pulsed turbidite" model provides an explanation for the S and Mo enrichments observed in the organic-rich layers. Geologically instantaneous emplacement of the turbidites would have rapidly attenuated oxygen influx, thus inducing anoxic conditions in the organic-rich sediments that immediately underlie each turbidite. The resulting sulphate reduction and precipitation of FeS would then have provided the conditions necessary for Mo accumulation. Pulsed, rapid sedimentation can adequately explain the observed Mo enrichment in these sediments. It is not necessary to invoke oxygen depletion of the bottom-water as having influenced the Mo content. This observation potentially complicates the use of sedimentary Mo concentrations as a proxy for bottom-water anoxia in environments where large changes in sedimentation rate occur. 5.4 Trace metal geochemistry While the rapid emplacement of turbidites on top of relatively organic-rich, hemipelagic deposits can explain the formation of the sulphide layers, it cannot explain why some redox-sensitive metals are enriched while others are apparently not. The precipitation of FeS clearly indicates that anoxic conditions developed. Under such conditions the accumulation of authigenic Mo, Re, U, Cd and Ag is to be expected, but only Mo and Re, 161 and to a lesser extent Ag, are significantly enriched in the sulphide layers. An explanation for the apparent contrasts among the redox-sensitive metals lies in how efficiently the metals are fixed under reducing conditions, as well as in the relative contributions of authigenic and lithogenic sources. The enrichment of Mo in anoxic sediments is directly linked to sulphate reduction and the formation Fe-sulphides. At an H2S concentration of ~11 pM the principal dissolved Mo species switches from Mo042" to MoS42". Thiomolybdate is highly particle reactive and readily adsorbes on to Fe-bearing particles such as Fe-sulphides (Helz et al., 1996; Erickson and Helz, 2000). The link between S and Mo in the sulphide layers is clearly indicated by their very strong positive correlation (r2 = 0.77; Fig. 5.5). The high Mo concentrations seen in the sulphide layers of Core JT96-02 is therefore explained by the downward diffusion of Mo from the overlying seawater, through the newly emplaced turbidite, to the underlying zone of sulphate reduction where it accumulated. In contrast to Mo, Re reduction and precipitation, possibly as Re02 or a Re-organic complex, commences when suboxic conditions are attained and continues during sulphate reduction (Crusius et al., 1996; Crusius and Thomson, 2000a). There is evidence to suggest that the precipitation of Re is kinetically slow (Crusius and Thomson, 2000a) and that it may be hindered by the presence of dissolved sulphide (Colodner et al., 1992). This could account for the numerous examples of preferential enrichment of Mo relative to Re in anoxic marine sediments (e.g., Crusius et al., 1996; Crusius et al., 1999; Crusius and Thomson, 2000a). However, within the sulphide layers of Core JT96-02 no such preferential enrichment is observed. The primary source for both authigenic Re and Mo in sediments is seawater and if these metals accumulate quantitatively, that is if their concentrations in the porewater are zero or nearly so at a common precipitation point, then the sedimentary Re/Mo ratio should be very similar to that in seawater (-0.7 ng/pg). Higher ratios would indicate that suboxic conditions prevailed under which Re but not Mo would accumulate, while lower ratios would result from preferential Mo accumulation under oxic or anoxic conditions 162 (Crusius et al., 1996). In Core JT96-02 Re/Mo ratios range from 0.4 to 3.0 ng/pg and average ~1 ng/pg in both the sulphide layers and intervening clay (Fig. 5.4f). In general these values suggest that: i) anoxic conditions developed rapidly within the organic-rich layers, before substantial Re enrichment could occur under suboxic conditions; and ii) suboxic conditions either did not develop, or did not persist, within the turbidites for long enough to allow preferential enrichment of Re. Furthermore, the similarity to the seawater Re/Mo ratio indicates that removal of Mo and Re from seawater into the sulphide layers was quantitative. The higher Re/Mo ratios near the base of each sulphide layer suggest that suboxic conditions prevailed. This is consistent with the low sulphur content, indicating that minimal sulphate reduction occurred, most probably due to the lower concentration of organic carbon. The precipitation of Cd and Ag in anoxic sediments does not involve a change in redox state. However, these metals are indirectly redox-sensitive because they precipitate as discrete sulphide minerals in the presence of trace amounts of dissolved sulphide (Koide et al, 1986; Rosenthal et al., 1995). Despite this, there is little if any observable Cd enrichment within the sulphide layers, and only a very slight Ag enrichment. The contrast between the distributions of this pair of metals and that of Mo reflects the relative importance of seawater and lithogenic metal sources. The contrast can be visualized by comparing the mean concentration of the metals in seawater to the content typical of lithogenic detritus (Table 5.1). The resulting unitless ratio emphasizes that Mo, unlike either Ag or Cd, is highly enriched in seawater relative to lithogenic materials. Since seawater is the primary source for the authigenic accumulation of metals, it follows that where geochemical conditions are appropriate (i.e., where sufficient H2S occurs in porewaters), Mo enrichment should be relatively more obvious than that of Ag and Cd. A similar reasoning applies to U. Like Re, U is reduced and will precipitate in suboxic and anoxic sediments (Klinkhammer and Palmer, 1991). However, the high lithogenic background concentration in JT96-02, and 163 Table 5.1. Seawater and lithogenic concentrations of the redox-sensitive trace metals. Metal Seawater concentration1 Lithogenic concentration SW: L ratio (SW, ug/g) (L, ug/g) Mo 10.266 (Collier, 1985) 0.6 (McKay, unpubl. data)2 17.1 Re 0.008 (Anbar etai., 1992) 0.0005 (Koide et al., 1986) 16.0 Ag 0.003 (Zhang et al., 2001) 0.1 (Bowen, 1966) <0.1 Cd 0.112 (Brulandetai, 1994) 0.2 (McKay, unpubl. data)2 0.6 U 3.240 (Chen etai, 1986) 21 !-6 1 Seawater data for the non-conservative elements Ag and Cd are for -1000 m water depth. 2 Data are for surface sediments at Site JT96-02 (see Chapter 4). Values are similar to those proposed for terrigenous sediments derived from Vancouver Island (Morford et al, 2001). 5 Lithogenic values for U vary greatly depending on the sediment composition and grain-size. Morford et al. (2001) suggest a lithogenic value of 1 ug/g which is lower than the average shale value of ~4 ug/g (Turekian and Wedepohl, 1961). We assume an intermediate value of 2 ng/g which is similar to the average value in the gray clay. 164 correspondingly low seawater to lithogenic ratio (Table 5.1), means that the authigenic U signal that must be present in these deposits is obscured. Finally, one other factor could influence the accumulation of certain trace metals in continental margin settings such as that described here. While Mo, Re and U have essentially conservative (i.e., invariant) concentrations in the open-ocean water column, both Ag and Cd distributions are strongly influenced by biological cycling as both elements are taken up by phytoplankton in surface waters. Dissolved Ag distributions tend to follow Si, implying that the element is incorporated into the "hard parts" and regenerated in relatively deep waters (Flegal et al., 1995; Zhang et al., 2001), while Cd is known to be incorporated into the "soft parts" and is regenerated at shallower depths (Boyle, 1988). Thus, the concentrations of Ag and Cd in the bottom water, and hence the seawater to lithogenic ratio, will vary with water depth. Furthermore, the biogenic flux of Cd and Ag to the sediment must be considered as contributions from this source could also obscure the authigenic signal. 5.5 Summary The relatively low organic carbon and carbonate contents, as well as lower 513Corg and 515N values that characterize the gray clay imply a more terrigenous provenance. The fining upward trend exhibited by each clay layer suggests that they are turbidites. Based on its texture and radiocarbon age, as well what is known about the deglacial history of western Canada, this material is most probably rock-flour generated as Cordilleran ice retreated from the nearby continental shelf. In contrast, the relatively organic-rich sediments that immediately underlie each turbidite layer have a more marine character as indicated by higher 513Corg and 815N values and most probably represent the background hemipelagic sedimentation. 165 The formation of sulphide-rich layers in Core JT96-02pc was the result of pulsed turbidite deposition that slowed the influx of oxygen into the underlying organic-rich sediments and thus allowed porewater anoxia to develop. This led to sulphate reduction and Fe-sulphide precipitation and in turn Mo enrichment. Concentration profiles for Re and Ag generally parallel that of Mo; however, Cd and U profiles do not. Contrasts among the profiles can be attributed to differences in the relative contributions of dissolved metals from the overlying bottom water (i.e., the authigenic flux) and background concentrations associated with the lithogenic flux. Where the seawater to lithogenic concentration ratio is relatively high (e.g., Mo and Re) the authigenic component can be readily identified, but where the ratio is low (e.g., Cd and U) the authigenic component may be obscured. In the latter instance observable enrichments would require time to develop. It follows that sediments characterized by a relatively high Mo content, but low concentrations of other redox-sensitive trace metals, and a Re/Mo ratio similar to that of seawater, must have experienced rapid onset of anoxic conditions and relatively rapid burial below the depth of active metal influx. The episodic deposition of turbidites set against a backdrop of hemipelagic sedimentation created the necessary conditions for the development of Mo concentration spikes. The enrichments thus occurred largely independently of changes in bottom water oxygen content and organic carbon flux to the sediment. This observation potentially complicates the use of sedimentary Mo concentration as a proxy for bottom-water anoxia in environments that are characterized by abrupt changes in sedimentation rate. It also emphasizes the importance of understanding the sedimentalogy if trace metal data are to be utilized as paleo-proxies. 166 5.6 References Adelson, J.M. Helz, G.R., Miller, C.V., 2001. Reconstructing the rise of recent coastal anoxia; molybdenum in Chesapeake Bay sediments. Geochimica et Cosmochimica Acta 65, 237-252. Anabar, A.D., Creaser, R.A., Papanastassiou, D.A., Wasserburg, G.J., 1992. Rhenium in seawater: Confirmation of generally conservative behavior. Geochimica et Cosmochimica Acta 56, 4099-4103. Blaise, B., Clague, J.J., Mathewes, R.W., 1990. Time of maximum Late Wisconsin glaciation, west coast of Canada. Quaternary Research 34, 282-295. Boyle, E.A., 1988. Cadmium: Chemical tracer of deepwater paleoceanography. Paleoceanography 4, 471-489. Bowen, H.J.M., 1966. Trace Metals in Biogeochemistry, Academic Press, New York. Bruland, K.W., Orians, K.J., Cowen, J.P., 1994. Reactive trace metals in the stratified central North Pacific. Geochimica et Cosmochimica Acta 58, 3171-3182. Calvert, S.E., 1976. The mineralogy and geochemistry of near-shore sediments. In: J.P. Riley, J.P., Chester, R. (Eds.), Chemical Oceanography, Vol. 6. Academic Press, London, 187-280pp. Calvert, S.E., 1990. Geochemistry and origin of the Holocene sapropel in the Black Sea. In: V. Ittekkot, V., Kempe, S., Michaelis, W., Spitzy, A. (Eds.), Facets of Modern Biogeochemistry. Springer-Verlag, Berlin, 326-352pp.. Calvert, S.E., Fontugne, M.R., 2001. On the late Pleistocene-Holocene sapropel record of climatic and oceanographic variability in the eastern Mediterranean. Paleoceanography 16, 78-94. Calvert, S.E., Pedersen, T.F., 1993. Geochemistry of Recent oxic and anoxic marine sediments: Implications for the geological record. Marine Geology 113, 67-88. Chen, J.H., Edwards, R.L., Wasserburg, G.J., 1986. 238U, 234U and 232Th in seawater. Earth and Planetary Science Letters 80, 241-251. Clague, J.J., James, T.S., 2002. History and isostatic effects of the last ice sheet in southern British Columbia. Quaternary Science Reviews 21, 71-87. Colley, S., Thomson, J., Wilson, T.R.S., Higgs, N.C, 1984. Post-depositional migration of elements during diagenesis in brown clay and turbidite sequences in the North East Atlantic. Geochimica et Cosmochimica Acta 48, 1223-1235. Collier, R.W., 1985, Molybdenum in the Northeast Pacific Ocean. Limnology and Oceanography 30, 1351-1354. Colodner, D.C, Boyle, E.A., Edmond, J.M., Thomson, J., 1992. Post-depositional mobility of platinum, iridium and rhenium in marine sediments. Nature 358, 402-404. 167 Crusius, J., Thomson, J., 2000a. Comparative behaviour of authigenic Re, U, and Mo during reoxidation and subsequent long-term burial in marine sediments. Geochimica et Cosmochimica Acta 64, 2233-2242. Crusius, J., Thomson, J., 2000b. Behaviour of Re and Ag on oxidation of SI, the most recent Mediterranean sapropel. AGU Fall Meeting, F614 (OS52B-10). Crusius, J., Calvert, S., Pedersen, T., Sage, D., 1996. Rhenium and molybdenum enrichments in sediments as indicators of oxic, suboxic and sulfidic conditions of deposition. Earth and Planetary Science Letters 145, 65-78. Crusius, J., Pedersen, T.F., Calvert, S.E., Cowie, G.L., Oba, T., 1999. A 36 kyr geochemical record from the Sea of Japan of organic matter flux variations and changes in intermediate water oyxgen concentrations. Paleoceaography 14, 248-259. Dean, W.E., Gardner, J.V., Piper, D.Z., 1997. Inorganic geochemical indicators of glacial-interglacial changes in productivity and anoxic on the California continental margin. Geochimica et Cosmochimica Acta 61, 4507-4518. Dean, W.E., Piper, D.Z., Peterson, L.C., 1999. Molybdenum accumulation in Cariaco basin sediment over the past 24 k.y.: A record of water-column anoxia and climate. Geology 27, 507-510. Erickson, B.E., Helz, G.R., 2000. Molybdenum(VI) speciation in sulfidic waters: Stability and lability of thiomolybdates. Geochimica et Cosmochimica Acta 64, 1149-1158. Flegal, A.R., Sanudo-Wilhelmy, S.A., Scelfo, G.M., 1995. Silver in the eastern Atlantic Ocean. Marine Chemistry 49, 315-320. Francois, R., 1988. A study on the regulation of the concentrations of some trace metals (Rb, Sr, Zn, Pb, Cu, V, Cr, Ni, Mn and Mo) in Saanich Inlet sediments, British Columbia, Canada. Marine Geology 83, 285-308. Helz, G.R., Miller, C.V., Charnock, J.M., Mosselmans, J.F.W., Pattrick, R.A.D., Garner, CD., Vaughaan, D.J., 1996. Mechanism of molybdenum removal from the sea and its concentration in black shales: EXAFS evidence. Geochimica et Cosmochimica Acta 60, 3631-3642. Ivanochko, T.S., 2001. Productivity influences on oxygenation of the Santa Barbara Basin, California during the Late Quaternary. M.Sc. Thesis, The University of British Columbia, 133pp. Josenhans, H.W., Fedje, D.W., Conway, K.W., Barrie, J.V., 1995. Post glacial sea levels on the Western Canadian continental shelf: Evidence for rapid change, extensive subaerial exposure, and early human habitation. Marine Geology 125, 73-94. Klinkhammer, G.P., Palmer, M.R., 1991. Uranium in the oceans: Where it goes and why. Geochimica et Cosmochimica Acta 55, 1799-1806. Koide, M., Hodge, V.F., Yang, J.S., Stallard, M., Goldberg, E.G., Calhoun, J., Bertine, K.K., 1986. Some comparative marine chemistries of rhenium, gold, silver and molybdenum. Applied Geochemistry 1, 705-714. 168 Morford, J.L., Russell, AD., Emerson, S., 2001. Trace metal evidence for changes in the redox environment associated with the transition from terrigenous clay to diatomaceous sediment, Saanich Inlet, BC. Marine Geology 174, 355-369. Nameroff, T.J., Balistrieri, L.S., Murray, J.W., 2002. Suboxic trace metal geochemistry in the eastern tropical North Pacific. Geochimica et Cosmochimica Acta 66, 1139-1158. Rosenthal, Y., Lam, P., Boyle, E.A., Thomson, J., 1995. Authigenic cadmium enrichments in suboxic sediments: Precipitation and postdepositional mobility. Earth and Planetary Science Letters 132, 99-111. Thomson, J., Higgs, N.C, Croudace, I.W., Colley, S., Hydes, D.J., 1993. Redox zonation of elements at an oxic/post-oxic boundary in deep-sea sediments. Geochimica et Cosmochimica Acta 57, 579-595. Thomson, J., Higgs, N.C, Wilson, T.R.S., Croudace, I.W., De Lange, G.J., van Santvoort, P.J.M., 1995. Redistribution and geochemical behaviour of redox-sensitive elements around SI, the most recent eastern Mediterranean sapropel. Geochimica et Cosmochimica Acta 59, 3487-3501. Turekian, K.K., Wedepohl, K.H., 1961. Distribution of the elements in some major units of the earth's crust. Geological Society of America Bulletin 72, 175-192. Zhang, Y., Amakawa, H., Nozaki, Y., 2001. Oceanic profiles of dissolved silver: Precise measurements in the basins of western North Pacific, Sea of Okhotsk and the Japan Sea. Marine Chemistry 75, 151-163. Zheng, Y., Anderson, R.F., van Geen, A., Kuwabara, J., 2000. Authigenic molybdenum formation in marine sediments: A link to pore water sulfide in the Santa Barbara Basin. Geochimica et Cosmochimica Acta 64, 4165-4178. 169 6. Summary 6.1 Sedimentation on the Vancouver Island Margin The last deglacial was a period of dramatic change off Vancouver Island. Glaciers, which at their maximum extent 15 to 14 14C kyr ago had extended on to the shelf, began rapidly to retreated (Barrie and Conway, 1999; Clague and James, 2002). At this time linear sedimentation rates on the mid-slope (920 m) increased from 47 cm/kyr to over 150 cm/kyr and remained high until -12.5 14C kyr when deglaciation of the shelf was essentially complete (Barrie and Conway, 1999) and the primary source of detritus (i.e., glacial outwash) was gone. Isostatic rebound caused a brief period of relative sea level fall during the deglaciation but then the sea began to rise rapidly reaching, and in many areas exceeding, modern sea level at -10 14C kyr. Flooding of the shelf and the resultant trapping of sediments within fjords caused a substantial decrease in sedimentation rates on the mid- and lower slope (to 5 to 20 cm/kyr). Holocene sedimentation rates on the shelf and upper slope are even lower (<2 cm/kyr) as a result of high wave and current energies (Bornhold and Yorath, 1984; Bornhold and Barrie, 1991). The exception is Station 01 (120 m water depth) which is located within a depression and may be a site of sediment focusing. 6.2 The sea-surface temperature record Alkenone palaeothermometry has yielded evidence of large and rapid shifts in sea-surface temperature (SST) during the deglacial (1°C in 40 to 80 years; Kienast and McKay, 2001). Based on radiocarbon dating and striking similarities to the 8180 temperature record of the GISP-2 ice core these SST fluctuations off Vancouver Island were essentially synchronous with changes in atmospheric temperature over Greenland (Kienast and McKay, 2001). Similar millennial-scale climate oscillations have been recognized in a number of palaeo-records from 170 the Northeast Pacific (e.g., Thunell and Mortyn, 1995; Hendy and Kennett, 1999; Mix et al., 1999) and Western North America (e.g. Benson et al., 1997; Grigg and Whitlock, 2002). One of the most widely recognized events, the "early Holocene" thermal maximum, is clearly evident in Core JT96-09 between 11 and 10 calendar kyr B.P. as it is in many other palaeo-records from the region (e.g., Sabin and Pisias, 1996; Doose et al., 1997; Pellatt and Mathewes, 1997; Heusser et al., 1998). The palaeotemperature profile from Core JT96-09 (Kienast and McKay, 2001) clearly defines warm and cold events that are inferred to represent the Boiling and Allerod warm periods (-14.3 to 13.5 kyr B.P. and 13.5 to 12.7 kyr B.P., respectively) and Younger Dryas (-12.7 to 11.0 kyr B.P.) cold event. This observation has allowed the palaeoceanographic data presented in this thesis to be placed in a more global context. 6.3 The marine productivity record One of the primary objectives of this thesis was to determine how marine primary production along the continental margin of the northeastern Pacific has varied on glacial-interglacial and shorter timescales. The organic carbon content and/or mass accumulation rate are commonly used as proxies for palaeoproductivity but in continental margin sediments a significant fraction of the total organic matter may be of terrestrial origin. Radiocarbon dating of bulk organic carbon, isotopic analysis (813C and 815N) and organic geochemistry were thus used to quantify the abundance of terrestrial organic matter in Core JT96-09. This information was used to partition the total sedimentary organic carbon content into marine and terrestrial fractions, which permitted the marine organic carbon flux to the sediment to be computed. As much as 70 % of the organic carbon found in late glacial and early deglacial (i.e., Boiling) sediments is terrestrial in origin. Prior to 14.3 calendar kyr B.P. marine organic carbon accumulation was generally low, a reflection of low primary production and resulting low export production, because glacial-mode atmospheric circulation (i.e., northerly winds driven by the Aleutian Low) did not favour coastal upwelling. In this respect the northern 171 California Current System (CCS) was similar to the central CCS, where a marked decrease in productivity was documented during the last glacial (Dymond et al., 1992; Lyle et al., 1992; Ortiz et al., 1997; Dean and Gardner, 1998; Mix et al., 1999; Kienast et al., 2002). At the start of the deglaciation (i.e., the Boiling) the accumulation of marine organic matter began to increase. However, this material was diluted by abundant terrestrial organic detritus generated as glaciers retreated from the continental shelf. At the start of the Allerod (-13.5 calendar kyr B.P.) terrestrial input dropped abruptly and marine organic matter accumulation increased 6-fold relative to glacial values. Other palaeoproductivity proxies (i.e., % biogenic Ba, % opal and alkenone abundances) also indicate that export production was high throughout this period. Increased marine productivity during the deglacial is also documented in palaeo-records from the central and southern CCS (Lyle et al., 1992; Gardner et al., 1997; Dean and Gardner, 1998; Mix et al., 1999; Kienast et al., 2002), as well as the Gulf of Alaska (de Vernal and Pedersen, 1997) and western Pacific (Keigwin et al., 1992). Given that this marine productivity pulse was a wide spread phenomenon it is possible that it had an impact on atmospheric CO2 concentrations. However, since increased productivity was probably the direct result of the initiation and/or intensification of upwelling, any sequestration of atmospheric C02 via the "biological pump" may have been countered by the release of C02 as upwelled water warmed. 13 A more detailed investigation of 8 Corg records from throughout the region may help to clarify how C02 fluxes changed. There was a brief return to glacial-like conditions (i.e., relatively low productivity) during the Younger Dryas and a corresponding decrease in marine organic matter accumulation. In the Holocene, the accumulation of marine organic matter was, and still is, lower than during the Allerod. This is largely the result of low organic matter preservation due to low sedimentation rates (i.e., long oxygen exposure times) and intense biological recycling. Since primary production off Vancouver Island is controlled by upwelling the data collectively suggest that in the last 16 kyr upwelling was most intense during the Allerod, an interpretation that is supported by a slight decrease in the benthic-planktonic age difference. 172 6.4 Redox-sensitive trace metals and their palaeo-applications Temporal changes in organic carbon flux to the sediment had a direct impact on sedimentary redox conditions and may also have influenced OMZ intensity. Such changes can be inferred from redox-sensitive trace metal data. At present sediments on the Vancouver Island Margin become suboxic within millimetres of the sediment-water interface but there is no evidence that fully anoxic conditions develop within the upper 50 cm (i.e., no Mo enrichment). This observation was unexpected given the highly productive surface waters and relatively low bottom water oxygen concentrations in the region. Suboxic conditions are maintained because low sedimentation rates and extensive bioturbation support a small, but continual, influx of oxygen and other oxidants from the overlying water column. Low sedimentation also favours substantial accumulation of Re decimetres below the sediment-water interface. Unfortunately, such diagenetic enrichment has overprinted the early to mid-Holocene palaeo-record. Where sedimentation rates are relatively high, as they are at station 01 (~ 40 cm/kyr) for example, trace metal concentrations are more likely to reflect the early diagenetic conditions that are in turn influenced by the depositional environment. It is thus assumed in this thesis that trace metal data in rapidly deposited glacial and deglacial sediments can be used as palaeo-environmental proxies. There is however one caveat. Sedimentation must be steady. In Core JT96-02 deglacial sedimentation was characterized by the episodic emplacement of centimetre-scale turbidites. Pulsed sedimentation led to the development of anoxic conditions by substantially reducing the rate of oxygen influx into relatively organic-rich sediments. The resulting Re and Mo enrichments occurred without any decrease in bottom water oxygen concentration and/or increase in organic carbon flux to the sediment, two factors that commonly lead to trace metal enrichment. In Core JT96-09 there is no evidence of abrupt changes in sedimentation during the deglacial with the exception of one turbidite deposited at ~13.5 calendar kyr. B.P. (i.e., the start of the Allerod). 173 The enrichment of redox sensitive trace metals during the Allerod, in particular the high Mo concentrations, indicate that anoxic conditions existed in near-surface sediments at this time. Benthic foraminifera species data (i.e., the dominance of Bolivina spp.) corroborate the trace metal results. Regardless of the exact cause (i.e., increased carbon flux and/or decreased ventilation) the development of anoxic conditions must have been associated with intensification of the oxygen minimum zone (OMZ). 6.5 Fluctuation in OMZ intensity Trace metal and benthic foraminifera species data suggest that the OMZ off Vancouver Island was more intense (i.e., more oxygen-depleted) between 13.5 to 12.7 calendar kyr B.P. (i.e., the Allerod) and possibly again between 11 and 8 kyr B.P. Radiocarbon dating of benthic-planktonic foraminiferal pairs indicate that by -16 kyr B.P. ventilation of the intermediate water mass (920 m water depth) off Vancouver Island was similar to that at present (~ 900 years). There is no evidence indicating that ventilation decreased during the Allerod (i.e., benthic-planktonic age differences did not increase) but there is substantial evidence for an increased biogenic flux to the sediment. Increased export production, most probably the direct result of increased primary production in surface waters, was the principal cause of OMZ intensification off Vancouver Island. The OMZ along the entire margin of the northeastern Pacific from Mexico to Vancouver Island intensified during the deglacial; however, off Vancouver Island intensification appears to lag -1500 years behind that off California. If OMZ intensification was controlled primarily by changes in ventilation of NPIW the entire northeastern margin of the Pacific should have responded at approximately the same time. The observed lag suggests that other factors, possibly regional changes in productivity and/or the transportation of oxygen-depleted water from south by the California Undercurrent, controls OMZ intensity. 174 6.6 Biogenic Ag? Intriguing results were obtained from the measurement of sedimentary Ag concentrations in near-surface sediments. Unlike the other redox-sensitive trace metals, for example Cd that has a similar geochemical behaviour, the primary control on Ag concentrations is not sedimentary redox conditions. Rather, it seems that Ag accumulates in, and is transferred to the sediment by, settling organic particles in much the same way that Ba is, as argued in Chapter 4. The measurement of sedimentary Ag concentrations may thus provide another palaeo productivity proxy. The potential of such a proxy is clearly seen when one looks at the Holocene record of Core JT96-09. Silver concentrations are consistently high throughout the Holocene while the other redox-sensitive trace metals, for example Re, are not (Fig. 6.1). As a preliminary inference, high Ag concentrations may reflect the relatively high primary productivity that has existed throughout the Holocene but which is not preserved in the sediments due to the intense degradation of organic matter in these slowly accumulating deposits. A number of major questions remain to be answered if Ag is to be used as a palaeoproductivity proxy. First, the hypothesis that Ag precipitates as Ag2S within anoxic microenvironments in setting organic particles must be tested. Second, the conditions under which this "biogenic" Ag is preserved need to be clarified. The work presented in this thesis has shown that sediments need only be weakly suboxic to preserve the signal, but the question remains about its preservation in oxic sediments. Third, the relative importance of the "direct biogenic" flux (i.e., Ag directly incorporated by organisms) relative to the "indirect biogenic" flux (i.e., Ag which accumulates as organic particles settle through the water column) must be evaluated. 6.7 Future work 175 Metal Concentrations (Ag - ng/g; Re - ng/g x 10) Fig. 6.1. A comparison of Re and Ag concentrations over the last 16 kyr B.P. in Core JT96-09. Note that in the weakly suboxic Late Holocene sediments the concentration of Re is low yet the Ag concentration is relatively suggesting that something other than redox is influencing Ag accumulation. 176 This thesis offers the first comprehensive palaeoceanographic investigation of the northeastern Pacific Ocean off the west coast of Vancouver Island. It provides a record for the past 16 kyr and in particular a very detailed record for the last deglaciation. Results indicate that in many respects the northern CCS behaved in a similar manner as the central and southern CCS. For example, both primary productivity and the accumulation of marine organic matter were substantially reduced in most regions during the deglacial and enhanced during the deglacial. However, the relatively short record present here must be extended back through Oxygen Isotope Stages 3 and 4 and if possible into OIS 5. Of particular interest is whether or not the Dansgaard-Oeschger events that are recognized in southern records (e.g., Santa Barbara Basin - Behl and Kennett, 1996; Hendy and Kennett, 1999) are evident in the sedimentary record off Vancouver Island. 177 6.7 References Barrie, J. V., Conway, K.W., 1999. Late Quaternary glaciation and postglacial stratigraphy of the Northern Pacific Margin of Canada. Quaternary Research 51, p. 113-123. Behl, R.J., Kennett, J.P., 1996. Brief interstadial events in the Santa Barbara basin, NE Pacific, during the past 60 kyr. Nature 379, 243-246. Benson, L., Burdett, J., Lund, S., Kashgarian, M., Mensing, S., 1997. Nearly synchronous climate change in the Northern Hemisphere during the last deglacial termination. Nature 388, 263-265. Bornhold, B.D., Barrie, J.V., 1991. Surficial sediments on the western Canadian continental shelf. Continental Shelf Research 11, 685-699. Bornhold, B.D., Yorath, C.J., 1984. Surficial geology of the continental shelf, northwestern Vancouver Island. Marine Geology 57, 89-112. Clague, J.J., James, T.S., 2002. History and isostatic effects of the last ice sheet in southern British Columbia. Quaternary Science Reviews 21, 71-87. de Vernal, A., Pedersen T.F., 1997. Micropaleontology and palynology of core PAR87A-10: A 23,000 year record of paleoenvironmental changes in the Gulf of Alaska, northeast North Pacific. Paleoceanography 12, 821-830. Dean, W.E., Gardner, J.V., 1998. Pleistocene to Holocene contrasts in organic matter production and preservation on the California continental margin. GSA Bulletin 110, 888-899. Dymond, J., Suess, E., Lyle, M., 1992. Barium in deep-sea sediment: A geochemical proxy for paleoproductivity. Paleoceanography 7,163-181. Gardner, J.V., Dean, W.E., Dartnell, P., 1997. Biogenic sedimentation beneath the California Current system for the past 30 kyr and its paleoceanographic significance. Paleoceanography 12, 207-225. Grigg, L.D., Whitlock, C, 2002. Patterns and causes of millennial-scale climate change in the Pacific Northwest during Marine Isotope Stages 2 and 3. Quaternary Science Reviews 21, 2067-2083. Hendy, I., Kennett, J.P., 1999. Latest Quatnary North Pacific surface-water responses imply atmosphere-driven climate instability. Geology 27,291-294. Keigwin, L.D., Jones, G.D., Froelich, P.N., 1992. A 15,000 year paleoenvironmental record fromMeiji Seamount, far northwestern Pacific. Earth and Planetary Science Letters 111, 425-440. Kienast, S.S., McKay, J.L., 2001. Sea surface temperatures in the subarctic Northeast Pacific reflect millennial-scale climate oscillations during the last 16 kyrs. Geophysical Research Letters 28,1563-1566. Kienast, S.S., Calvert, S.E., Collier, R.W., Pedersen, T.F., 2002. Nitrogen isotope and productivity variations along the North East Pacific margin over the last 120 kyr: Surface and subsurface palaeoceanography. Paleoceanography 17, doi: 10.1029/200IPA000650. 178 Lyle, M., Zahn, R., Prahl, F., Dymond, J., CoUer, R., Pisias, N., Suess, E., 1992. Paleoproductivity and carbon burial across the California Current: The mulitracers transect, 42°N. Paleoceanography 7, 251-272. Mix, A.C., Lund, D.C., Pisias, N.G., Boden, P., Bornmalm, L., Lyle, M., Pike, J., 1999. Rapid climate oscillations in the northeast Pacific during the last deglaciation reflect northern and southern hemisphere sources in Mechanisms of Global Climate Change at Millennian Time Scales, Geophysical Monograph 112 edited by P.U. Clark, R.S. Webb and L.D. Keigwin. American Geophysical Union, pp. 127-148. Ortiz, J., Mix, A., Hostetler, S., Kashgarian, M., 1997. The California Current of the last glacial maximum: Reconstruction at 42°N based on multiple proxies. Paleoceanography 12,191-205. Thunell, R.C., Mortyn, P.G., 1995. Glacial climate instability in the Northeast Pacific Ocean. Nature 376, 504-506. 179 Appendix I. Methods a) Total carbon, total nitrogen and total sulphur were measured by high temperature combustion using a Carlo Erba NA1500 elemental analyzer. Two marine sediment standards from the National Research Council of Canada (NRC standards PACS-1 and MESS-1) were analyzed with each batch of samples and yielded relative standard deviations (RSD, la) of 3 %, 5 % and 6 % for C, N and S, respectively. The accuracy was within 7 % of the recommended values for all three elements. b) Percent carbonate carbon (i.e., inorganic carbon) was measured using a Coulometrics 5010 coulometer. Approximately 50 mg of sample was reacted with 10% HC1. The C02 gas produced was bubbled into a dark blue solution of ethanolamine causing the pH to decrease and the blue colour to fade. The transmittance of the solution was continually measured and when it began to increase OH" ions were generated electrically. The amount of C02 produced, and thus amount of CaC02 in the sample, was directly proportional to the current required to return the ethanol amine solution to its original colour. Results for a calcium carbonate standard run multiple times with each batch of samples yielded an RSD of ~1 % (i.e., 11.93 ±0.13, n=141). c) The amount of organic carbon in samples was calculated by difference (organic carbon = total carbon - inorganic carbon) and the aggregate RSD for these data is ~4 %. d) Biogenic silica (i.e., opal) content was determined by alkaline dissolution following the procedure of Mortlock and Froelich (1989). The concentration of Si was measured by spectrophotometry and then converted to weight % opal by multiplying by 2.4 (this assumes 180 10% water in opal; Mortlock and Froelich, 1989). The RSD (la) for two in-house standards (SNB and JV5) and repeat samples was <4 %. e) Organic geochemical analysis to determine the concentrations of C37 alkenones and C29 n-alkanes was carried out using the solvent extraction method of Villanueva et al. (1997). The organic fraction was extracted from 3 to 10 mg of freeze-dried sediment using dichloromethane (CH2C1) followed by saponification with a mixture of 6 % KOH in methanol. The organic fraction was then extracted with n-hexane and run through a silica column. The fraction containing the n-alkanes and C37 alkenones was eluted from the column with a mixture of dichloromethane and n-hexane. Analyses were made by manual-injection on an HP 5880 gas chromatograph. The n-alkanes and C37 alkenones were identified using an internal standard (i.e., mixture of C19, C36 and C40), and the identification of C37 alkenones was confirmed by GC-MS. The concentrations of C29 n-alkanes and C37 alkenones were calculated by assuming that their concentrations were proportional to the chromatogram peak area and that the response factor was the same as for the internal standards. Only the di- and tri-unsaturated C37 alkenones were quantified, as the amount of tetra-unsaturated C37 alkenone was generally below the detection limit of 10 ng. f) The carbon- and nitrogen-isotopic composition of organic matter were obtained by continuous-flow mass spectrometry using a Fisons NA1500 elemental analyzer attached to a VG Prism mass spectrometer. Samples for carbon isotopic analysis of organic matter (813C0rg) were pretreated with 10 % HC1 to remove carbonate material and then dried at 50°C overnight. These samples were not washed with distilled water prior to drying. Nitrogen isotope results (815N) were obtained for untreated bulk sediment samples. Data are reported in the standard 8-notation relative to Vienna Peedee belemnite (VPDB) for carbon and atmospheric N2 for nitrogen. The isotopic results for an in-house sediment standard are ±0.1 181 %o for carbon and ±0.2 %o for nitrogen. Repeat analyses of samples are generally better than this. g) Prior to isotopic analysis benthic and planktonic foraminifera were sonicated in methanol and roasted under vacuum for 30 minutes at 430°C. The samples were then reacted in a common orthophosphoric acid bath at 90°C and analyzed using a VG Prism mass spectrometer in dual inlet mode. Data are reported in the standard 5-notation relative to VPDB. The reproducibility of three in-house standards (Mexical, UQ6 and a foraminiferal standard), that were calibrated using the international standard NBS-19, was ±0.15 %o for 5180 and ±0.07 for 513C. h) Major and minor element concentrations were measured by X-ray fluorescence (XRF) on a Philips PW2400 wavelength-dispersive sequential automatic spectrometer according to the method of Calvert (1990). Samples for major element analysis were combined with Spectroflux 105 in a 9:1 ratio and melted at 1100°C. To correct for the loss of volatiles during melting Spectroflux 100 was added to bring the samples back to their original weight. The samples were then re-melted and cast into glass disks. Minor elements were measured on pressed pellets backed with borax. The precision of the XRF method is estimated to be 5 % and 15 % for major and minor elements, respectively. Unless otherwise stated, all major and minor element data were corrected for the presence of sea salt in the dried sediment. The salt content was determined by measuring the chlorinity by titration with AgN03. A salinity of 35 %o was assumed. i) The concentration of biogenic barium was calculated using the equation of Dymond et al. (1992): Bio-barium = Total Ba - (Al x Ba/Allith) 182 A Ba/Al lithogenic ratio (Ba/Aliith) of 0.0027 was estimated from the exponential regression of the Ba/Al ratios of surface sediments versus water depth following the method of Klump et al. (2000). This method assumes that the fraction of biogenic Ba increases seaward and that close to land (i.e., ~0 m water depth) all of the Ba is terrigenous in origin. The Ba/Aliith estimated using the regression method was confirmed by chemically extracting the bio-barium using a 2M solution of NH4CI (Schenau et al., 2001) and then measuring the barium content of the residue (i.e., the lithogenic Ba). j) Trace metal concentrations (i.e., Re, Ag, Cd, Mo and U) were measured by isotope-dilution inductively-coupled plasma mass spectrometery. Sample preparation involved adding known amounts of isotopically-enriched spike solutions to ~20 mg of powdered sediment. Samples were then microwave digested in a mixture of concentrated HNO3, HC1 and HF. The digests were evaporated on a hotplate overnight and then re-digested in 5N HC1. Aliquots were taken for Mo and U analysis and the remaining sample was run through an anion exchange column (Dowex 1-X8 resin) to remove Zr and Nb that form compounds (i.e., ZrO, ZrOH, NbO) which interfere with the analysis of Ag, and Mo that forms MoO" which interferes with Cd analysis. To check precision a University of British Columbia sediment standard (SNB) was analyzed with each batch of samples. The resulting RSD (la) for Re, U, Mo, Cd and Ag analyses were 11 %, 9 %, 7 %, 10 % and 8 %, respectively. Accuracy, which was assessed by measuring the concentrations of these metals in the NRC sediment standard MESS-1, was 8 % or better for Re, Mo and U, and -14 % for Cd. The detection limits, calculated as 3x the relative standard deviation of the blank, were estimated to be <1 ppb for Re, 0.04 ppm for U, 0.5 ppm for Mo, 0.1 ppm for Cd and 57 ppb for Ag (data provided by K. Gordon). 183 References Calvert, S.E., 1990. Geochemistry and origin of the Holocene sapropel in the Black Sea. In: Ittekkot, V., Kemp, S., Michaelis, W., Spitzy, A. (Eds.), Facets of Modern Biogeochemistry. Springer-Verlag, Berlin, 326-352pp. Dymond, J., Suess, E., Lyle, M., 1992. Barium in deep-sea sediments: A geochemical proxy of paleoproductivity. Paleoceanography 7, 163-181. Klump, J., Hebbeln, D., Wefer, G., 2000. The impact of sediment provenance on barium-based productivity estimates. Marine Geology 169, 259-271. Mortlock, R.A., Froelich, P.N., 1989. A simple method for the rapid determination of biogenic opal in pelagic marine sediments. Deep Sea Research 36,1415-1426. Schenau, S.J., Prins, M.A., De Lange, G.J., Monnin, C, 2001. Barium accumulation in the Arabian Sea: Controls on barite preservation in marine sediemnts. Geochimica et Cosmochimica Acta 65, 1545-1556. Villanueva, J., Pelejero, C, Grimalt, J.O., 1997. Clean-up procedures for the unbiased estimation of C37 alkenone sea surface temperatures and terrigenous n-alkane inputs in paleoceanography. Journal of Chromatography 757, 145-151. 184 J5 2 c E O E V 00 00 00 <N •S e ^ s =§ & * 6 CL o O S-185 Table A2. Core Descriptions. Core' Latitude Longitude Water Depth (m) Bottom Water 02(ml/l) Depth in Core (cm)2 Core Description JT96-01mc 48° 45.95' N 125° 29.57' W 120 2.4 0-38 olive green mud (38 cm) JT96-04mc 49° 00.71'N 126° 49.82'W 407 1.0 0-9 9-19 olive green muddy sand (9 cm) gray clay (10 cm) JT96-06bc 48° 58.73'N 126° 52.68' W 720 0.4 0-18 18-35 olive green sandy mud (18 cm) gray clay (17 cm) JT96-09mc 48° 54.76' N 126° 53.44' W 920 0.3 0-40 olive green mud (40 cm) JT96-09pc II 0-0.2 0.2-51 51 - 136 136-152 152-320 320 - 360 360 - 374 brownish "fluff layer (0.2 cm) olive green mud (50.8 cm) gray green clay (85 cm) sandy mud (16 cm) gray clay (168 cm) gray clay interbedded with mm-scale sandy layers (40 cm) gray clay (14 cm) JT96-02mc 49° 12.81'N 127° 18.57'W 1340 0.4 3 0-18 olive green mud (18 cm) JT96-02pc II II 0-16 16-34 34-90 90-211 olive green mud (16 cm) greenish gray clay (18 cm) interbedded sand and clay with erosional base (56 cm) gray clay with abundant sulphide-rich layers (121 cm) JT96-05bc 49° 07.91'N 127° 33.12'W 1750 1.2 0-1 1 -48 brownish "fluff layer (1 cm) olive green mud (47 cm) ' mc = multicores; be = boxcores; pc = piston cores. 2 uncorrected core depth. 'dissolved Oj concentration was measured at 1240 m. 186 Table A3. CTD Data for Stations JT96-01,2,4,6 and 9. CTD Data for Station JT96-01 Depth Temperature Salinity Sigma-t (m) (°C) (psu) 0.0. 14.21 1.8 14.27 2.0 14.27 3.2 14.27 4.0 14.26 5.0 14.26 6.0 14.30 7.6 14.01 8.0 14.11 9.4 12.87 10.0 11.31 11.8 10.94 12.0 11.05 13.8 11.07 14.0 11.07 15.0 10.56 16.2 10.18 17.0 10.11 18.6 10.04 19.0 10.02 20.0 9.97 21.0 9.93 22.0 9.85 23.2 9.85 25.0 9.80 26.0 9.78 27.0 9.75 28.2 9.66 29.0 9.63 30.0 9.58 31.8 9.55 32.0 9.53 33.0 9.50 34.0 9.49 35.2 9.45 37.4 9.38 38.2 9.36 39.0 9.12 40.0 9.04 41.8 8.95 42.0 8.96 43.0 8.95 44.0 8.99 45.0 8.90 46.6 8.84 47.0 8.83 48.0 8.83 49.0 8.80 50.4 8.74 55.4 8.36 60.0 8.19 65.0 8.12 70.0 8.07 75.0 8.06 81.0 7.99 86.0 7.97 91.0 7.94 96.4 7.87 101.0 7.83 106.0 7.82 112.0 7.82 31.18 23.19 31.19 23.19 31.20 23.20 31.20 23.20 31.19 23.19 31.21 23.20 31.20 23.19 31.14 23.21 31.03 23.10 30.58 23.00 31.00 23.61 31.05 23.71 31.07 23.71 31.26 23.85 31.27 23.86 31.34 24.01 31.44 24.15 31.41 24.13 31.43 24.16 31.43 24.16 31.47 24.20 31.53 24.26 31.63 24.35 31.62 24.34 31.65 24.37 31.65 24.38 31.67 24.39 31.73 24.46 31.76 24.48 31.79 24.52 31.82 24.54 31.82 24.55 31.83 24.56 31.83 24.56 31.85 24.58 31.92 24.65 31.93 24.66 32.16 24.88 32.18 24.91 32.25 24.97 32.26 24.98 32.30 25.01 32.34 25.04 32.39 25.09 32.43 25.13 32.44 25.14 32.43 25.14 32.46 25.16 32.49 25.20 32.63 25.36 32.78 25.50 32.85 25.56 32.90 25.61 32.91 25.63 32.97 25.69 32.99 25.70 33.02 25.73 33.08 25.79 33.11 25.82 33.12 25.82 33.08 25.79 Table A3, (continued) CTD Data for Station JT96-02 Depth Temperature Salinity Sigma-(m) (°C) (psu) 0.0 13.54 31.63 23.68 1.0 13.54 31.63 23.68 2.0 13.54 31.63 23.68 3.0 13.54 31.63 23.68 5.2 13.54 31.63 23.68 6.0 13.54 31.63 23.68 8.0 13.55 31.63 23.68 9.0 13.52 31.63 23.68 10.8 13.50 31.63 23.69 11.0 13.51 31.63 23.68 12.0 13.50 31.63 23.68 13.0 13.49 31.63 23.69 14.0 13.45 31.64 23.70 15.0 13.39 31.63 23.71 17.2 12.16 31.75 24.04 20.0 11.33 31.74 24.18 21.0 11.21 31.76 24.22 22.0 11.05 31.78 24.26 23.0 10.96 31.78 24.28 24.0 10.81 31.90 24.40 25.0 10.48 31.88 24.44 26.2 10.19 31.90 24.50 27.0 10.18 31.90 24.50 28.0 10.17 31.90 24.50 29.0 10.16 31.93 24.53 30.0 10.17 31.91 24.51 31.0 10.19 31.92 24.52 34.4 10.38 32.04 24.58 35.0 10.40 32.05 24.59 36.6 10.42 32.07 24.60 37.0 10.43 32.07 24.60 39.2 10.47 32.10 24.61 40.0 10.47 32.12 24.62 41.0 10.46 32.12 24.63 42.0 10.52 32.19 24.67 43.0 10.64 32.29 24.73 44.6 10.06 32.48 24.97 45.0 9.96 32.33 24.88 46.6 9.93 32.18 24.76 47.0 9.75 32.28 24.87 48.6 9.67 32.21 24.83 49.0 9.64 32.27 24.88 50.6 9.52 32.20 24.85 55.0 9.13 32.33 25.01 60.8 8.97 32.38 25.07 65.6 8.91 32.49 25.17 71.0 8.73 32.56 25.25 77.0 8.57 32.63 25.33 82.0 8.41 32.68 25.39 88.0 8.39 32.78 25.48 93.4 8.32 32.89 25.57 98.0 8.17 32.94 25.63 104.6 8.06 32.99 25.69 109.0 8.04 33.03 25.72 114.2 8.02 33.08 25.77 119.0 7.95 33.20 25.87 124.2 7.92 33.22 25.89 129.0 7.84 33.31 25.97 134.2 7.82 33.38 26.03 139.0 7.82 33.36 26.01 146.2 7.82 33.22 25.90 151.0 7.82 33.19 25.88 156.6 7.82 33.13 25.83 161.0 7.82 33.05 25.77 166.0 7.82 32.92 25.66 171.8 7.82 32.80 25.57 176.0 7.82 32.74 25.53 181.0 7.82 32.66 25.47 187.2 7.82 32.56 25.38 192.0 7.82 32.48 25.32 197.0 7.82 32.41 25.27 202.8 7.82 32.24 25.14 207.0 7.82 32.18 25.08 212.0 7.82 32.08 25.01 Depth Temperature Salinity Sigma-(m) (°C) (psu) 218.6 7.82 32.00 24.94 223.0 7.82 31.95 24.90 228.0 7.82 31.87 24.85 234.0 7.82 31.82 24.80 239.0 7.82 31.81 24.80 245.2 7.82 31.78 24.77 250.0 7.82 31.73 24.73 257.4 7.82 31.70 24.71 262.0 7.82 31.68 24.70 268.6 7.82 31.58 24.61 273.2 7.82 31.55 24.59 278.0 7.82 31.52 24.57 283.0 7.82 31.50 24.55 288.6 7.82 31.50 24.55 293.0 5.16 34.04 26.89 298.0 5.15 34.04 26.90 304.2 5.10 34.04 26.90 309.0 5.06 34.03 26.90 314.0 5.05 34.04 26.91 319.0 5.03 34.05 26.92 324.2 5.02 34.04 26.91 329.0 4.98 34.04 26.92 334.6 4.94 34.05 26.93 339.0 4.92 34.06 26.94 344.0 4.92 34.06 26.94 350.4 4.89 34.06 26.94 355.0 4.85 34.06 26.95 360.0 4.84 34.07 26.95 365.8 4.82 34.07 26.96 370.0 4.82 34.07 26.96 375.0 4.82 34.07 26.96 382.0 4.80 34.07 26.96 387.0 4.79 34.06 26.96 394.0 4.76 34.07 26.97 399.0 4.71 34.07 26.97 404.0 4.73 34.08 26.97 409.0 4.74 34.09 26.98 415.2 4.74 34.09 26.99 420.0 4.72 34.09 26.99 425.0 4.70 34.09 26.99 431.0 4.70 34.10 27.00 436.0 4.68 34.10 27.00 441.0 4.67 34.11 27.00 447.8 4.66 34.11 27.01 452.0 4.66 34.11 27.01 457.0 4.64 34.11 27.01 462.8 4.64 34.13 27.02 467.4 4.62 34.13 27.03 472.0 4.61 34.13 27.03 477.0 4.60 34.14 27.04 483.4 4.58 34.14 27.04 488.2 4.54 34.15 27.05 493.0 4.53 34.15 27.05 499.4 4.53 34.15 27.06 504.0 4.50 34.16 27.06 509.0 4.49 34.16 27.07 515.6 4.46 34.17 27.08 520.0 4.45 34.17 27.08 525.0 4.44 34.17 27.08 530.0 4.42 34.17 27.09 535.4 4.41 34.18 27.09 540.0 4.39 34.17 27.09 545.0 4.38 34.18 27.09 551.8 4.36 34.18 27.09 556.0 4.34 34.18 27.10 561.0 4.32 34.18 27.10 566.0 4.30 34.18 27.10 572.0 4.29 34.18 27.11 577.0 4.28 34.19 27.11 582.0 4.28 34.19 27.11 587.0 4.26 34.19 27.12 592.0 4.25 34.19 27.12 597.0 4.17 34.20 27.13 602.0 4.13 34.20 27.13 Depth (m) Temperature (°C) Salinity (psu) Sigma-608.0 4.12 34.20 27.14 613.0 4.10 34.20 27.14 618.0 4.09 34.20 27.14 624.2 4.05 34.20 27.14 629.0 4.03 34.20 27.14 636.0 3.97 34.20 27.15 641.0 3.98 34.20 27.16 648.4 3.99 34.21 27.16 653.0 3.98 34.22 27.17 659.2 3.99 34.22 27.17 664.0 3.99 34.22 27.17 671.4 4.03 34.24 27.18 676.0 4.03 34.24 27.18 681.0 4.03 34.24 27.18 686.2 4.03 34.25 27.19 691.4 4.02 34.25 27.19 696.0 4.02 34.25 27.19 702.8 3.97 34.26 27.20 707.0 3.95 34.26 27.20 712.0 3.94 34.26. 27.20 719.0 3.90 34.27 27.21 724.0 3.89 34.27 27.21 729.0 3.88 34.27 27.22 734.8 3.88 34.27 27.22 739.0 3.87 34.28 27.22 744.0 3.84 34.28 27.23 750.2 3.83 34.28 27.23 755.0 3.81 34.28 27.23 760.0 3.81 34.29 27.24 765.0 3.82 34.30 27.25 770.6 3.81 34.30 27.25 775.0 3.81 34.30 27.25 780.0 3.78 34.31 27.26 786.4 3.75 34.30 27.25 791.0 3.68 34.30 27.26 796.0 3.65 34.30 27.27 802.4 3.67 34.31 27.27 807.0 3.67 34.31 27.27 812.0 3.67 34.31 27.27 819.0 3.66 34.32 27.28 824.2 3.66 34.32 27.28 830.4 3.60 34.32 27.28 835.0 3.59 34.31 27.28 840.0 3.58 34.32 27.28 846.2 3.57 34.32 27.29 851.0 3.57 34.33 27.29 856.0 3.56 34.33 27.30 861.8 3.54 34.33 27.30 867.4 3.53 34.33 27.30 874.0 3.52 34.33 27.30 879.0 3.52 34.33 27.31 885.4 3.51 34.34 27.31 890.0 3.51 34.34 27.31 895.0 3.50 34.34 27.31 901.2 3.48 34.35 27.32 906.0 3.48 34.35 27.32 911.0 3.46 34.35 27.32 916.8 3.45 34.35 27.33 921.0 3.45 34.36 27.33 926.0 3.44 34.36 27.34 931.4 3.44 34.37 27.34 936.0 3.44 34.37 27.34 941.0 3.46 34.37 27.34 947.2 3.45 34.38 27.35 952.2 3.44 34.38 27.35 957.0 3.44 34.38 27.35 962.6 3.53 34.40 27.35 967.0 3.52 34.41 27.37 973.2 3.53 34.41 27.37 978.0 3.52 34.42 27.37 984.4 3.51 34.42 . 27.37 989.0 3.51 34.42 27.37 994.0 3.51 34.42 27.37 1000.8 3.50 34.42 27.37 188 Table A3, (continued) CTD Data for Station JT96-04 Depth Temperature Salinity Sigma-t Depth Temperature Salinity Sigma-(m) CC) (psu) (m) CC) (psu) 0.0 14.68 31.30 23.19 185.6 6.74 33.89 26.58 1.6 14.69 31.29 23.18 190.0 6.68 33.91 26.60 2.0 14.69 31.29 23.18 195.0 6.60 33.93 26.63 3.0 14.69 31.29 23.18 200.4 6.53 33.94 26.65 4.6 14.69 31.28 23.18 206.0 6.40 33.95 26.68 5.0 14.67 31.28 23.18 211.0 6.29 33.96 26.70 7.2 13.67 31.65 23.66 216.4 6.19 33.97 26.72 8.0 13.56 31.65 23.69 221.0 6.11 33.97 26.73 9.0 13.41 31.65 23.72 227.2 6.06 33.98 26.74 10.0 13.21 31.62 23.74 232.0 6.00 33.98 26.75 11.8 12.96 31.63 23.79 237.8 5.92 33.99 26.76 12.2 12.93 31.64 23.80 242.0 5.89 33.98 26.76 13.0 12.42 31.68 23.94 247.0 5.82 33.98 26.77 14.0 11.73 31.73 24.10 252.2 5.79 33.98 26.77 15.0 11.56 31.75 24.15 257.0 5.70 33.98 26.79 16.8 11.05 31.82 24.29 263.2 5.65 33.98 26.79 17.0 10.98 31.82 24.30 268.2 5.60 33.98 26.80 18.0 10.79 31.86 24.36 274.2 5.52 33.98 26.81 19.0 10.59 31.87 24.41 279.0 5.50 33.98 26.81 20.0 10.45 31.92 24.48 284.2 5.43 34.00 26.83 21.0 10.25 31.97 24.54 291.2 5.41 34.00 26.83 22.0 10.11 32.01 24.60 297.2 536 34.00 26.84 23.0 10.07 32.12 24.69 302.0 5.34 34.00 26.84 24.2 9.87 32.10 24.71 307.4 5.34 34.00 26.85 25.0 9.72 32.12 24.75 312.0 5.34 34.01 26.85 26.4 9.47 32.12 24.79 318.0 5.34 34.01 26.85 27.0 9.55 32.13 24.79 323.0 5.33 34.01 26.85 28.8 9.38 32.18 24.85 328.6 5.34 34.01 26.85 29.0 9.39 32.19 24.86 333.0 5.34 34.01 26.85 30.0 9.37 32.24 24.90 339.0 5.35 34.01 26.85 31.0 9.31 32.24 24.91 344.0 5.32 34.02 26.86 32.2 9.45 32.29 24.92 349.0 5.27 34.03 26.88 33.0 9.59 32.33 24.93 354.8 5.25 34.03 26.88 34.0 9.52 32.33 24.95 359.0 5.24 34.03 26.88 35.0 9.63 32.41 24.99 365.0 5.22 34.03 26.88 36.0 9.45 32.43 25.04 370.0 5.19 34.04 26.89 37.0 9.34 32.45 25.07 375.4 5.15 34.04 26.90 38.0 9.29 32.46 25.08 380.2 5.11 34.05 26.91 39.0 9.25 32.46 25.09 386.2 5.08 34.06 26.92 40.0 9.22 32.47 25.11 391.0 5.04 34.06 26.92 41.0 9.17 32.48 25.12 396.2 5.02 34.06 26.93 42.0 9.14 32.48 25.13 401.0 4.99 34.06 26.93 43.6 9.09 32.49 25.14 406.0 4.96 34.07 26.94 44.0 9.00 32.50 25.16 411.0 4.93 34.08 26.95 45.0 8.94 32.51 25.18 46.6 8.81 32.53 25.22 47.0 8.80 32.54 25.22 48.0 8.80 32.54 25.22 49.8 8.77 32.56 25.24 50.0 8.77 32.55 25.24 55.0 8.48 32.63 25.35 60.0 8.40 32.68 25.39 67.8 8.18 32.83 25.55 72.0 8.13 32.87 25.58 77.0 7.89 32.92 25.66 82.0 7.97 32.99 25.70 87.0 7.92 33.03 25.74 93.2 7.92 33.07 25.77 98.0 7.80 33.25 25.93 103.2 7.74 33.37 26.03 109.4 7.74 33.43 26.08 114.0 7.69 33.50 26.14 119.0 7.69 33.54 26.17 125.2 7.75 33.59 26.20 130.0 7.80 33.64 26.23 136.2 7.66 33.70 26.30 141.0 7.52 33.72 26.33 147.4 7.39 33.76 26.39 152.0 7.24 33.78 26.43 157.0 7.09 33.80 26.46 163.6 6.95 33.84 26.51 168.0 6.91 33.85 26.53 174.4 6.86 33.86 26.54 179.0 6.79 33.88 26.56 189 Table A3, (continued) CTD Data for Station JT96-06 Depth Temperature Salinity Sigma-t (m) (°C) (psu) 0.0 13.12 31.78 23.88 1.0 13.12 31.73 23.84 3.2 13.11 31.79 23.89 4.0 13.10 31.80 23.89 6.4 13.12 31.79 23.88 7.0 13.12 31.78 23.88 8.0 13.12 31.79 23.88 10.2 13.12 31.79 23.88 11.0 13.12 31.79 23.89 13.6 13.03 31.79 23.90 14.0 13.03 31.79 23.90 15.0 12.98 31.81 23.93 16.6 12.86 31.79 23.93 17.0 12.85 31.79 . 23.94 18.0 12.37 31.82 24.05 19.0 12.14 31.80 24.08 20.6 11.48 31.86 24.25 21.0 11.38 31.86 24.26 22.0 11.37 31.85 24.26 23.0 11.33 31.86 24.27 25.2 11.02 31.88 24.35 26.0 10.97 31.88 24.36 27.0 10.83 31.92 24.41 28.6 10.22 32.01 24.58 29.0 10.20 32.01 24.59 30.0 9.98 32.05 24.66 31.0 9.95 32.06 24.66 32.8 9.86 32.10 24.71 33.0 9.85 32.11 24.72 34.0 9.71 32.16 24.79 36.6 9.61 32.21 24.83 37.0 9.61 32.21 24.83 38.0 9.60 32.21 24.84 39.0 9.56 32.22 24.86 40.4 9.50 32.24 24.88 41.0 9.44 32.25 24.90 42.0 9.38 32.28 24.93 45.4 9.28 32.32 24.97 48.2 9.17 32.37 25.03 49.0 9.19 32.38 25.04 50.0 9.14 32.40 25.06 57.2 8.74 32.53 25.23 62.0 8.56 32.60 25.30 69.0 8.46 32.65 25.36 74.0. 8.30 32.75 25.46 80.8 8.23 32.82 25.53 85.0 8.11 32.94 25.64 90.0 8.05 32.99 25.69 97.8 7.85 33.18 25.87 102.0 7.83 33.26 25.94 107.4 7.80 33.36 26.02 112.0 7.74 33.43 26.08 117.0 7.68 33.49 26.13 122.0 7.59 33.56 26.20 127.0 7.51 33.61 26.25 132.0 7.44 33.64 26.29 137.0 7.39 33.71 26.35 142.6 7.29 33.75 26.40 147.0 7.24 33.77 26.42 152.2 7.19 33.80 26.44 157.4 6.97 33.85 26.52 162.0 6.91 33.86 26.53 168.2 6.87 33.87 26.54 173.0 6.80 33.89 26.57 178.4 6.78 33.89 26.58 183.1 6.72 33.90 26.59 188.0 6.62 33.93 26.62 193.1 6.58 33.93 26.63 198.0 6.53 33.93 26.64 203.1 6.47 33.95 26.66 209.8 6.37 33.97 26.69 214.0 6.31 33.99 26.71 219.0 6.20 33.98 26.72 224.0 6.11 33.98 26.73 Depth Temperature Salinity Sigma-l (m) CO (psu) 229.6 6.06 33.98 26.74 234.0 6.03 33.99 26.75 240.0 5.97 33.99 26.76 245.0 5.84 33.99 26.77 250.0 5.82 33.99 26.78 256.0 5.80 34.00 26.78 261.0 5.72 34.00 26.80 266.0 5.64 34.00 26.81 271.6 5.62 34.00 26.81 276.0 5.60 34.00 26.82 281.0 5.59 34.00 26.82 286.0 5.53 34.01 26.82 292.6 5.45 34.01 26.84 297.0 5.43 34.01 26.84 302.0 5.42 34.01 26.84 307.0 5.39 34.01 26.85 312.0 5.38 34.02 26.85 317.8 5.33 34.02 26.86 322.2 5.32 34.02 26.86 327.0 5.29 34.02 26.87 332.0 5.26 34.03 26.88 337.0 5.23 34.03 26.88 343.4 5.22 34.03 26.88 348.4 5.22 34.04 26.89 353.0 5.18 34.05 26.90 358.0 5.18 34.05 26.90 364.2 5.13 34.06 26.91 369.0 5.11 34.05 26.91 374.0 5.08 34.06 26.92 379.0 5.00 34.05 26.93 384.0 4.99 34.06 26.93 390.0 4.97 34.06 26.94 395.0 4.92 34.07 26.95 402.6 4.88 34.07 26.95 407.0 4.84 34.07 26.96 412.0 4.82 34.07 26.96 418.2 4.82 34.09 26.97 423.0 4.81 34.09 26.98 428.0 4.79 34.09 26.98 434.8 4.77 34.10 26.99 439.0 4.76 34.10 26.99 444.0 4.75 34.10 26.99 450.2 4.73 34.10 26.99 455.0 4.71 34.11 27.00 460.0 4.67 34.11 27.00 466.8 4.65 34.11 27.01 471.0 4.64 34.11 27.01 476.0 4.61 34.11 27.01 482.6 4.56 34.12 27.03 487.0 4.54 34.12 27.03 492.0 4.53 34.13 27.03 499.0 4.52 34.13 27.04 504.0 4.51 34.14 27.05 510.0 4.51 34.15 27.06 515.0 4.53 34.15 27.06 522.6 4.54 34.16 27.06 527.0 4.56 34.16 27.06 532.0 4.57 34.16 27.06 537.6 4.58 34.17 27.06 542.0 4.59 34.17 27.06 547.0 4.59 34.17 27.06 554.2 4.59 34.17 27.06 559.0 4.58 34.17 27.07 565.4 4.58 34.18 27.08 570.0 4.56 34.19 27.08 575.0 4.54 34.19 27.08 580.0 4.55 34.19 27.09 585.0 4.55 34.20 27.09 590.0 4.54 34.20 27.09 595.0 4.52 34.21 27.10 602.0 4.50 34.21 27.11 607.0 4.49 34.21 27.11 613.8 4.48 34.22 27.11 618.0 4.46 34.22 27.12 Depth Temperature Salinity Sigma-t (m) (°C) (psu) 623.0 4.43 34.22 27.12 629.0 4.42 34.23 27.13 634.0 4.41 34.23 27.13 639.0 4.41 34.23 27.13 644.2 4.40 34.24 27.14 649.0 4.38 34.24 27.14 654.0 4.38 34.24 27.14 659.0 4.37 34.24 27.14 664.2 4.35 34.25 27.15 669.0 4.32 34.25 . 27.16 674.0 4.31 34.25 27.16 679.4 4.28 34.26 27.17 684.0 4.28 34.26 27.17 689.0 4.26 34.27 27.18 694.0 4.25 34.27 27.18 699.0 4.23 34.28 27.19 190 Table A3, (continued) CTD Data for Station JT96-09 Depth Temperature Salinity Sigma-t Depth Temperature Salinity Sigma-t Depth Temperature Salinity Sigma-i (m) (°C) (psu) (m) CC) (psu) (m) (°C) (psu) 0.0 14.59 31.52 23.38 208.0 6.66 33.94 26.63 599.0 4.34 34.17 27.09 1.0 14.59 31.52 23.38 213.0 6.59 33.94 26.64 604.8 4.32 34.17 27.09 3.0 14.60 31.52 23.38 218.2 6.53 33.94 26.65 609.0 4.32 34.17 27.09 4.0 14.59 31.53 23.38 223.0 6.50 33.94 26.65 615.8 4.32 34.17 27.10 6.2 14.59 31.52 23.38 228.4 6.45 33.95 26.66 620.2 4.31 34.17 27.10 7.0 14.58 31.54 23.40 233.0 6.43 33.94 26.66 625.0 4.29 34.18 27.10 8.0 14.52 31.57 23.43 238.8 6.38 33.95 26.68 630.6 4.29 34.18 27.10 9.4 14.32 31.58 23.48 243.0 6.31 33.95 26.69 635.0 4.26 34.19 27.11 10.0 14.32 31.58 23.48 250.4 6.21 33.95 26.70 641.8 4.23 34.19 27.12 12.8 14.29 31.58 23.48 255.0 6.17 33.95 26.70 646.0 4.22 34.19 27.12 13.0 14.27 31.59 23.50 260.4 6.11 33.95 26.71 651.0 4.21 34.19 27.12 15.4 14.18 31.62 23.54 265.0 6.06 33.96 26.72 657.6 4.21 34.20 27.13 16.0 14.19 31.63 23.54 271.2 5.98 33.97 26.74 662.0 4.19 34.20 27.13 17.0 14.11 31.69 23.61 276.0 5.93 33.97 26.75 667.0 4.18 34.21 27.14 18.8 13.95 31.75 23.68 282.2 5.91 33.96 26.75 672.8 4.17 34.21 27.14 19.0 13.97 31.72 23.66 287.0 5.91 33.96 26.75 677.0 4.15 34.22 27.15 20.0 13.91 31.75 23.69 292.6 5.87 33.97 26.75 683.2 4.14 34.22 27.15 21.8 13.77 31.92 23.86 297.0 5.82 33.97 26.76 688.0 4.14 34.23 27.16 22.0 13.73 31.95 23.89 303.0 5.78 33.98 26.77 695.0 4.14 34.23 27.16 23.0 13.04 32.12 24.16 308.0 5.76 33.98 26.78 700.2 4.10 34.24 27.17 24.0 12.78 32.19 24.26 313.0 5.75 33.99 26.79 707.2 4.10 34.25 27.18 25.0 12.44 32.54 24.60 319.0 5.68 33.99 26.79 712.0 4.09 34.25 27.18 26.0 12.34 32.41 24.51 324.0 5.65 34.00 26.80 717.0 4.09 34.26 27.19 27.0 12.23 32.33 24.47 329.0 5.62 33.99 26.81 723.0 4.08 34.26 27.19 28.8 11.71 32.43 24.65 335.0 5.59 34.00 26.81 728.0 4.06 34.27 27.20 29.0 11.72 32.44 24.65 340.0 5.56 34.00 26.82 733.0 4.06 34.27 27.20 30.0 11.64 32.50 24.71 345.0 5.54 34.01 26.83 738.6 4.05 34.28 27.21 31.0 11.51 32.46 24.71 351.8 5.50 34.01 26.83 743.0 4.05 34.28 27.21 32.2 11.29 32.56 24.83 356.0 5.48 34.01 26.83 748.2 4.03 34.29 27.22 33.0 11.19 32.69 24.95 361.0 5.45 34.02 26.85 754.2 4.03 34.29 27.22 34.0 11.12 32.46 24.78 366.0 5.44 34.02 26.85 759.0 4.02 34.30 27.22 36.2 11.00 32.53 24.85 372.0 5.42 34.02 26.85 764.2 4.00 34.30 27.23 37.0 10.94 32.51 24.85 377.0 5.40 34.03 26.86 770.2 3.99 34.30 27.23 39.6 10.81 32.47 24.84 382.0 5.35 34.03 26.87 775.0 3.96 34.31 27.24 40.0 10.80 32.47 24.84 387.8 5.30 34.04 26.88 780.2 3.95 34.32 27.25 41.0 10.84 32.47 24.83 392.0 5.29 34.03 26.87 786.6 3.94 34.32 27.25 42.8 10.63 32.47 24.87 397.0 5.25 34.04 26.88 791.0 3.94 34.32 27.25 43.0 10.61 32.47 24.88 402.0 5.20 34.04 26.89 796.2 3.93 34.32 27.26 44.0 10.55 32.47 24.88 408.0 5.15 34.04 26.90 801.8 3.92 34.33 27.26 45.0 10.49 32.47 24.90 413.0 5.12 34.05 26.91 806.0 3.90 34.33 27.27 46.0 10.41 32.48 24.92 418.0 5.11 34.04 26.90 811.0 3.90 34.34 27.27 47.0 10.34 32.47 24.92 424.6 5.10 34.04 26.91 816.0 3.90 34.34 27.27 48.0 10.27 32.47 24.93 429.0 5.10 34.05 26.91 822.2 3.89 34.34 27.27 49.8 10.07 32.59 25.06 434.0 5.05 34.05 26.92 827.0 3.89 34.34 27.27 50.0 10.03 32.50 24.99 439.0 5.04 34.06 26.92 832.0 3.89 34.34 27.27 55.0 9.73 32.50 25.04 445.2 5.02 34.06 26.93 841.0 3.89 34.34 27.28 60.6 9.40 32.49 25.09 450.0 5.00 34.06 26.93 846.0 3.88 34.34 27.28 65.0 9.27 32.51 25.13 456.4 4.96 34.07 26.95 852.6 3.88 34.34 27.28 72.0 9.04 32.55 25.19 461.0 4.93 34.09 26.96 857.0 3.89 34.34 27.28 77.0 8.72 32.65 25.32 466.0 4.88 34.08 26.96 862.0 3.88 34.35 27.28 82.8 8.37 32.74 25.45 472.4 4.85 34.09 26.97 867.6 3.85 34.35 27.29 87.0 8.17 32.82 25.54 477.0 4.82 34.10 26.98 872.2 3.83 34.36 27.29 92.0 8.28 32.89 25.57 484.2 4.80 34.10 26.98 877.0 3.80 34.36 27.30 97.4 8.08 33.11 25.78 489.0 4.75 34.10 26.99 885.4 3.78 34.37 27.31 102.0 7.88 33.28 25.94 495.0 4.70 34.10 26.99 890.2 3.76 34.37 27.31 108.0 7.92 33.36 26.00 500.0 4.68 34.10 27.00 895.0 3.73 34.38 27.32 113.0 7.89 33.44 26.06 505.6 4.68 34.10 27.00 900.0 3.72 34.38 27.32 119.4 7.83 33.55 26.16 510.0 4.68 34.10 27.00 124.0 7.71 33.64 26.24 515.0 4.70 34.12 27.01 129.6 7.67 33.68 26.29 521.2 4.66 34.12 27.02 134.0 7.61 33.71 26.32 526.0 4.67 34.13 27.02 140.8 7.51 33.78 26.39 531.6 4.66 34.13 27.03 145.0 7.40 33.81 26.42 536.2 4.64 34.14 27.03 150.0 7.36 33.81 26.43 541.0 4.62 34.14 27.03 155.4 7.33 33.83 26.45 546.0 4.58 34.14 27.04 160.0 7.28 33.84 26.46 551.2 4.57 34.15 27.05 166.0 7.17 33.86 26.50 556.2 4.55 34.15 27.05 171.0 7.12 33.87 26.51 562.2 4.50 34.15 27.06 177.2 7.02 33.88 26.54 567.0 4.44 34.15 27.07 182.0 6.96 33.90 26.56 572.2 4.44 34.15 27.06 187.8 6.88 33.91 26.58 578.0 4.42 34.16 27.07 192.0 6.83 33.92 26.59 583.0 4.40 34.16 27.07 197.0 6.78 33.93 26.60 589.0 4.38 34.16 27.08 203.8 6.71 33.93 26.62 594.0 4.37 34.16 27.08 191 Table A4. Dissolved oxygen data for Stations JT96-01,2,4, 5,6 and 9. Station JT96-01 Station JT96-05 Depth Oxygen Oxygen Depth Oxygen Oxygen (m) (ml/L) (umol/L) (m) (ml/L) (umol/L) 10 7.69 343.49 100 4.93 220.25 20 5.88 262.56 300 2.31 103.00 30 5.44 242.79 450 1.24 55.18 50 4.49 200.48 650 0.72 32.19 75 3.02 134.73 800 0.41 18.39 100 2.72 121.39 900 0.52 22.99 115 2.38 106.22 1000 0.45 20.23 1100 0.41 18.39 1250 0.42 18.85 1400 0.63 28.05 1720 1.19 53.34 Station JT96-02 Station JT96-06 Depth Oxygen Oxygen Depth Oxygen Oxygen (m) (ml/L) (umol/L) (m) (ml/L) (umol/L) 25 5.92 264.40 10 6.34 283.25 50 6.06 270.38 50 4.39 195.88 75 5.56 248.30 100 4.44 198.18 100 5.01 223.47 200 2.12 94.72 200 2.18 97.48 300 1.56 69.43 400 0.95 42.30 400 1.03 45.98 600 0.52 22.99 500 0.72 32.19 800 0.36 16.09 550 0.60 26.67 1000 0.37 16.55 600 0.52 22.99 1200 0.37 16.55 650 0.46 20.69 1240 0.43 19.31 700 0.41 18.39 Station JT96-04 Station JT96-09 Depth Oxygen Oxygen Depth Oxygen Oxygen (m) (ml/L) (umol/L) (m) (ml/L) (umol/L) 10 6.76 301.64 10 6.19 276.35 20 6.31 281.87 20 6.20 276.81 30 5.72 255.20 30 6.79 303.02 50 4.99 222.55 50 6.72 299.80 75 5.30 236.81 75 6.28 280.49 100 n.a. n.a. 100 4.33 193.13 150 2.91 130.13 200 2.73 121.85 200 2.08 92.88 400 1.29 57.48 300 1.59 70.81 600 0.58 25.75 350 1.35 60.24 800 0.35 15.63 410 0.98 43.68 900 0.32 14.25 192 Table AS. Magnetic Susceptibility Data for Sediment Cores from Stations JT96-02,5, 6 and 9. Piston Core JT96-02 Piston Core JT95-05 (I) Piston Core JT96-05 (2) Piston Core JT96-05 (4) Trigger Core JT96-05 (3) Depth Mag. Susc. Depth Mag. Susc. Depth Mag. Susc. Depth Mag. Susc. Depth Mag. Su (cm) (cm) (cm) (cm) (cm) 5 0.93 5 0.75 5 1.89 5 1.69 5 0.99 10 1.22 10 0.92 10 2.07 10 2.44 10 1.10 15 1.63 15 0.90 15 2.35 15 3.48 15 1.16 20 2.16 20 0.90 20 2.74 20 3.48 20 1.27 25 2.27 25 0.92 25 3.18 25 3.30 25 1.34 30 2.24 30 0.88 30 3.54 30 3.47 30 1.47 35 2.47 35 0.86 35 4.01 35 3.91 35 1.58 40 3.13 40 0.89 40 3.53 40 5.69 40 1.59 45 3.83 45 1.11 45 3.29 45 4.99 45 1.62 50 3.71 50 1.04 50 3.83 50 2.56 50 1.81 55 4.05 55 1.10 55 4.16 55 2.01 55 1.92 60 4.43 60 1.16 60 4.09 60 1.93 60 2.06 65 4.42 65 1.16 65 5.23 65 1.17 65 2.26 70 5.16 70 1.21 70 4.90 70 0.86 70 2.33 75 5.37 75 1.37 75 4.97 75 0.91 75 2.38 80 5.48 80 1.47 80 5.12 80 0.74 80 2.57 85 4.35 85 1.46 85 4.95 85 0.58 85 2.49 90 2.62 90 1.31 90 4.77 90 0.53 90 2.68 95 2.22 95 1.48 95 4.89 95 0.59 95 2.68 100 2.26 100 1.75 100 5.57 100 0.70 100 2.44 105 2.40 105 1.89 105 5.45 105 0.47 105 1.84 110 2.17 110 1.95 110 4.64 110 0.47 110 1.27 115 2.27 115 2.02 115 4.13 115 0.72 115 1.22 120 2.25 120 2.10 120 3.94 120 0.97 120 1.28 125 2.37 125 2.68 125 4.26 125 1.07 125 1.30 130 2.34 130 3.10 130 4.36 130 1.07 130 1.30 135 2.25 135 3.49 135 4.38 135 0.72 140 2.31 140 3.32 140 3.91 140 0.67 145 2.32 145 4.24 145 3.66 145 0.76 150 2.30 150 4.72 150 0.80 155 2.25 155 2.78 155 0.68 160 2.35 160 0.62 165 2.40 165 0.58 170 2.26 170 0.62 175 2.12 175 0.73 180 2.34 180 0.85 185 2.23 185 0.72 190 2.21 190 0.62 195 2.45 195 0.67 200 2.57 200 0.64 205 2.32 205 0.30 193 Table AS. (continued) Piston Core JT96-06 Piston Core JT96-09 Depth Mag. Susc. (cm) 5 2.52 10 2.37 15 1.71 20 2.41 25 2.58 30 2.75 35 2.70 40 3.52 45 3.19 50 3.24 55 2.56 60 2.19 65 2.19 70 2.44 75 2.59 80 2.14 85 1.83 90 1.78 95 1.86 100 1.82 105 2.23 110 2.34 115 2.35 120 2.46 125 2.46 130 2.43 135 2.48 140 2.72 145 2.77 150 2.58 155 2.34 160 2.25 165 2.12 170 2.52 175 2.81 180 2.80 185 2.43 190 2.05 195 2.78 200 3.10 205 2.95 210 2.84 215 2.45 220 2.57 225 2.48 230 2.46 235 2.60 240 2.86 245 2.81 250 2.57 255 2.62 260 2.90 265 2.68 270 2.94 275 2.92 280 3.10 285 3.43 290 3.30 295 4.31 300 5.22 305 4.48 310 3.00 Depth Mag. Susc. (cm) 5 1.20 10 1.47 15 1.50 20 1.41 25 1.25 30 1.04 35 0.92 40 0.90 45 1.06 50 1.69 55 2.19 60 2.54 65 2.99 70 2.85 75 1.33 80 1.51 85 1.60 90 1.57 95 1.58 100 1.58 105 1.62 no 1.66 115 1.68 120 1.69 125 1.73 130 2.02 135 3.01 140 4.20 145 4.45 150 2.63 155 2.32 160 2.30 165 2.63 170 . 2.73 175 2.39 180 2.15 185 2.52 190 2.77 195 2.81 200 2.91 205 2.90 210 2.74 215 2.72 220 2.74 225 2.29 230 n.a. 235 2.77 240 2.68 245 2.51 250 2.74 255 2.62 260 2.66 265 2.60 270 3.09 275 2.70 280 2.53 285 2.63 290 2.63 295 2.61 300 2.54 305 2.77 310 2.93 315 2.78 320 2.61 325 2.57 330 2.91 335 2.81 340 2.75 345 2.86 350 2.89 355 3.07 360 2.65 365 2.56 370 2.40 Table A6. Radiocarbon Data for Planktonic and Benthic Foraminifera, and Bulk Organic Carbon. Core / Sample Depth Corrected Depth Calendar Age "C Age "C Age Benthic Foraminifera "C Age CAMS#* (cm) (cm) (kyr B.P.) Mixed Planktonics Uvigerina spp. Bolivina argentea Organic Carbon JT96-01 35-38 mc 36.5 0.92 1 1780 ±30 67353 JT96-02 101-103 pc 102.0 14.09 2 13280 ±60 70907 201-203 pc 202.0 15.68 2 14240 ±60 70908 JT96-05 46-48 be 47.0 10.29 ' 9960 ±50 60245 JT96-09 38-40 mc 39.0 39.0 9830 ±110 9920 ± 40 40427, 56704 20-21 pc 20.5 32.5 10050 ±110 40426 35-36 pc 35.5 47.5 10.03' 9760 ±70 10710 ±60 10650 ±90 60240, 60241, 36-37 pc 36.5 48.5 11390 ±80 78516 45-46 pc 45.5 57.5 9360 ±240 10910 ±70' 78456, 78457 65-66 pc 65.5 77.5 12.24' 11210 ± 120 12380 ±280' 42160, 78458 66-67 pc 66.5 78.5 14790 ±450 78517 75-76 pc 75.5 87.5 12.73 1 11500 ±110 12110 ±80 60242, 62465 90-91 pc 90.5 102.5 12.84 ' 11600 ±80 12700 ±70' 60243, 78459 96-97 pc 96.5 108.5 15660 ±70 62468 100-101 pc 100.5 112.5 13.17 ^ 12460 ± 120 13170 ±100 40428, 48152 130-131 pc 130.5 142.5 13.43 2 12640 ±90 13500 ±80 62466, 62797 165-166 pc 165.5 161.5 10090 ±50 13210 ±150 60244, 62469 166-167 pc 166.5 162.5 9859 ±^0 24480 ± 120 62467,62471 170-171 pc 170.5 166.5 13240 ±200 62798 265-266 pc 265.5 261.5 14.14' 13410 ±80 14290 ±110 40430,42161 266-267 pc 266.5 262.5 24500 ±120 62469 290-291 pc 290.5 286.5 14.30' 13520 ±70 14350 ± 120 40431,42162 350-351 pc 350.5 346.5 15.57' 14140 ±70 14830 ± 280 40432, 78460 351-352 pc 351.5 347.5 21440 ± 100 57327 370-371 pc 370.5 366.5 20320 ± 90 57104 Tul96-05 31-32pc 31.5 77.5 3760 ± 40 58708 199-200 pc 199.5 234.5 10.13 ' 9800 ±90 11150 ±70 58463, 58460 200-201 pc 200.5 235.5 11200 ±340 11140 ± 60 57558,57324 290-291 pc 290.5 306.5 28120 ±390 28980 ± 290 58463, 58461 360-361 pc 360.5 371.5 32590 ± 1000 32900 ± 740 58465, 58462 480-481 pc 480.5 481.5 38400 ± 1200 38300 ± 2600 57325, 57326 1 Calculated using Calib 4.3 (Stuiver et al.. 1998) and a reservoir age of 800 yrs. 2 Calculated using Calib 4.3 (Stuiver et al.. 1998) and a reservoir age of 1100 yrs. 3 For sample 45-46 the species was Bolivina spicca (not Bolivina argentea ) and for samples 65-66 and 90-91 a mixture of Bolivina species were used. 4 Lawrence Livermore laboratory reference number 195 Table A7.1,0Pb data for multicores (mc) JT96-01, 02 and 09 Multicore JT96-01 Sample Depth Total 2,°Pb Excess 21°Pb (cm) (dpm/g) (dpm/g) 0-1 mc 0.5 24.47 22.65 2-3 mc 2.5 20.27 18.45 4-5 mc 4.5 20.32 18.50 6-8 mc 7.0 21.20 19.37 10-12 mc 11.0 18.94 17.12 14-16 mc 15.0 18.28 16.46 18-20mc 19.0 13.21 11.38 25-30 mc 27.5 4.00 2.17 35-38 mc 36.5 1.82 0.00 Multicore JT96-02 Sample Depth Total !l0Pb Excess ""Pb (cm) (dpm/g) (dpm/g) 0-1 mc 0.5 2-3 mc 2.5 4-5 mc 4.5 6-8 mc 7.0 10-12 mc 11.0 14-16 mc 15.0 18-20mc 19.0 58.84 56.06 27.50 24.71 17.74 15.95 9.19 6.41 3.80 1.02 2.22 -0.57 2.79 0.00 Multicore JT96-09 Sample Depth Total !'"Pb Excess 3j0Pb (cm) (dpm/g) (dpm/g) 0-2 mc 1.0 23.64 22.04 2-3 mc 2.5 38.23 36.63 4-5 mc 4.5 17.50 15.90 6-8 mc 7.0 9.41 7.51 8-10 mc 9.0 12.16 10.56 14-16 mc 15.0 6.62 5.02 18-20 mc 19.0 1.93 0.34 24-26 mc 25.0 1.37 -0.22 28-30 mc 29.0 1.60 0.00 34-36 mc 35.0 1.54 -0.06 38-40 mc 39.0 1.77 0.17 Table A8. Stable Isotope Data for Foraminifera from Cores JT96-02pc, 06pc, 09mc and 09pc, and Tul96-05pc, 03tc and 03pc. Data for Piston Core (pc) JT96-02. Sample Depth Calendar Age Uvigerina spp. Bolivina argentea Bolivina pacifica N. pachyderma G. bulloides (cm) (kyrs) 5"C 5"Q 8"C 8"Q 5"C S"Q 5"C 8"0 8"C 8"Q 1-3 pc 2.0 11 -13 pc 12.0 21-23 pc 22.0 31-33 pc 32.0 41-43 pc No sample 51-53 pc No sample 61-63 pc No sample 71-73 pc No sample 81-83 pc No sample 91-93 pc 92.0 101-103 pc 102.0 111-113 pc 112.0 121-123 pc 122.0 131-133 pc 132.0 141-143 pc 142.0 151-153 pc 152.0 -1.13 4.40 -1.30 3.94 161-163 pc 162.0 -1.21 4.42 -1.21 4.08 171-173 pc 172.0 -1.10 4.43 181-183 pc 182.0 -1.05 4.39 -1.55 3.81 191-193 pc 192.0 210-203 pc 212.0 -1.49 4.06 Data for Piston Core (pc) JT96-06. Sample Depth Calendar Age Uvigerina spp. Bolivina argentea Bolivina pacifica N. pachyderma C. bulloides (cm) (kyis) S"C 8"0 8"C 8"Q 8"C 8'"0 8"C 8"0 8"C TO 0-1 pc 0.5 10-11 pc 11.5 20-21 pc 21.5 -0.79 3.66 30-31 pc 31.5 -0.99 3.55 40-41 pc 41.5 -0.89 3.26 -1.12 3.55 50-51 pc 51.5 60-61 pc 61.5 -0.65 3.24 -1.14 3.29 70-71 pc 71.5 -0.9 3.64 80-81 pc 81.5 -0.78 3.75 90-91 pc 91.5 -0.87 3.22 100-101 pc 101.5 -0.75 3.97 -0.95 3.66 110-111 pc 111.5 -0.82 3.39 120-121 pc 121.5 130-131 pc 131.5 -0.56 3.08 -0.55 3.71 140-141 pc 141.5 150-151 pc 151.5 -0.95 3.30 160-161 pc 161.5 -1.14 3.24 170-171 pc 171.5 -1.20 3.75 180-181 pc 181.5 -0.95 3.44 190-191 pc 191.5 200-201 pc 201.5 -1.32 3.15 210-211 pc 211.5 -1.35 3.70 220-221 pc 221.5 230-231 pc 231.5 -1.07 2.74 240-241 pc 241.5 250-251 pc 251.5 -0.47 2.79 260-261 pc 261.5 -1.35 3.62 270-271 pc 271.5 -0.94 3.88 280-281 pc 281.5 -1.26 3.87 290-291 pc 291.5 -1.27 3.87 300-301 pc 301.5 -1.33 3.95 310-311 pc 309.5 -1.24 3.99 197 Table A8. (continued) Data for Multicore (mc) and Piston Core (pc) JT96-09. Sample Depth * Calendar Age Uvigerina spp. Bolivina argentea Bolivina pacifica N. pachyderma G. bulloides (cm, corr) (kyrs) 8"C 8"0 5,3C S"0 8"C S"0 S"C S"0 8IJC 8"0 24-26 mc 25.0 5.28 -1.23 3.02 15-16 pc 27.5 5.81 -0.83 2.96 30-32 mc 31.0 6.54 -0.94 2.96 -0.97 2.87 20-21 pc 32.5 6.86 -0.59 3.02 -0.89 2.93 34-36 mc 35.0 7.39 -0.75 3.00 25-26 pc 37.5 7.92 -0.81 3.17 -1.27 2.89 -1.04 2.86 38-40 mc 39.0 8.23 -0.66 3.16 -0.84 3.20 30-31 pc 42.5 8.97 -0.70 3.21 -1.12 3.06 35-36 pc 47.5 10.03 -0.85 3.16 -0.65 3.09 40-41 pc 52.5 10.40 -0.90 3.21 45-46 pc 57.5 10.76 -1.57 3.14 0.42 1.25 50-51 pc 62.5 11.13 -1.15 3.64 -1.21 3.30 0.49 1.10 55-56 pc 67.5 11.50 -1.37 3.49 65-66 pc 77.5 12.24 -1.20 3.34 75-76 pc 87.5 12.73 -0.88 3.61 -1.03 3.42 0.32 1.54 0.03 1.80 80-81 pc 92.5 12.77 -1.22 3.43 -2.31 3.53 -0.09 1.03 85-86 pc 97.5 12.80 -1.06 3.54 -2.15 3.39 90-91 pc 102.5 12.84 -1.42 3.55 -0.06 1.36 95-96 pc 107.5 13.00 -2.56 3.84 0.09 1.23 100-101 pc 112.5 13.17 -2.33 3.39 0.14 1.23 105-106 pc 117.5 13.21 -2.36 3.85 0.05 1.32 110-111 pc 122.5 13.26 0.06 1.07 115-116 pc 127.5 13.30 -2.68 3.84 0.13 1.53 -0.26 1.50 120-121 pc 132.5 13.35 -1.16 3.61 -2.53 3.59 125-126 pc 137.5 13.39 -1.18 3.75 0.01 1.21 130-131 pc 142.5 13.43 -1.48 3.58 0.22 1.36 135-136 pc 147.5 13.46 -1.03 3.69 -2.40 3.70 -0.27 2.12 150-151 pc 148.5 13.47 -1.12 3.92 -0.23 1.84 155-156 pc 151.5 13.49 -1.08 3.94 160-161 pc 156.5 13.52 -Ul 3.66 0.26 2.88 165-166 pc 161.5 13.55 -1.03 4.00 0.08 2.42 -0.14 2.75 170-171 pc 166.5 13.58 -1.04 4.02 -0.45 0.71 -0.22 0.59 175-176 pc 171.5 13.61 -1.00 4.03 -0.14 2.68 185-186 pc 181.5 13.67 -1.08 4.26 0.50 2.57 0.28 1.64 190-191 pc 186.5 13.70 -1.35 4.01 -0.32 2.46 195-196 pc 191.5 13.72 -1.20 4.16 200-201 pc 196.5 13.75 -1.39 3.94 -0.27 2.15 -0.44 2.66 210-211 pc 206.5 13.81 -1.21 3.98 -0.37 2.62 215-216 pc 211.5 13.84 -0.23 2.35 220-221 pc 216.5 13.87 -1.28 4.17 260-261 pc 256.5 14.11 -1.09 4.04 -1.21 3.99 -0.28 2.58 265-266 pc 261.5 14.14 -0.15 2.20 275-276 pc 271.5 14.21 -1.09 4.09 -0.21 2.43 -0.42 2.70 280-281 pc 276.5 14.24 -1.01 4.15 0.01 2.48 -0.24 2.60 290-291 pc 286.5 14.30 -1.01 4.11 -0.37 2.34 -0.70 2.44 300-301 pc 296.5 14.52 -0.95 4.14 -0.30 2.88 -0.60 2.85 310-311 pc 306.5 14.73 -1.10 4.14 -1.17 4.17 -0.70 2.45 320-321 pc 316.5 14.94 -1.12 4.12 -0.24 2.57 -0.48 2.18 330-331 pc 326.5 15.15 -0.07 2.42 -0.58 2.49 340-341 pc 336.5 15.36 -1.10 4.14 -0.24 2.41 -0.48 2.55 350-351 pc 346.5 15.57 -1.36 4.08 -0.32 2.50 -0.57 2.67 360-361 pc 356.5 15.78 -1.24 4.10 -1.23 4.10 -0.13 2.48 -0.55 2.47 370-371 pc 366.5 15.99 -0.03 2.64 -0.86 2.54 * Corrected Depth 198 Table A8. (continued) Data for Trigger Core (tc) and Piston Core (pc) Tul96-03. Sample Depth (cm) Calendar Age (kyrs) Uvigerina spp. 8"C S"Q Bolivina argentea 8"C 8"Q Bolivina pacified 8"C 8"Q N. pachyderma 8"C 8"Q G. bulloides 8"C 8"Q 5-6 tc 10-11 tc 15-16 tc 20-21 tc 25-26 tc 10-11 pc 15-16 pc 20-21 pc 25-26 pc 30-31 pc 40-41 pc 50-51 pc 60-61 pc 65-66 pc 70-71 pc 80-81 pc 90-91 pc 100-101 pc 110-111 pc 120-121 pc 130-131 pc 140-141 pc 150-151 pc 160-161 pc 165-166 pc 170-171 pc 180-181 pc 190-191 pc 200-201 pc 210-211 pc 220-221 pc 230-231 pc 240-241 pc 250-251 pc 260-261 pc 270-271 pc 280-281 pc 290-291 pc 300-301 pc 310-311 pc 320-321 pc 330-331 pc 340-341 pc 350-351 pc 360-361 pc 370-371 pc 380-381 pc 390-391 pc 400-401 pc 410-411 pc 420-421 pc 430-431 pc 5.5 10.5 15.5 20.5 25.5 10.5 15.5 No sample 25.5 30.5 40.5 No sample No sample 65.5 70.5 80.5 90.5 100.5 110.5 120.5 130.5 140.5 No sample No sample 165.5 170.5 180.5 190.5 200.5 210.5 220.5 230.5 240.5 250.5 No sample 270.5 280.5 290.5 300.5 310.5 320.5 330.5 340.5 350.5 360.5 370.5 380.5 390.5 400.5 410.5 420.5 430.5 -1.20 -1.21 3.39 4.00 -1.05 3.87 -0.97 3.68 -0.95 3.63 -0.84 3.94 199 Table A8. (continued) Data for Piston Core (pc) Tul96-05. Sample Depth * Calendar Age Uvigerina spp. Bolivina argentea Bolivina pacijica N. pachydermia G. bulloides (cm.corr.) (kyrs) 8"C 8"Q S"C 8"Q 8"C 5"Q 5"C S"Q 5"C 8"Q 10-11 pc 56.5 2.35 -0.84 2.38 20-21 pc 66.5 2.76 -1.29 3.09 30-31 pc 76.5 3.18 40-41 pc 86.5 3.62 50-51 pc No sample 60-61 pc 65.5 4.01 70-71 pc 105.5 4.45 -1.30 3.22 -0.95 3.20 80-81 pc 115.5 4.89 90-91 pc 125.5 5.33 -1.49 2.65 -0.98 3.03 100-101 pc 135.5 5.77 -1.26 3.04 110-111 pc 145.5 6.21 -1.27 2.90 120-121 pc 155.5 6.65 130-131 pc 165.5 7.09 -1.52 2.70 -0.99 3.00 140-141 pc 175.5 7.53 -1.18 2.26 0.60 0.82 0.33 0.75 150-151 pc 185.5 7.97 -1.45 2.81 -1.23 2.91 0.26 1.05 160-161 pc 195.5 8.41 -1.39 2.51 -1.20 2.65 170-171 pc 205.5 8.85 -1.01 3.02 180-181 pc 215.5 9.29 -1.19 2.94 190-191 pc 225.5 9.73 -1.45 3.32 -1.11 3.21 0.48 1.28 200-201 pc 235.5 10.17 -1.49 3.52 -1.23 3.25 0.18 1.09 -0.06 1.05 210-211 pc 245.5 10.61 -1.26 3.58 -0.54 3.19 0.33 1.56 220-221 pc No sample 230-231 pc Turbidite -1.24 3.86 240-241 pc 256.5 27.89 0.01 2.92 -0.38 3.04 250-251 pc 266.5 28.70 0.15 2.89 260-261 pc 276.5 29.50 -1.44 4.31 -0.01 2.86 -0.27 2.84 270-271 pc 286.5 30.30 0.29 3.37 -0.11 2.28 280-281 pc 296.5 31.11 0.19 2.64 290-291 pc 306.5 31.91 -1.34 4.21 -0.07 2.42 -0.05 2.81 300-301 pc 316.5 32.71 -1.20 4.05 0.11 2.70 310-311 pc 326.5 33.52 0.04 2.59 320-321 pc 336.5 34.32 -1.42 4.08 0.09 2.58 -0.05 2.60 330-331 pc No sample 340-341 pc 351.5 35.53 -0.09 2.38 350-351 pc 361.5 36.33 -1.36 3.94 0.07 2.71 -0.05 2.24 360-361 pc 371.5 37.13 -1.70 4.14 -0.36 2.00 -0.43 2.25 370-371 pc 381.5 37.75 -1.71 4.28 -0.38 1.86 -0.86 1.96 380-381 pc 391.5 38.37 -1.62 4.09 -0.46 2.55 -0.82 2.32 390-391 pc 401.5 38.98 -1.41 4.17 -0.21 2.74 400-401 pc 411.5 39.60 -1.55 3.77 0.01 2.73 410-411 pc 421.5 40.22 -1.40 4.32 -0.39 2.61 420-421 pc No sample 430-431 pc 431.5 40.83 -1.54 4.21 -0.22 2.18 440-441 pc 441.5 41.45 -1.64 4.09 0.34 2.05 450-451 pc 451.5 42.07 -1.40 4.26 -0.32 2.75 -0.83 2.45 460-461 pc 461.5 42.68 -1.58 4.20 -0.10 2.66 -0.33 2.82 470-471 pc 471.5 43.30 -1.51 4.26 0.05 2.50 -0.20 2.05 480-481 pc 481.5 43.92 -1.52 4.20 -0.11 2.44 -0.38 2.10 490-491 pc 491.5 44.53 -1.45 4.16 0.04 2.65 500-501 pc 501.5 45.15 -1.63 4.19 -0.33 2.57 -0.30 2.14 200 Table A9. Geochemical Data for Multicore JT96-01. General Geochemical Data for Multicore (mc) JT96-01. Sample Depth Calendar Age cr 8"N 5"C„, N„ C S,„ CaCO, c„. c,^ Opal Ban, (cm) (kyrs) (wt. %) (%o vs air) ( %o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (ng/g) 0-1 mc 0.5 0.01 4.52 4.81 -21.81 0.24 2.18 0.31 0.72 2.09 8.6 10.9 61 l-2mc 1.5 0.04 2.56 0.22 1.95 0.24 0.60 1.88 8.7 2-3 mc 2.5 0.06 2.73 0.22 1.97 0.24 0.64 1.89 8.7 3-4 mc 3.5 0.09 2.30 0.21 1.92 0.25 0.68 1.84 8.6 4-5 mc 4.5 0.11 2.18 0.22 1.95 0.25 0.63 1.87 8.7 5-6 mc 5.5 0.14 2.46 4.77 -22.15 0.21 1.98 0.26 0.63 1.91 9.0 9.2 80 6-8 mc 7.0 0.18 2.16 0.21 1.87 0.26 0.62 1.80 8.7 8-10 mc 9.0 0.23 1.88 0.21 1.87 0.26 0.68 1.78 8.7 10-12 mc 11.0 0.28 2.00 4.81 -22.25 0.21 1.93 0.28 0.75 1.84 8.9 9.7 85 12-14 mc 13.0 0.33 1.80 0.20 1.87 0.27 0.72 1.78 8.8 14-16 mc 15.0 0.38 2.02 5.29 -22.30 0.20 1.92 0.30 0.71 1.83 9.0 10.2 65 16-18 mc 17.0 0.43 1.67 0.19 1.85 0.32 0.65 1.77 9.1 18-20mc 19.0 0.48 1.95 4.82 -22.33 0.20 1.87 0.33 0.66 1.79 9.2 40 20-25 mc 22.5 0.57 1.90 5.16 -22.31 0.19 1.80 0.41 0.68 1.72 9.2 10.4 73 25-30 mc 27.5 0.70 1.87 4.86 -22.25 0.17 1.71 0.49 0.82 1.61 9.3 29 30-35 mc 32.5 0.82 1.89 5.51 -22.06 0.17 1.64 0.54 0.80 1.55 9.2 8.2 72 35-38 mc 37.5 0.92 1.81 5.17 -22.07 0.16 1.73 0.57 0.82 1.63 10.0 61 Major Element Data for Multicore (mc) JT96-01. Sample Depth AljOj FeA K20 MgO PA SiOj CaO MnO Ti02 Na,0 (cm) (%) (%) (%) (%) (%) (%) <%) (%) (%) (%) 0-1 mc 0.5 14.00 6.34 1.44 2.10 0.26 61.22 2.38 0.074 0.82 -0.28 1-2 mc 1.5 2-3 mc 2.5 3-4 mc 3.5 4-5 mc 4.5 5-6 mc 5.5 14.03 6.32 1.62 1.86 0.22 60.80 2.46 0.059 0.85 0.99 6-8 mc 7.0 8-10 mc 9.0 10-12 mc 11.0 13.02 6.28 1.64 1.84 0.20 57.51 2.42 0.074 0.83 1.15 12-14 mc 13.0 14-16 mc 15.0 14.93 6.27 1.69 2.25 0.21 64.14 2.46 0.063 0.86 2.04 16-18 mc 17.0 18-20 mc 19.0 15.00 6.43 1.74 2.24 0.20 64.03 2.43 0.061 0.87 2.11 20-25 mc 22.5 14.75 6.29 1.74 2.19 0.19 63.17 2.44 0.057 0.87 2.01 25-30 mc 27.5 15.16 6.33 1.74 2.43 0.19 64.07 2.62 0.065 0.87 2.44 30-35 mc 32.5 15.18 6.33 1.71 2.44 0.19 64.59 2.58 0.059 0.87 2.46 35-38 mc 37.5 15.16 6.35 1.71 2.47 0.19 64.94 2.63 0.064 0.88 2.56 Sample Depth V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br (cm) (ng/g) (ng/g) (ng/g) (ng/g) (ng/g) (ng/g) (ng/g) (ng/g) (Pg/g) (Pg/g) (ng/g) (ng/g) (ng/g) 0-1 mc 0.5 142 87 532 7 40 30 96 8 76 231 17 158 411 20 145 1-2 mc 1.5 2-3 mc 2.5 3-4 mc 3.5 4-5 mc 4.5 5-6 mc 5.5 148 92 449 11 38 29 97 5 70 238 18 169 430 19 111 6-8 mc 7.0 8-10 mc 9.0 10-12 mc 11.0 152 96 454 11 37 29 97 8 68 235 19 166 410 17 104 12-14 mc 13.0 14-16 mc 15.0 151 92 454 12 37 29 98 6 69 239 19 164 438 20 101 16-18 mc 17.0 18-20mc 19.0 146 86 418 12 38 30 102 7 67 234 19 168 415 19 99 20-25 mc 22.5 152 92 438 11 39 28 97 10 67 236 19 167 442 13 91 25-30 mc 27.5 147 94 430 12 39 28 94 11 69 246 20 174 408 14 71 30-35 mc 32.5 150 94 443 10 39 27 87 10 67 ' 242 19 169 452 13 70 35-38 mc 37.5 152 . 94 442 13 38 27 88 8 65 248 19 172 440 12 66 201 Table A9. (continued) Trace Metal Data for Multicore (mc) JT96-01. Sample Depth Ag Cd Re Mo U (cm) (ng/g) (Pg/g) (ng/g) (ug/g) (ug/g) 0-1 mc 0.5 84 0.15 2.13 0.62 1.12 1-2 mc 1.5 81 0.23 2.70 0.66 1.28 2-3 mc 2.5 94 0.20 2.91 0.72 1.26 3-4 mc 3.5 89 0.23 2.53 0.68 1.38 4-5 mc 4.5 97 0.23 3.50 0.74 1.58 5-6 mc 5.5 87 0.18 3.30 0.68 1.28 6-8 mc 7.0 94 0.25 2.34 0.81 1.32 8-10 mc 9.0 86 0.21 2.87 0.95 1.40 10-I2mc 11.0 85 0.23 3.40 1.02 1.44 12-14 mc 13.0 87 0.20 2.90 0.99 1.34 14-16 mc 15.0 79 0.23 3.87 1.22 1.31 16-18 mc 17.0 96 0.21 3.77 0.96 1.61 18-20 mc 19.0 86 0.20 3.29 0.95 1.49 20-25 mc 22.5 92 0.23 4.10 1.22 1.63 25-30 mc 27.5 85 0.20 5.32 1.30 1.84 30-35 mc 32.5 82 0.22 5.31 1.35 1.96 35-38 mc 37.5 75 0.25 5.91 1.42 2.12 202 Table A10. Geochemical Data for Multicore and Piston Core JT96-02. General Geochemical Data for Multicore (mc) and Piston Core (pc) JT96-02. Sample Depth Calendar Age cr 5"N SI3C0„ N,« C,„ S,„ CaCOj %C„, C„/N Opal Ba,,. (m) (kyrs) (wt. %) (%o vs air) (%o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) Wg) 0-1 mc 0.5 0.02 5.02 6.60 -21.50 0.40 3.49 0.38 0.83 3.39 8.5 8.8 378 l-2mc 1.5 0.06 4.24 0.39 3.33 0.35 0.83 3.23 8.4 2-3 mc 2.5 0.10 3.67 0.39 3.34 0.27 0.63 3.26 8.4 3-4 mc 3.5 0.14 3.23 0.39 3.29 0.27 0.54 3.23 8.3 4-5 mc 4.5 0.19 3.57 0.39 3.22 0.27 0.40 3.17 8.2 5-6 mc 5.5 0.23 3.38 6.24 -21.06 0.38 3.25 0.29 0.33 3.21 8.5 8.7 341 6-8 mc 7.0 0.29 2.67 0.37 3.14 0.25 0.33 3.10 8.5 8-10 mc 9.0 0.37 2.39 0.35 2.99 0.25 0.31 2.95 8.4 10-12 mc 11.0 0.45 2.35 6.84 -21.11 0.31 2.78 0.23 0.25 2.75 8.8 8.6 368 12-14 mc 13.0 0.53 2.55 0.33 2.86 0.25 0.32 2.82 8.6 14-16 mc 15.0 0.62 2.57 6.73 -21.09 0.29 2.56 0.28 0.33 2.52 8.8 9.9 371 16-18 mc 17.0 0.70 2.09 0.29 2.53 0.28 0.37 2.49 8.7 18-20mc 19.0 0.78 2.36 6.69 -21.14 0.28 2.48 0.32 0.42 2.43 8.8 8.1 345 Sample Depth Calendar Age ci" S"N 8"C„ N„ C, S,, CaC02 %C„, C^/N Opal Ba^ (m) (kyrs) (wt. %) (%o vs air) (%. vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (ng/g) 5-6 pc 5.5 2.50 6.84 -21.04 0.32 2.83 0.25 0.25 2.80 8.81 10-11 pc 10.5 2.17 6.62 -21.12 0.24 2.18 0.22 0.25 2.15 8.96 15-16 pc 15.5 2.19 6.99 -22.05 0.16 1.59 0.32 0.89 1.48 9.33 20-21 pc 20.5 1.70 6.83 -22.75 0.07 1.06 0.26 1.67 0.86 12.81 25-26 pc 25.5 1.66 6.84 -22.74 0.09 1.08 0.24 1.92 0.85 9.57 30-31 pc 30.5 1.64 6.40 -22.81 0.09 1.13 0.26 1.67 0.93 10.09 35-36 pc 35.5 1.20 6.05 -22.55 0.08 0.94 0.18 1.67 0.74 9.62 40-41 pc 40.5 1.07 6.68 -22.72 0.08 0.97 0.25 1.67 0.77 9.70 45-46 pc 45.5 0.44 -23.96 0.01 0.24 0.08 0.80 0.15 10.43 50-51 pc 50.5 0.71 -23.26 0.05 0.71 0.22 2.17 0.45 9.04 55-56 pc 55.5 0.57 -23.47 0.03 0.55 0.14 2.15 0.30 10.96 60-61 pc 60.5 0.48 -23.71 0.02 0.49 0.12 2.27 0.22 9.95 65-66 pc 65.5 0.73 -23.22 0.04 0.61 0.22 1.74 0.40 10.10 70-71 pc 70.5 0.48 -23.47 0.03 0.59 0.16 1.67 0.39 14.44 75-76 pc 75.5 0.50 -23.46 0.03 0.52 0.16 2.22 0.25 9.77 80-81 pc 80.5 0.47 -23.63 0.01 0.28 0.07 1.46 0.10 7.85 85-86 pc 85.5 0.42 -23.91 0.01 0.26 0.08 1.42 0.09 7.33 90-91 pc 90.5 0.96 5.08 -24.41 0.05 0.76 0.11 3.00 0.40 8.15 95-96 pc 95.5 0.89 -24.65 0.05 0.82 0.16 3.33 0.42 8.53 100-101 pc 100.5 0.97 -24.51 0.05 0.72 0.12 2.92 0.37 7.61 105-106 pc 105.5 0.83 -24.63 0.04 0.72 0.09 3.33 0.32 7.85 110-111 pc 110.5 1.01 5.31 -24.45 0.05 0.77 0.10 3.21 0.39 8.37 115-116 pc 115.5 0.81 -24.41 0.05 0.84 0.23 3.33 0.44 9.39 120-121 pc 120.5 0.90 -24.60 0.05 0.80 0.11 3.25 0.41 8.59 125-126 pc 125.5 0.86 -24.78 0.05 0.77 0.08 3.33 0.37 7.60 130-131 pc 130.5 0.98 4.75 -24.49 0.05 0.75 0.11 2.92 0.40 7.65 135-136 pc 135.5 1.04 -23.73 0.08 1.21 0.36 4.20 0.71 8.90 140-141 pc 140.5 0.99 -24.82 0.05 0.73 0.11 3.00 0.37 8.11 145-146 pc 145.5 0.95 -23.88 0.07 0.98 0.12 3.33 0.58 8.46 150-151 pc 150.5 0.79 5.17 -24.57 0.05 0.74 0.11 3.33 0.34 7.03 155-156 pc 155.5 1.20 -24.23 0.06 0.82 0.11 3.33 0.42 7.05 160-161 pc 160.5 1.16 -23.56 0.07 1.17 0.47 4.17 0.67 9.16 165-166 pc 165.5 0.82 -24.28 0.04 0.76 0.12 3.33 0.36 8.07 170-171 pc 170.5 0.99 5.45 -23.82 0.06 0.94 0.13 3.50 0.52 8.68 175-176pc 175.5 0.85 -24.53 0.05 0.76 0.10 3.33 0.36 7.52 180-181 pc 180.5 0.71 -24.47 0.04 0.72 0.10 3.17 0.34 9.44 185-186pc 185.5 0.79 -24.55 0.04 0.70 0.09 3.29 0.31 8.13 190-191 pc 190.5 0.91 5.50 -24.39 0.04 0.76 0.09 3.08 0.39 9.51 195-196 pc 195.5 0.97 -23.34 0.10 1.48 0.43 4.17 0.98 9.96 200-201 pc 200.5 0.88 -23.50 0.09 1.36 0.40 4.08 0.87 9.21 205-206 pc 205.5 0.86 -24.25 0.05 0.89 0.13 3.33 0.49 9.39 203 Table A10. (continued) Major Element Data for Multicore (mc) and Piston Core (pc) JT96-02. Sample Depth A1A FeA KjO MgO PA SiO, CaO MnO TiOj Na,0 (m) (%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 0-1 mc 0.5 12.36 6.57 1.72 2.77 0.26 50.27 2.58 0.066 0.76 1.00 1-2 mc 1.5 2-3 mc 2.5 3-4 mc 3.5 4-5 mc 4.5 5-6 mc 5.5 16.71 7.99 2.18 3.05 0.23 63.94 2.52 0.064 0.94 4.18 6-8 mc 7.0 8-10 mc 9.0 10-12 mc 11.0 14.72 6.61 2.26 2.76 0.21 59.15 2.37 0.054 0.84 4.00 12-14 mc 13.0 14-16 mc 15.0 13.05 6.48 2.07 2.74 0.19 53.46 2.29 0.063 0.84 1.48 16-18 mc 17.0 18-20mc 19.0 14.38 6.30 2.24 2.81 0.19 59.39 2.50 0.053 0.83 3.15 Sample Depth A1;0) Fe20, K.,0 MgO PA SiO; CaO MnO TA Na:0 (m) (%) (%) (%) (%) <%) (%) (%) (%) (%) (%) 5-6 pc 5.5 14.65 6.28 1.74 3.02 0.19 58.42 2.16 0.050 0.78 1.88 10-11 pc 10.5 11.89 6.84 2.03 2.83 0.18 52.46 2.45 0.052 0.78 0.55 15-16 pc 15.5 15.67 6.61 2.24 2.97 0.20 59.09 2.98 0.073 0.81 4.25 20-21 pc 20.5 15.56 6.62 2.17 3.06 0.20 59.49 3.83 0.083 0.79 3.74 25-26 pc 25.5 16.20 7.30 2.53 3.24 0.20 60.17 3.79 0.082 0.84 4.20 30-31 pc 30.5 16.55 6.94 2.51 3.04 0.20 62.55 3.92 0.082 0.82 5.08 35-36 pc 35.5 15.49 7.31 2.14 2.89 0.18 62.85 3.94 0.082 0.84 3.34 40-41 pc 40.5 15.52 5.84 1.93 2.81 0.18 61.63 3.69 0.071 0.76 3.99 45-46 pc 45.5 50-51 pc 50.5 55-56 pc 55.5 60-61 pc 60.5 65-66 pc 65.5 70-71 pc 70.5 75-76 pc 75.5 80-81 pc 80.5 85-86 pc 85.5 90-91 pc 90.5 16.14 7.16 2.84 2.94 0.23 56.41 3.62 0.081 1.02 3.56 95-96 pc 95.5 16.26 7.46 2.88 2.98 0.25 58.07 4.00 0.091 1.08 3.41 100-101 pc 100.5 14.76 7.53 3.04 2.89 0.26 54.24 3.77 0.081 1.10 2.14 105-106 pc 105.5 15.03 7.08 2.89 2.84 0.26 56.59 4.08 0.091 1.09 3.28 110-111 pc 110.5 15.69 7.08 2.83 3.06 0.24 56.03 3.42 0.090 0.97 2.53 115-116 pc 115.5 14.98 7.04 3.22 2.78 0.24 56.07 4.02 0.091 0.99 4.47 120-121 pc 120.5 16.57 7.35 3.20 2.92 0.25 58.93 3.71 0.091 1.05 4.12 125-126pc 125.5 13.75 7.16 2.84 2.81 0.24 51.85 3.80 0.091 1.03 1.48 130-131 pc 130.5 16.99 7.38 3.36 2.94 0.26 60.51 3.65 0.081 1.06 4.66 135-136 pc 135.5 16.38 7.33 2.33 3.51 0.22 57.08 4.39 0.100 0.92 3.19 140-141 pc 140.5 14.49 6.92 2.96 2.91 0.25 53.95 3.77 0.092 1.04 1.90 145-146 pc 145.5 16.36 7.34 2.99 3.06 0.23 58.26 3.93 0.081 0.99 4.39 150-151 pc 150.5 15.33 7.07 3.16 2.82 0.26 57.64 3.90 0.101 1.06 3.08 155-156 pc 155.5 16.33 7.19 3.07 2.98 0.26 57.26 3.52 0.082 0.99 3.53 160-161 pc 160.5 15.58 7.35 2.44 3.14 0.22 55.65 4.44 0.092 0.93 3.37 165-166 pc 165.5 16.85 6.93 2.85 3.15 0.24 59.77 3.63 0.100 0.98 3.36 170-171 pc 170.5 17.10 7.10 3.13 3.00 0.25 59.94 3.91 0.081 0.98 4.09 175-176pc 175.5 16.38 6.80 3.05 2.90 0.24 57.45 3.39 0.091 0.98 3.73 180-181 pc 180.5 14.56 6.78 2.95 2.79 0.25 57.36 4.08 0.091 1.02 2.82 185-186 pc 185.5 14.16 7.23 3.25 2.75 0.25 54.65 3.79 0.091 1.02 2.65 190-191 pc 190.5 16.34 7.37 3.35 2.88 0.26 57.84 3.67 0.092 1.04 3.89 195-196 pc 195.5 15.50 7.84 2.60 3.13 0.22 56.62 5.04 0.102 0.90 4.08 200-201 pc 200.5 15.97 6.87 2.12 2.98 0.22 57.76 4.53 0.091 0.86 4.42 205-206 pc 205.5 17.60 7.50 2.82 3.25 0.26 60.45 4.02 0.110 1.06 3.41 204 Table A10. (continued) Minor Element Data for Multicore (mc) and Piston Core (pc) JT96-02. Sample Depth V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br I (m) (pg/g) (ug/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (ng/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (ug/g) (pg/g) 0- 1 mc 0.5 141 113 521 20 58 46 126 99 227 16 134 687 24 135 968 1- 2 mc 1.5 2- 3 mc 2.5 3- 4 mc 3.5 4- 5 mc 4.5 5- 6mc 5.5 161 126 430 19 60 47 132 95 222 18 145 759 20 116 669 6- 8 mc 7.0 8-10 mc 9.0 10-12mc 11.0 167 130 446 20 59 44 129 86 228 19 146 736 18 99 472 12-14 mc 13.0 14-16 mc 15.0 164 131 451 20 59 45 132 86 230 20 154 697 17 75 347 16-18 mc 17.0 18-20mc 19.0 159 125 455 19 56 42 125 83 237 20 156 705 18 81 357 Sample Depth V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br I (m) (pg/g) (ug/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (pg/g) 5-6 pc 5.5 165 131 459 19 57 43 10-11 pc 10.5 170 143 499 20 53 39 15-16 pc 15.5 200 141 650 21 57 38 20-21 pc 20.5 196 129 711 21 50 34 25-26 pc 25.5 204 126 727 22 51 37 30-31 pc 30.5 198 131 740 21 50 35 35-36 pc 35.5 185 128 663 20 45 30 40-41 pc 40.5 182 119 669 18 44 28 45-46 pc 45.5 125 86 498 16 30 13 50-51 pc 50.5 148 102 616 23 39 25 55-56 pc 55.5 142 107 598 22 35 19 60-61 pc 60.5 138 109 596 23 33 16 65-66 pc 65.5 147 110 629 19 40 22 70-71 pc 70.5 126 96 559 28 27 11 75-76 pc 75.5 145 108 650 23 36 19 80-81 pc 80.5 118 83 503 22 26 11 85-86 pc 85.5 118 92 558 24 29 12 90-91 pc 90.5 199 81 740 23 31 32 95-96 pc 95.5 200 80 807 23 31 34 100-101 pc 100.5 205 76 733 25 32 35 105-106 pc 105.5 181 70 784 16 33 33 110-111 pc 110.5 169 76 703 17 34 35 115-116 pc 115.5 172 71 762 20 35 34 120-121 pc 120.5 172 74 690 18 34 34 125-126 pc 125.5 172 69 758 19 31 33 130-131 pc 130.5 173 75 659 22 32 34 135-136 pc 135.5 181 97 744 19 49 36 140-141 pc 140.5 174 72 755 20 32 36 145-146 pc 145.5 174 90 698 14 43 36 150-151 pc 150.5 171 73 757 16 34 33 155-156 pc 155.5 166 75 672 17 35 35 160-161 pc 160.5 178 91 764 22 46 39 165-166 pc 165.5 175 72 741 18 31 37 170-171 pc 170.5 169 87 702 14 43 36 175-176 pc 175.5 163 74 728 16 35 34 180-181 pc 180.5 166 74 742 17 33 34 185-186 pc 185.5 168 75 737 19 33 35 190-191 pc 190.5 172 72 723 20 33 35 195-196 pc 195.5 187 110 824 19 57 38 200-201 pc 200.5 179 113 810 18 52 37 205-206 pc 205.5 180 78 729 19 37 36 122 85 227 19 141 779 22 94 483 118 79 244 20 158 737 22 61 309 108 79 255 22 137 673 26 21 95 97 74 294 22 143 625 25 -3 36 101 80 286 23 131 637 23 -1 36 100 79 298 22 137 659 22 -4 37 93 70 303 23 156 613 21 7 65 89 68 306 21 154 630 20 6 . 38 51 2 343 17 128 510 16 17 14 77 4 318 20 158 562 14 23 25 67 2 325 21 184 528 14 18 13 63 1 333 21 196 506 16 17 16 75 5 332 20 166 559 15 23 25 51 3 364 17 151 513 15 15 22 68 8 346 21 196 532 13 17 16 49 4 361 16 146 491 13 15 18 51 4 357 17 158 533 12 14 21 107 126 224 34 181 664 30 -4 8 106 122 231 33 181 674 29 -4 15 110 123 229 34 179 647 27 -3 18 107 7 116 240 33 186 608 15 -5 16 108 6 133 220 34 176 614 23 -7 9 107 6 122 232 32 182 615 19 -5 16 109 6 133 223 35 183 595 23 -5 3 110 7 128 229 35 194 619 20 -8 5 109 7 128 222 33 182 576 20 -6 14 105 21 102 261 31 165 626 15 1 20 107 5 126 225 34 188 624 23 -10 8 105 5 115 239 32 169 611 22 -3 14 103 4 123 224 34 187 604 25 -6 14 108 4 132 216 35 177 613 28 -16 11 104 13 106 257 30 170 587 17 -7 8 106 4 124 222 34 182 635 22 -5 12 106 11 126 232 33 177 624 14 -6 15 105 5 132 215 34 180 628 21 -5 11 100 7 115 240 33 203 590 15 -3 10 105 4 125 226 35 191 594 24 -4 10 105 7 129 217 36 182 591 19 -4 -3 101 21 83 289 27 157 603 12 3 11 100 8 78 286 25 145 582 14 2 10 111 7 117 237 34 178 586 16 -3 8 205 Table A10. (continued) Trace Metal Data for Multicore (mc) and Piston Core (pc) JT96-02. Sample Depth Ag Cd Re Mo U (m) (ng/g) (ug/g) (ng/g) (pg/g) (pg/g) 0-1 mc 0.5 330 0.21 3.53 0.69 1.21 1-2 mc 1.5 335 0.20 3.92 0.65 1.27 2-3 mc 2.5 420 0.23 4.20 0.68 1.36 3-4 mc 3.5 462 0.30 4.02 0.73 1.54 4-5 mc 4.5 458 0.24 4.16 0.67 1.54 5-6 mc 5.5 475 0.37 4.57 0.68 1.62 6-8 mc 7.0 476 0.51 5.29 0.69 1.80 8-10 mc 9.0 473 0.47 8.54 0.67 1.74 10-12 mc 11.0 430 0.43 11.06 0.61 1.64 12-14 mc 13.0 428 0.59 15.12 0.90 1.79 14-16 mc 15.0 434 0.64 25.02 0.85 2.11 16-18 mc 17.0 422 0.61 30.18 0.82 2.83 18-20 mc 19.0 384 0.45 16.89 0.86 2.83 Table A10. (continued) General geochemical data for the sulphide layers (in bold print) and intervening sediments, Piston Core JT96-02. Sample Depth (cm) $I5N S1^ Ntot1 (permil vs air) (permil vs VPDB) (wt.%) 1 134.15 4.79 -24.57 0.033 2 134.45 4.61 -25.09 0.033 3 134.75 4.94 -24.26 0.039 4 135.05 5.40 -23.46 0.063 5 135.35 530 -23.49 0.065 6 135.65 5.65 -23.61 0.059 7 135.95 5.24 -23.57 0.056 8 136.25 5.25 -23.64 0.056 9 136.55 5.62 -23.75 0.049 10 136.85 5.15 -23.87 0.047 11 137.15 4.98 -23.80 0.045 12 137.45 5.02 -23.90 0.039 13 137.75 4.96 -24.22 0.033 14 138.05 0.031 15 138.35 0.030 16 138.65 0.030 17 138.95 0.031 18 139.25 4.77 -24.78 0.032 19 139.55 0.032 20 139.85 0.030 21 140.15 0.032 22 140.45 -24.54 0.030 23 140.75 4.71 0.029 24 141.05 0.030 25 141.35 0.029 26 141.65 0.030 27 141.95 4.73 -24.48 0.062 28 142.25 4.91 -24.49 0.028 29 142.55 4.50 -24.40 0.031 30 142.85 4.27 -24.58 0.028 31 143.15 4.72 -24.45 0.029 32 143.45 4.45 -24.15 0.037 33 143.75 5.56 -23.39 0.063 34 144.05 5.81 -2338 0.062 35 144.35 5.56 -23.40 0.058 36 144.65 5.48 -23.56 0.057 37 144.95 5.20 -23.44 0.053 38 145.25 5.08 -23.76 0.051 39 145.55 5.40 -23.57 0.047 40 145.85 5.02 -23.84 0.043 41 146.15 4.94 -23.96 0.037 42 146.45 5.01 -24.14 0.037 43 146.75 4.61 -24.86 0.034 44 147.05 0.027 45 147.35 0.033 46 147.65 0.029 47 147.95 4.53 -24.51 0.031 48 148.25 0.031 49 148.55 0.032 50 148.85 4.76 -24.45 0.031 51 149.15 0.032 52 149.45 0.029 53 149.75 0.034 54 150.05 0.029 55 150.35 0.038 56 150.65 3.43 0.037 57 150.95 0.038 58 151.25 0.047 59 151.55 0.046 60 151.85 0.042 61 152.15 0.060 62 152.45 0.075 63 152.75 0.071 64 152.05 0.067 65 15335 0.065 66 153.65 0.061 67 153.95 0.056 68 154.25 0.055 69 154.55 0.048 70 154.85 0.044 Ctot1 Stot' %Ccarb %carb %Corg Corg/N (wt.%) (wt.%) (wt.%) 0.679 0.096 0.415 3.46 0.264 7.95 0.689 0.098 0.411 3.42 0.278 8.41 0.726 0.285 0.415 3.46 0.311 8.01 1.117 0.379 0.546 4.55 0.571 9.02 1.128 0.440 0.540 4.50 0.588 8.99 1.035 0.418 0.498 4.15 0.537 9.06 0.989 0.297 0.485 4.04 0.504 9.00 0.970 0.191 0.483 4.02 0.487 8.72 0.904 0.138 0.466 3.88 0.438 8.88 0.841 0.112 0.436 3.63 0.405 8.63 0.810 0.113 0389 3.24 0.421 9.30 0.702 0.092 0.358 2.98 0.344 8.81 0.642 0.086 0.356 2.97 0.286 8.72 0.629 0.073 0.368 3.07 0.261 8.36 0.617 0.072 0.390 3.25 0.227 7.51 0.637 0.080 0.367 3.06 0.270 8.87 0.651 0.076 0.386 3.22 0.265 8.46 0.646 0.075 0.374 3.12 0.272 8.54 0.648 0.083 0.396 3.30 0.252 7.93 0.653 0.079 0.397 3.31 0.256 8.64 0.652 0.075 0.404 3.37 0.248 7.80 0.645 0.078 0.393 3.27 0.252 8.49 0.644 0.079 0.374 3.12 0.270 9.23 0.643 0.077 0385 3.21 0.258 8.69 0.645 0.074 0.419 3.49 0.226 7.77 0.646 0.074 0.413 3.44 0.233 7.74 0.645 0.082 0.393 3.27 0.252 4.07 0.646 0.078 0.399 3.32 0.247 8.82 0.651 0.103 0.423 3.52 0.228 7.47 0.651 0.074 0.414 3.45 0.237 8.49 0.667 0.078 0.440 3.67 0.227 7.82 0.730 0.232 0.421 3.51 0.309 8.32 1.098 0.460 0.538 4.48 0.560 8.84 1.058 0.283 0.483 4.02 0.575 9.22 0.996 0.258 0.495 4.12 0.501 8.66 0.979 0.269 0.461 3.84 0.518 9.09 0.951 0.199 0.471 3.92 0.480 9.06 0.883 0.126 0.425 3.54 0.458 9.00 0.842 0.116 0.411 3.42 0.431 9.12 0.751 0.123 0372 3.10 0.379 8.86 0.686 0.091 0.380 3.17 0.306 8.29 0.682 0.079 0.374 3.12 0.308 8.40 0.629 0.079 0.361 3.01 0.268 7.77 0.506 0.069 0.342 2.85 0.164 6.06 0.619 0.079 0.355 2.96 0.264 8.05 0.613 0.076 0.350 2.92 0.263 9.16 0.623 0.080 0.365 3.04 0.258 8.37 0.631 0.083 0.372 3.10 0.259 8.32 0.640 0.081 0.392 3.27 0.248 7.77 0.642 0.073 0.386 3.22 0.256 8.19 0.652 0.075 0.399 3.32 0.253 7.92 0.639 0.077 0.389 3.24 0.250 8.48 0.647 0.087 0.402 3.35 0.245 7.28 0.639 0.075 0.385 3.21 0.254 8.63 0.694 0.042 0.414 3.45 0.280 7.29 0.667 0.043 0.388 3.23 0.279 7.59 0.710 0.052 0.415 3.46 0.295 7.80 0.851 0.090 0.448 3.73 0.403 8.55 0.823 0.085 0.461 3.84 0.362 7.86 0.787 0.264 0.449 3.74 0338 8.07 1.039 0.589 0.494 4.12 0.545 9.02 1.227 0381 0.556 4.63 0.671 8.98 1.185 0.280 0.559 4.66 0.626 8.82 1.171 0.239 0.532 4.43 0.639 9.48 1.139 0.218 0.546 4.55 0.593 9.13 1.082 0.191 0.507 4.22 0.575 9.39 1.001 0.155 0.499 4.16 0.502 8.98 0.965 0.131 0.468 3.90 0.497 8.97 0.861 0.143 0.446 3.72 0.415 8.66 0.812 0.151 0.436 3.63 0376 8.56 1 Data have not been salt corrected. 207 Table A10. (continued) Trace metal1 and major element data1 for the sulphide layers (in bold) and intervening sediments, Piston Core JT96-02. Sample Depth (cm) Ag Cd Re Mo U A1:0, Fe20, KjO MgO Na:0 PjO, SiOj CaO MnA TiO, (ng/g) (ug/g) (ng/g) (ng/g) (ug/g) wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% wt.% 1 134.15 102 0.25 1.26 1.03 1.76 16.09 7.05 2.85 3.14 5.21 0.253 57.17 3.61 0.1 1.02 2 134.45 112 0.20 0.85 0.87 1.67 16.25 7.13 2.98 3.19 4.82 0.253 57.24 3.58 0.101 1.02 3 134.75 106 0.15 2.04 1.40 1.63 16 7.39 2.96 3.18 5.12 0.245 56.54 3.65 0.1 1 4 135.05 108 0.18 1.65 2.05 1.51 15.63 7.53 232 3.37 5.17 0.222 55.06 4.57 0.1 0.9 5 135.35 113 0.28 1.38 2.11 1.56 15.53 7.5 237 3.4 539 0.22 55.49 4.57 0.097 0.89 6 135.65 109 0.19 0.93 1.71 1.53 7 135.95 115 0.15 1.55 1.52 1.59 15.69 7.32 2.52 3.28 5.77 0.223 55.89 4.25 0.092 0.93 8 136.25 114 0.14 1.18 1.00 1.78 15.59 7.12 2.63 3.3 5.12 0.227 56.11 4.23 0.094 0.93 9 136.55 100 0.23 1.31 0.86 1.63 10 136.85 105 0.14 2.21 0.75 1.63 16 7.18 2.57 3.28 5.11 0.232 56.65 3.81 0.089 0.96 11 137.15 98 0.15 1.11 0.78 1.7 12 137.45 75 0.20 1.94 0.79 1.71 13 137.75 80 0.30 0.74 0.79 1.68 16.18 7.27 2.93 3.12 5.28 0.243 57.64 3.38 0.086 1.01 14 138.05 82 0.13 0.64 0.73 1.69 16.27 7.03 2.96 3.11 6.96 0.245 57.66 3.38 0.091 1.01 15 138.35 88 0.10 0.68 0.70 1.95 16 138.65 89 0.12 0.56 1.05 1.85 16.28 7.1 2.96 3.11 5.81 0.249 58.01 3.44 0.092 1.01 17 138.95 93 0.14 0.84 0.72 1.69 18 139.25 82 0.12 0.43 0.76 1.69 19 139.55 81 0.19 0.34 0.81 1.83 16.22 7.09 2.96 3.13 6.16 0.249 57.59 3.48 0.095 1.01 20 139.85 99 0.09 0.51 0.79 1.74 21 140.15 108 0.12 0.76 1.67 22 140.45 93 0.16 0.84 0.79 1.96 15.92 7.14 2.79 3.12 4.73 • 0.245 57.8 3.58 0.096 1.01 23 140.75 92 0.26 0.61 0.77 2.01 15.91 7.09 2.85 3.07 4.85 0.25 58.3 3.66 0.098 1.02 24 141.05 91 0.26 0.63 0.75 1.76 25 141.35 90 0.42 0.61 0.76 2.18 15.84 7.05 2.83 3.06 5.28 0.251 58.12 3.67 0.098 1.02 26 141.65 96 0.72 0.77 1.91 27 141.95 77 0.39 0.65 0.79 1.87 28 142.25 83 0.20 0.78 0.61 1.8 15.49 7.02 2.63 3.07 4.37 0.251 59.11 3.84 0.099 1.04 29 142.55 92 0.20 0.73 0.83 1.91 30 142.85 91 0.29 0.98 0.75 1.9 15.45 6.96 2.64 3.07 6.96 0.249 59.09 3.92 0.1 1.05 31 143.15 95 0.33 1.41 1.66 1.93 32 143.45 116 0.34 0.84 1.35 2.11 15.82 7.33 2.65 3.16 4.86 0.237 57.39 3.78 0.099 1 33 143.75 112 0.15 1.76 2.27 1.84 15.6 7.63 2.29 3.4 5.8 0.218 55.34 4.47 0.099 0.89 34 144.05 117 0.41 1.76 2.52 1.7 16.34 7.69 2.49 3.49 5.65 0.225 58.46 4.52 0.098 0.92 35 144.35 112 035 1.55 2.18 1.78 36 144.65 118 0.27 1.13 2.00 1.74 15.81 7.28 2.47 3.34 5.57 0.218 56.15 4.12 0.092 0.92 37 144.95 116 0.38 1.31 1.40 1.82 38 145.25 95 0.25 1.08 0.86 1.96 16.15 7.07 2.6 3.34 5.52 0.224 57.01 3.86 0.091 0.94 39 145.55 91 0.39 0.97 0.79 1.86 40 145.85 83 0.23 0.89 0.85 2.03 16.29 7.24 2.77 3.25 4.96 0.228 57.13 3.46 0.086 0.95 41 146.15 77 0.17 0.94 0.74 2.05 42 146.45 66 0.12 1.44 0.67 2.01 16.47 7.01 2.9 3.21 5.26 0.238 57.82 3.39 0.088 0.97 43 146.75 85 0.14 0.71 0.73 2.05 44 147.05 65 0.11 0.44 0.78 1.9 45 147.35 71 0.13 0.54 0.64 1.81 16.39 7.09 2.96 3.11 6.87 0.242 57.86 3.26 0.089 0.99 46 147.65 80 0.21 0.73 0.69 1.92 16.26 7.03 3.01 3.12 4.91 0.24 57.87 3.34 0.088 0.99 47 147.95 73 0.27 0.51 0.71 1.94 48 148.25 70 0.13 0.53 0.71 1.91 16.28 6.98 2.97 3.12 5.63 0.243 57.93 3.41 0.092 1 49 148.55 74 0.18 0.94 0.70 2.22 50 148.85 81 0.32 0.66 0.71 1.99 51 149.15 82 0.30 0.58 0.72 1.98 16.22 7 2.93 3.13 4.94 0.245 57.92 3.48 0.095 0.99 52 149.45 82 0.23 0.60 0.73 2.04 53 149.75 94 0.24 0.71 1.99 54 150.05 91 0.39 0.89 0.77 1.82 16.28 7.01 2.91 3.09 4.98 0.241 57.87 3.47 0.098 1 55 150.35 97 0.19 1.76 0.74 1.89 56 150.65 98 0.15 0.62 0.83 1.61 15.85 6.97 2.97 3.05 5.03 0.249 57.93 3.62 0.099 1.01 57 150.95 86 0.14 0.47 0.72 1.5 58 151.25 97 0.16 0.66 0.93 1.5 15.76 7.04 2.62 3.19 5.15 0.239 58.02 4.14 0.096 0.99 59 151.55 99 0.15 0.82 0.85 1.64 60 151.85 114 0.17 0.82 1.37 1.7 15.66 7.26 2.62 3.15 4.85 0.242 57.96 3.91 0.099 1 61 152.15 117 0.17 1.21 2.30 1.47 15.61 7.89 2.51 3.25 5.68 0.224 55.63 4.16 0.101 0.92 62 152.45 103 0.20 1.37 1.95 139 63 152.75 97 0.22 1.21 1.57 1.53 15.77 7.21 2.31 3.43 5.27 0.214 55.95 4.54 0.097 0.89 64 152.05 96 0.20 1.42 1.59 1.62 65 15335 98 0.17 1.21 1.31 1.59 66 153.65 96 0.18 0.83 1.03 1.6 16.01 7.15 2.48 332 5.48 0.223 56.22 4.24 0.094 0.93 67 153.95 90 0.13 0.79 0.98 1.63 68 154.25 84 0.11 0.71 0.73 1.6 69 154.55 75 0.14 0.73 0.71 1.62 16.01 7.08 2.88 3.24 5.19 0.237 56.75 3.76 0.089 0.95 70 154.85 71 0.16 0.98 0.68 1.74 16.48 7.1 2.84 3.23 5.01 0.24 57.33 3.68 0.089 0.97 1 Data have not been salt corrected. 208 Table A10. (continued) Minor element data ' for the sulphide layers (in bold print) and intervening sediments. Piston Core JT96-02. Sample Depth (cm) V Cr Mn Co Ni Cu Zn Rb Sr Y Zr Ba Pb Nb (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) (ppm) 1 134.15 184.4 76.1 799.2 33.6 28.6 32.8 100.2 99.4 214 30.9 169.1 662.3 27.4 13.9 2 134.45 170.9 67.3 727.4 29.7 29 30.6 96.3 97.3 206.1 29.1 162.8 579.9 21.2 13.4 3 134.75 180.6 75.9 767.2 34.6 31.6 30.2 102.4 97.7 209.7 27.2 161.4 668.1 27.1 13.2 4 135.05 205.9 113.7 870.2 41.8 51.5 39.8 105.5 78.2 260 26.9 145.1 784.6 37 13.3 5 135.35 190.5 104.8 758.5 37.6 46.6 32.9 92.6 71.2 249.2 25.1 142 654.3 39.3 13.7 6 135.65 7 135.95 191.8 98.3 754J 35.6 43 36.3 102.2 86.9 254.6 27.4 161.4 688 47.9 12.9 8 136.25 195 101.9 763 34.4 40.3 33.7 102.2 87.6 251.5 27.9 158.4 701.5 40 13 9 136.55 10 136.85 181.7 84 679 32 32.3 29.2 90.8 86.2 222.2 27.5 155.1 616.4 21.3 11 11 137.15 12 137.45 13 137.75 189.5 79.6 708.7 34.8 28.5 28.2 99 101.9 211.2 29.3 166.9 721.6 19.1 14.6 14 138.05 189.9 81.1 738.3 36.1 30.2 35.5 108.6 113.6 227.7 32.6 185.6 755.3 27.2 14.7 15 138.35 16 138.65 181.2 71 712.7 32.9 27.2 29.9 94.7 101.4 209.3 28.4 170.1 646 21.6 13.7 17 138.95 18 139.25 19 139.55 177.6 71.3 717.9 33.3 26.3 28.8 91.4 99.4 203.9 28.4 164.4 640.9 26.3 14.2 20 139.85 21 140.15 22 140.45 190.3 84.8 799.8 36.7 29.6 33.8 98 101.3 220.7 31.9 173 726.7 21.3 17.4 23 140.75 175.8 64.8 716.7 30.6 26.8 33.6 91.6 98 217.1 30.6 172.7 618.3 20.4 13.1 24 141.05 25 141.35 180.4 71.4 760 31.8 28.7 31.5 96.9 100.1 224.2 31.2 180.3 647.8 28.3 14.3 26 141.65 27 141.95 28 142.25 182.5 70.4 775.6 31.8 27.9 29.5 94.5 96 233.7 29.8 184.1 639.4 31.8 14.6 29 142.55 30 142.85 187.9 71.7 784.1 32.5 28.6 30.7 94.6 93.6 235.5 29 189.1 628.9 19.1 15.3 31 143.15 32 143.45 174.9 70.4 731.6 31.6 31.6 31.5 88.7 91.5 216.3 27.7 165.7 594.1 22.5 13.5 33 143.75 218.6 125.7 897.1 46.6 49.2 35 99.7 76.8 258 27.6 147.8 865.8 44.1 11.6 34 144.05 181.8 101.3 733.6 34.4 44.6 34.4 93.1 76.8 252.3 25.4 150.9 637.6 21.3 12.5 35 144.35 36 144.65 191.5 105.9 771.9 36.7 42.8 36.1 97.1 84.6 254.1 26.6 161.1 728 23.8 11.5 37 144.95 38 145.25 182.2 89.4 698J 33 34.5 28.6 85.1 83.3 220.6 25.8 145.8 639.6 23.9 12.4 39 145.55 40 145.85 175.5 78.9 656.8 32.2 31.7 28.3 87.4 94 210 26.6 156.7 634.1 20.6 12.5 41 146.15 42 146.45 174.6 73.5 659.3 31 29.9 29.1 91.2 102.6 208.6 28.7 162.3 621.2 27.8 13 43 146.75 44 147.05 45 147.35 173.6 69.8 655.6 31.5 27.9 27.9 90.8 101.3 204.1 31.4 161.7 639 28.8 14.7 46 147.65 174.7 71.4 684.8 32.5 29 31.2 97,7 112.1 215.2 33.4 176.5 641.9 26.9 13.9 47 147.95 48 148.25 176.9 71 708.4 32.7 26.3 27.7 93.2 104.3 206.6 27.3 167.5 633.8 24.4 13.8 49 148.55 50 148.85 51 149.15 167.7 69 698.4 30.8 26.1 26.2 85.7 95.5 200.7 29.2 162.5 611.2 18.5 12.6 52 149.45 53 149.75 54 150.05 169.7 69 733.3 31.4 26.5 28.3 91.2 101.4 205.4 30.8 166.9 624.5 29.1 14 55 150.35 56 150.65 182.9 72.5 760.7 32.3 27.7 29 94.4 102.1 218.3 30.3 176.3 675.5 24.4 14.1 57 150.95 58 151.25 185.7 82.6 767 33.9 32.5 32.3 89.4 88.5 235.8 28.1 171.3 661.9 23.1 13.8 59 151.55 60 151.85 186.2 82 777.5 35.5 31.8 31.6 94.5 90.4 229.1 28.9 176.1 659.6 25.8 13.7 61 152.15 187.8 97.1 804.9 41.9 41.5 33.3 95.5 85 235.4 26.7 151.5 728 21.8 11.3 62 152.45 63 152.75 188.3 100.3 758.2 34.4 42.1 33.6 87.2 73.4 248.6 24.6 139.9 655.7 30.2 10.1 64 152.05 65 153.35 66 153.65 184.4 95.1 736.1 33.8 37.2 35.5 91.6 81.9 234 25.7 151.6 655.7 34.8 11.8 67 153.95 68 154.25 69 154.55 174.8 81.6 683 32.4 31.5 27.9 89.6 90.6 213.7 26.9 150.9 637.8 30.2 12.9 70 154.85 179.2 80.7 696.4 33.4 31 29.4 93.2 97.9 214.8 27.5 161 655.2 18.5 13.4 1 Data have not been salt corrected. 209 Table Al 1 Geochemical Data Multicore JT96-04. General Geochemical Data for Multicore (mc) JT96-04. Sample Depth Calendar Age cr 8"N 8"C„ N„ C s,» CaCOj C^N Opal Baiu, (m) (kyrs) (wt. %) (%o vs air) (%o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (pg/g) 0-1 mc 0.5 0.69 -20.49 0.06 0.54 0.04 0.27 0.51 8.2 2.5 9 1-2 mc 1.5 0.64 0.06 0.53 0.03 0.17 0.51 8.4 2-3 mc 2.5 0.62 0.07 0.54 0.04 0.20 0.52 7.6 3-4 mc 3.5 0.58 0.06 0.49 0.03 0.22 0.46 7.6 4-5 mc 4.5 0.72 0.06 0.51 0.03 0.17 0.49 7.9 5-6 mc 5.5 0.66 -20.51 0.06 0.51 0.03 0.22 0.48 7.9 3.2 9 6-7 mc 6.5 0.57 0.07 0.57 0.05 0.19 0.55 8.2 7-8 mc 7.5 0.53 0.07 0.55 0.03 0.22 0.53 7.7 8-9 mc 8.5 0.62 0.07 0.59 0.04 0.32 0.55 8.0 9-10 mc 9.5 0.66 0.07 0.59 0.07 0.28 0.56 8.5 10-11 mc 10.5 0.74 -22.40 0.06 0.56 0.15 0.41 0.51 8.8 3.2 59 ll-12mc 11.5 0.82 0.04 0.60 0.35 1.23 0.45 10.2 12-13 mc 12.5 0.74 0.05 0.55 0.28 0.42 0.50 9.5 13-14 mc 13.5 0.80 0.06 0.61 0.25 0.52 0.54 9.3 14-16 mc 15.0 0.79 -23.60 0.05 0.58 0.29 0.94 0.46 9.5 2.9 16-18 mc 17.0 0.68 0.06 0.57 0.13 0.40 0.52 8.9 Major Element Data for Multicore (mc) JT96-04. Sample Depth AljOj Fe20, K,0 MgO PA SiO, CaO MnO TiOj Na30 (m) <%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 0-1 mc 0.5 11.70 8.18 1.63 2.27 0.15 66.37 3.07 0.060 0.56 5.94 1-2 mc 1.5 2-3 mc 2.5 3-4 mc 3.5 4-5 mc 4.5 5-6 mc 5.5 11.93 8.11 1.67 2.31 0.14 67.22 3.14 0.061 0.58 7.54 6-7 mc 6.5 7-8 mc 7.5 8-9 mc 8.5 9-10 mc 9.5 10-11 mc 10.5 14.33 7.61 1.77 2.86 0.17 63.55 2.95 0.079 0.72 8.27 ll-12mc 11.5 12-13 mc 12.5 13-14 mc 13.5 14-16mc 15.0 15.87 6.84 1.74 3.10 0.19 61.13 3.21 0.091 0.79 7.96 Minor Element Data for Multicore (mc) JT96-04. Sample Depth V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br (m) (Pg/g) (pg/g) (pg/g) (Pg/g) (Pg/g) (pg/g) (pg/g) (pg/g) (Pg/g) (pg/g) (pg/g) (pg/g) (ug/g) (Pg/g) (Pg/g) 0-1 mc 0.5 110 151 404 29 25 16 64 3 54 307 15 129 291 16 55 1-2 mc 1.5 2-3 mc 2.5 3-4 mc 3.5 4-5 mc 4.5 5-6 mc 5.5 112 133 397 26 26 15 62 2 53 307 16 146 293 14 41 6-7 mc 6.5 7-8 mc 7.5 8-9 mc 8.5 9-10 mc 9.5 10-11 mc 10.5 164 122 550 27 38 27 83 3 60 283 20 144 397 17 36 ll-12mc 11.5 12-13 mc 12.5 13-14 mc 13.5 14-16 mc 15.0 187 118 672 22 45 32 89 57 288 22 144 450 20 32 16-18 mc 17.0 210 Table Al 1. (continued) Trace Metal Data for Multicore (mc) JT96-04. Sample Depth Ag Cd Re Mo U (m) (ng/g) (Pg/g) (ng/g) (ug/g) (llg/g) 0-1 mc 0.5 n.a. 0.09 1.58 0.39 n.a. 1-2 mc 1.5 32 0.11 1.43 0.30 0.92 2-3 mc 2.5 n.a. 0.09 1.60 0.42 0.73 3-4 mc 3.5 27 0.08 1.37 0.35 0.71 4-5 mc 4.5 27 0.07 1.78 0.28 0.65 5-6 mc 5.5 32 0.10 1.66 0.34 0.92 6-7 mc 6.5 27 0.09 1.76 0.28 0.75 7-8 mc 7.5 26 0.10 2.72 0.31 0.88 8-9 mc 8.5 33 0.15 4.17 0.37 1.16 9-10 mc 9.5 47 0.25 4.59 0.40 1.14 10-11 mc 10.5 70 0.30 5.19 0.61 1.41 ll-12mc 11.5 79 0.17 1.91 0.80 1.87 12-13 mc 12.5 77 0.31 3.27 0.83 1.51 13-14 mc 13.5 76 0.24 3.55 0.85 1.36 14-16 mc 15.0 81 0.18 2.50 0.73 1.61 16-18 mc 17.0 62 0.25 5.06 0.47 1.04 Table A12. Geochemical Data for Box Core and Trigger Core JT96-05. General Geochemical Data for Box Core (be) and Trigger Core (tc) JT96-05. Sample Depth Calendar Age cr 5"N 8"C™ s«. CaCOj %C0„ C„/N Opal Ba^ (m) (kyrs) (wt. %) (%> vs air) (%o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (pg/g) 0-1 be 0.5 0.109 2.48 6.29 -21.27 0.20 1.71 0.21 0.62 1.64 8.3 6.5 560 1-2 be 1.5 0.328 1.94 0.24 2.00 0.19 0.57 1.94 8.2 556 2-3 be 2.5 0.547 1.97 0.23 1.94 0.18 0.55 1.88 8.2 559 3-4 be 3.5 0.766 1.78 0.21 1.79 0.16 0.46 1.74 8.2 4-5 be 4.5 0.985 1.75 0.20 1.70 0.16 0.28 1.67 8.3 5-6 be 5.5 1.204 1.84 6.55 -20.99 0.21 1.79 0.17 0.32 1.75 8.4 7.1 677 6-7 be 6.5 1.423 1.80 0.22 1.85 0.16 0.29 1.82 8.3 7-8 be 7.5 1.642 1.69 0.22 1.84 0.16 0.29 1.81 8.3 8-9 be 8.5 1.861 1.78 0.22 1.86 0.18 0.27 1.83 8.3 9-10 be 9.5 2.080 1.76 0.21 1.78 0.16 0.35 1.73 8.4 10-12 be 11.0 2.408 1.77 6.72 -21.05 0.21 1.83 0.17 0.37 1.78 8.4 6.7 664 12-14bc 13.0 2.846 1.59 0.21 1.78 0.15 0.27 1.75 8.4 14-16 be 15.0 3.284 1.60 6.48 -21.07 0.20 1.74 0.15 0.25 1.71 8.4 6.5 665 16-18 be 17.0 3.722 ' 1.41 0.20 1.73 0.13 0.22 1.70 8.6 18-20 be 19.0 4.160 1.36 0.20 1.74 0.14 0.29 1.71 8.5 20-22 be 21.0 4.598 1.32 6.86 -21.16 0.18 1.60 0.13 0.35 1.56 8.7 5.9 584 22-24 be 23.0 5.036 1.37 0.20 1.73 0.14 0.31 1.70 8.5 24-26 be 25.0 5.473 1.27 6.81 -21.26 0.16 1.47 0.12 0.30 1.43 8.7 5.5 552 26-28 be 27.0 5.911 1.24 0.18 1.60 0.12 0.31 1.56 8.7 28-30 be 29.0 6.349 1.15 0.15 1.35 0.10 0.43 1.30 8.8 30-32 be 31.0 6.787 1.18 6.95 -21.46 0.15 1.41 0.12 0.59 1.34 8.8 5.4 513 32-34 be 33.0 7.225 1.21 0.15 1.46 0.12 0.77 1.37 8.9 34-36 be 35.0 7.663 1.28 7.23 -21.62 0.16 1.51 0.12 1.11 1.38 8.7 4.6 496 36-38 be 37.0 8.101 1.26 0.16 1.61 0.12 1.32 1.45 8.9 38-40 be 39.0 8.539 1.23 0.15 1.64 0.12 2.42 1.35 8.9 40-42 be 41.0 8.976 1.29 7.94 -21.69 0.16 1.74 0.13 2.68 1.42 8.8 5.1 456 42-44 be 43.0 9.414 1.28 0.16 1.76 0.13 2.95 1.40 8.9 44-46 be 45.0 9.852 1.18 7.75 -21.90 0.14 1.65 0.20 3.08 1.28 9.0 4.8 385 46-48 be 47.0 10.290 1.12 0.14 1.61 0.24 2.56 1.30 9.6 Sample Depth Calendar Age cr 8"N S"C„„ N„ " C" s„ ** CaC02 %c„ C„/N Opal Ban. (m) (kyrs) (wt. %) (%, vs air) ( %o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (Pg/g) 11-12 tc 11.5 0.25 2.13 0.20 0.73 21-22 tc 21.5 0.28 2.51 0.30 0.71 31-32 tc 31.5 0.29 2.57 0.33 0.75 41-42 tc 41.5 0.30 2.60 0.45 0.79 51-52 tc 51.5 0.26 2.34 0.40 0.82 61-62 tc 61.5 0.24 2.14 0.30 0.60 71-72 tc 71.5 0.22 1.99 0.33 0.58 81-82 tc 81.5 0.20 1.80 0.18 0.54 91-92 tc 91.5 0.19 1.75 0.24 1.00 101-102 tc 101.5 0.17 1.67 0.20 1.54 111 -112 tc 111.5 0.17 1.90 0.18 3.53 121-122 tc 121.5 0.16 1.89 0.29 3.88 131-132tc 131.5 0.14 1.69 0.40 2.77 141-142 tc 141.5 0.10 1.28 0.36 3.12 Data have not been salt corrected. 212 Table A12. (continued) Major Element Data for Box Core (be) JT96-05. Sample Depth AljO, Fe,0, K,0 MgO PA Si02 CaO MnO Ti02 Na;0 (m) (%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 0-1 be 0.5 14.40 5.70 1.52 2.28 0.17 63.83 3.10 0.085 0.67 1.33 1-2 be 1.5 14.18 5.61 1.51 2.66 0.18 61.64 2.89 0.070 0.70 5.92 2-3 be 2.5 13.71 5.61 1.56 2.56 0.17 60.64 2.87 0.060 0.71 10.60 3-4 be 3.5 4-5 be 4.5 5-6 be 5.5 15.21 6.24 1.73 2.86 0.18 64.14 2.74 0.068 0.73 2.56 6-7 be 6.5 7-8 be 7.5 8-9 be 8.5 9-10 be 9.5 10-12 be 11.0 15.29 6.04 1.76 2.89 0.17 64.10 2.76 0.057 0.73 2.62 12-14 be 13.0 14-16 be 15.0 15.21 6.21 1.84 2.65 0.16 62.94 2.65 0.065 0.74 2.28 16-18 be 17.0 18-20 be 19.0 20-22 be 21.0 15.45 6.02 1.84 2.59 0.16 63.58 2.72 0.067 0.76 2.23 22-24 be 23.0 24-26 be 25.0 15.34 5.97 1.87 2.67 0.16 63.27 2.75 0.070 0.76 2.49 26-28 be 27.0 28-30 be 29.0 30-32 be 31.0 15.66 6.12 1.85 2.74 0.16 63.79 2.88 0.064 0.77 2.64 32-34 be 33.0 34-36 be 35.0 15.87 6.40 1.90 2.99 0.17 63.15 3.23 0.072 0.79 2.76 36-38 be 37.0 38-40 be 39.0 40-42 be 41.0 15.69 6.31 1.85 3.03 0.17 61.82 4.16 0.075 0.78 2.71 42-44 be 43.0 44-46 be 45.0 15.53 6.37 1.84 3.16 0.17 61.33 4.46 0.082 0.77 2.94 46-48 be 47.0 213 Tabic A12. (continued) Minor Element Data for Box Core (be) JT96-05. Sample Depth V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br I (m) (Pg/g) (ng/g) (ng/g) (ug/g) (ug/g) (pg/g) (pg/g) (pg/g) (ng/g) (|ig/g) (ng/g) (ng/g) (ng/g) (ng/g) (ng/g) (pg/g) 0-1 be 0.5 134 108 723 13 48 35 100 3 68 286 16 132 920 14 126 538 1-2 be 1.5 145 118 543 16 49 37 103 3 70 277 17 129 911 16 123 502 2-3 be 2.5 147 116 473 16 50 38 101 2 70 278 17 124 902 19 116 469 3-4 be 3.5 4-5 be 4.5 5-6 be 5.5 149 111 475 12 55 43 110 5 71 268 19 135 1057 7 122 438 6-7 be 6.5 7-8 be 7.5 8-9 be 8.5 9-10 be 9.5 10-12 be 11.0 153 116 472 12 54 41 112 3 70 271 18 136 1046 13 119 380 12-14 be 13.0 14-16 be 15.0 155 110 478 13 56 43 117 3 71 270 20 143 1045 11 104 304 16-18 be 17.0 18-20 be 19.0 20-22 be 21.0 173 122 503 11 54 41 117 3 70 271 19 145 970 8 95 234 22-24 be 23.0 24-26 be 25.0 164 115 503 12 56 41 115 2 70 272 21 144 935 13 90 204 26-28 be 27.0 28-30 be 29.0 30-32 be 31.0 166 117 525 13 55 41 111 4 70 276 20 142 904 7 82 174 32-34 be 33.0 34-36 be 35.0 169 114 539 15 55 42 113 3 71 275 22 147 892 10 83 ' 155 36-38 be 37.0 38-40 be 39.0 40-42 be 41.0 163 112 552 13 54 41 110 5 69 295 21 140 848 9 82 134 42-44 be 43.0 44-46 be 45.0 166 110 586 16 52 39 103 4 66 302 21 135 774 12 76 113 46-48 be 47.0 214 Table A12. (continued) Trace Metal Data for Box Core (be) JT96-05. Sample Depth Ag Cd Re Mo U (m) (ng/g) Wg) (ng/g) (pg/g) (Pg/g) 0-1 be 0.5 357 0.14 1.41 0.52 1.01 1-2 be 1.5 375 0.09 1.61 0.41 1.11 2-3 be 2.5 466 0.12 1.54 0.40 0.98 3-4 be 3.5 400 0.18 3.00 0.37 1.32 4-5 be 4.5 431 0.13 2.47 0.36 1.07 5-6 be 5.5 478 0.09 2.64 0.33 1.16 6-7 be 6.5 501 0.10 3.14 0.38 1.07 7-8 be 7.5 553 0.15 2.79 0.39 1.11 8-9 be 8.5 506 0.16 3.29 0.33 1.13 9-10 be 9.5 474 0.22 3.37 0.41 1.07 10-12 be 11.0 573 0.26 3.97 0.45 1.54 12-14 be 13.0 538 0.20 4.90 0.39 1.37 14-16 be 15.0 555 0.22 9.56 0.43 1.71 16-18 be 17.0 584 0.43 14.13 0.49 1.93 18-20 be 19.0 527 0.34 34.81 0.61 2.62 20-22 be 21.0 513 0.70 43.31 0.54 2.59 22-24 be 23.0 519 0.57 17.85 0.63 2.46 24-26 be 25.0 453 0.63 42.22 0.80 2.92 26-28 be 27.0 410 0.56 27.46 0.77 2.88 28-30 be 29.0 459 0.45 30.26 0.58 1.92 30-32 be 31.0 423 0.67 32.41 0.59 2.79 32-34 be 33.0 477 0.65 32.75 0.78 2.89 34-36 be 35.0 441 0.50 30.95 0.78 2.70 36-38 be 37.0 413 0.43 15.95 0.60 2.69 38-40 be 39.0 442 0.44 26.59 0.70 2.61 40-42 be 41.0 445 0.44 15.41 0.70 2.92 42-44 be 43.0 405 0.44 11.70 0.88 3.13 44-46 be 45.0 385 0.42 18.52 0.90 3.18 46-48 be 47.0 402 0.48 25.77 1.15 2.92 215 Table A13. Geochemical Data for Multicore, Box Core and Piston Core JT96-06. Basic Geochemical Data for Multicore (mc). Box Core (be) and Piston Core (pc) JT96-06. Sample Depth Calendar Age cr 5"N 8"C„ C,« S,„ CaCO, %C„„ C„/N Opal Bau„ (m) (kyrs) (wt. %) (%. vs air) (%> vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (pg/g) 0-2 mc 1.0 2.48 6.40 -21.20 0.24 2.09 0.17 1.04 1.97 8.4 2-4 mc 3.0 2.27 0.22 1.94 0.16 0.49 1.88 8.5 4-6 mc 5.0 2.21 6.24 -21.03 0.21 1.80 0.10 0.36 1.75 8.6 6-8 mc 7.0 1.25 0.18 1.59 0.11 0.27 1.56 8.6 8-10 mc 9.0 1.17 0.18 1.59 0.10 0.26 1.56 8.5 10-12 mc 11.0 1.27 6.35 -21.14 0.18 1.63 0.11 0.34 1.59 8.6 12-14 mc 13.0 1.22 0.17 1.52 0.09 0.27 1.48 8.7 14-16 mc 15.0 1.43 6.47 -21.13 0.17 1.52 0.10 0.28 1.49 8.8 16-18 mc 17.0 1.32 6.65 -21.16 0.15 1.40 0.11 0.35 1.36 8.8 Sample Depth Calendar Age Cr S"N 8"Cm N,a C,„ s,„ CaCO, %c„. C„/N Opal Baa. (m) (kyrs) (wt. %) ( %o vs air) (%o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (Pfi/g) 0-0.5 be 0.3 1.25 6.84 -21.17 0.24 2.15 0.19 1.22 2.00 8.3 6.3 164 0.5-1 be 0.8 0.92 0.23 2.06 0.11 1.12 1.93 8.2 1-2 be 1.5 1.01 0.22 1.93 0.11 0.73 1.84 8.4 2-3 be 2.5 0.80 0.23 1.95 0.10 0.57 1.88 8.3 3-4 be 3.5 0.86 0.23 1.98 0.12 0.52 1.92 8.3 4-5 be 4.5 0.84 6.86 -20.99 0.21 1.84 0.09 0.51 1.78 8.4 5.1 150 5-7 be 6.0 0.86 0.21 1.78 0.10 0.41 1.73 8.4 7-9 be 8.0 0.76 0.19 1.66 0.09 0.39 1.61 8.4 9-11 be 10.0 0.78 6.58 -21.16 0.17 1.53 0.08 0.40 1.48 8.5 4.6 125 11-13 be 12.0 0.83 0.18 1.58 0.09 0.40 1.53 8.4 13-15 be 14.0 0.87 6.45 -20.94 0.17 1.53 0.09 0.39 1.48 8.5 4.5 146 17-19 be 18.0 0.85 0.16 1.39 0.09 0.48 1.34 8.6 21-23 be 22.0 0.65 6.61 -23.71 0.07 0.84 0.34 1.70 0.64 9.5 4.3 205 25-27 be 26.0 0.62 6.31 -23.33 0.05 0.91 0.27 3.01 0.55 10.6 3.9 192 29-31 be 30.0 0.61 5.80 -23.50 0.06 0.86 0.25 2.69 0.54 9.8 33-35 be 34.0 0.79 5.94 -23.46 0.06 0.85 0.38 2.25 0.58 9.7 5.2 142 Sample Depth Calendar Age cr 8"N 8"C„ N,„ C s,„ CaCO; %C,„ Opal Ba.i. (m) (kyrs) (wt. %) ( %o vs air) (%o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt.%) (pg/g) l-2pc 0.5 1.49 -21.04 0.19 1.66 0.11 0.31 1.62 8.6 11-12 pc 11.5 1.33 -23.36 0.07 0.79 0.32 1.59 0.60 9.1 21-22 pc 21.5 1.01 -23.17 0.06 0.91 0.34 3.13 0.53 8.6 31-32 pc 31.5 1.12 -23.19 0.06 1.00 0.28 3.68 0.56 8.9 41-42 pc 41.5 1.06 -23.41 0.06 0.98 0.26 3.40 0.57 9.3 51-52 pc 51.5 0.97 -23.70 0.06 0.93 0.25 3.36 0.52 9.5 61-62 pc 61.5 1.27 -23.27 0.07 0.95 0.37 2.54 0.64 9.3 71-72 pc 71.5 1.14 -23.23 0.06 0.87 0.29 2.22 0.60 9.9 81-82 pc 81.5 1.16 -24.33 0.06 0.77 0.31 1.56 0.58 9.9 91-92 pc 91.5 1.16 -24.20 0.06 0.93 0.38 2.59 0.62 9.7 101-102 pc 101.5 1.20 -24.16 0.06 0.91 0.45 2.52 0.61 9.5 11! -112 pc 111.5 1.21 -24.56 0.06 0.72 0.26 1.21 0.57 10.0 121-122 pc 121.5 1.08 -24.58 0.05 0.69 0.29 1.27 0.54 10.2 131-132 pc 131.5 1.13 -24.58 0.05 0.69 0.27 1.34 0.53 10.0 141-142 pc 141.5 1.14 -24.53 0.06 0.69 0.35 1.32 0.53 9.6 151-152 pc 151.5 1.15 -24.68 0.05 0.70 0.28 2.05 0.46 9.7 161-162 pc 161.5 1.09 -24.46 0.05 0.73 0.29 1.68 0.53 9.7 171-172 pc 171.5 1.04 -24.28 0.05 0.72 0.28 1.59 0.53 9.9 181-182 pc 181.5 1.01 -24.63 0.05 0.67 0.28 1.24 0.52 10.3 191-192 pc 191.5 1.07 -24.55 0.04 0.77 0.16 3.32 0.37 8.7 201-202 pc 201.5 0.97 -25.02 0.05 0.65 0.31 1.27 0.50 10.3 211-212 pc 211.5 1.03 -24.59 0.05 0.69 0.29 1.46 0.51 10.1 221-222 pc 221.5 0.99 -24.63 0.05 0.66 0.31 1.30 0.50 9.8 231-232 pc 231.5 1.02 -24.51 0.05 0.71 0.33 1.36 0.55 10.2 241-241 pc 241.5 0.08 -24.52 0.05 0.66 0.34 1.34 0.50 10.1 251-252 pc 251.5 1.01 -24.61 0.05 0.71 0.24 1.58 0.52 10.1 261-262 pc 261.5 0.97 -24.65 0.06 0.69 0.28 1.26 0.54 9.8 271-272 pc 271.5 0.74 -24.44 0.06 0.72 0.33 1.56 0.53 9.3 281-282 pc 281.5 0.76 -24.35 0.05 0.78 0.30 2.25 0.51 9.5 291-292 pc 291.5 0.99 -24.54 0.06 0.71 0.28 1.37 0.54 9.9 301-302 pc 301.5 0.75 -23.69 0.04 0.89 0.11 4.07 0.40 9.1 309-310 pc 309.5 0.78 -23.81 0.04 0.90 0.20 4.53 0.36 8.5 216 Table A13. (continued) Major Element Data for Multicore (mc), Box Core (be) and Piston Core (pc) JT96-06. Sample Depth A1A FejOi K,0 MgO P20, SiO; CaO MnO Ti02 Na,0 (m) (%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 0-2 mc 1.0 13.87 7.01 1.53 2.55 0.20 63.00 3.14 0.069 0.72 1.53 2-4 mc 3.0 4-6 mc 5.0 14.38 7.29 1.69 2.57 0.19 65.53 2.89 0.063 0.74 2.12 6-8 mc 7.0 8-10 mc 9.0 10-12 mc 11.0 14.19 7.43 1.78 2.65 0.17 64.60 2.80 0.064 0.74 2.65 12-14 mc 13.0 14-16 mc 15.0 14.55 7.18 1.83 2.71 0.17 65.62 2.86 0.065 0.75 2.73 16-18 mc 17.0 13.23 6.87 1.78 2.62 0.17 62.38 2.86 0.063 0.73 4.21 Sample Depth Al,Oj Fe20, K20 MgO P.0, SiO, CaO MnO Ti02 Na,0 (m) (%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 0-0.5 be 0.3 13.33 6.69 1.51 2.46 0.21 61.92 3.22 0.065 0.72 7.25 0.5-1 be 0.8 1-2 be 1.5 2-3 be 2.5 3-4 be 3.5 4-5 be 4.5 13.78 6.70 1.60 2.52 0.17 63.04 2.81 0.060 0.73 5.70 5-7 be 6.0 7-9 be 8.0 9-11 be 10.0 13.57 7.16 1.61 2.56 0.19 63.58 2.76 0.063 0.72 5.58 U-13bc 12.0 13-15 be 14.0 13.64 6.65 1.66 2.54 0.17 63.94 2.91 0.066 0.73 6.43 17-19 be 18.0 21-23 be 22.0 16.67 6.65 2.09 3.39 0.19 60.17 3.20 0.091 0.80 5.10 25-27 be 26.0 15.97 6.30 2.01 3.19 0.19 61.03 3.87 0.088 0.78 5.04 29-31 be 30.0 33-35 be 34.0 16.55 6.94 1.95 3.47 0.19 59.01 3.73 0.101 0.81 5.07 Sample Depth A1,0, Fe.O, KjO MgO PjO, SiOi CaO MnO TiO, Na;0 (m) (%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 1-2 pc 0.5 13.07 7.22 1.70 2.61 0.17 61.64 2.69 0.061 0.72 4.82 11-12 pc 11.5 15.96 6.86 2.46 3.38 0.19 57.45 3.01 0.100 0.81 4.56 21-22 pc 21.5 15.37 6.59 2.15 3.23 0.19 58.99 3.95 0.093 0.79 3.05 31-32 pc 31.5 15.66 6.67 2.27 3.34 0.20 57.86 4.07 0.096 0.80 2.95 41-42 pc 41.5 15.53 6.49 2.19 3.22 0.19 59.32 3.92 0.103 0.79 2.90 51-52 pc 51.5 15.19 6.28 2.16 3.10 0.19 59.59 4.00 0.099 0.77 3.56 61-62 pc 61.5 15.84 6.78 2.03 3.37 0.18 58.15 3.62 0.101 0.79 3.31 71-72 pc 71.5 15.46 6.82 1.86 3.38 0.18 58.28 3.83 0.087 0.79 4.16 81-82 pc 81.5 16.25 7.19 1.87 3.49 0.18 57.60 2.94 0.094 0.81 4.55 91-92 pc 91.5 15.91 6.95 1.98 3.32 0.18 56.92 3.27 0.099 0.82 4.02 101-102 pc 101.5 16.01 6.92 1.94 3.37 0.18 56.78 3.19 0.090 0.81 4.59 111-112 pc 111.5 16.25 7.49 1.93 3.64 0.20 56.17 2.93 0.101 0.85 7.89 121-122 pc 121.5 15.91 7.49 1.96 3.48 0.20 56.09 3.02 0.108 0.86 3.57 131-132 pc 131.5 17.62 7.81 2.05 3.72 0.22 61.42 3.09 0.102 0.88 3.56 141-142 pc 141.5 16.07 7.45 1.88 3.42 0.20 57.56 3.15 0.107 0.85 5.08 151-152 pc 151.5 16.22 7.30 1.98 3.46 0.20 57.81 3.31 0.101 0.87 2.84 161-162 pc 161.5 16.10 7.43 1.89 3.49 0.20 57.65 3.33 0.104 0.85 2.91 171-172 pc 171.5 16.25 7.31 1.84 3.46 0.20 57.96 3.27 0.103 0.85 2.90 181-182 pc 181.5 16.14 7.39 1.84 3.43 0.20 57.79 3.09 0.097 0.84 2.93 191-192 pc 191.5 16.01 6.77 2.56 3.34 0.22 56.65 3.47 0.104 0.88 3.80 201-202 pc 201.5 16.02 7.50 1.75 3.49 0.20 57.50 3.18 0.105 0.87 3.68 211-212 pc 211.5 16.05 7.31 1.83 3.45 0.20 56.92 3.10 0.110 0.85 6.42 221-222 pc 221.5 16.05 7.45 1.84 3.45 0.20 57.15 3.06 0.102 0.85 3.60 231-232 pc 231.5 16.18 7.48 1.80 3.49 0.20 56.56 2.98 0.102 0.86 5.90 241-241 pc 241.5 251-252 pc 251.5 16.54 7.55 2.13 3.59 0.21 56.50 2.99 0.104 0.87 4.02 261-262 pc 261.5 16.74 7.80 1.93 3.59 0.20 58.35 3.00 0.102 0.87 3.77 271-272 pc 271.5 281-282 pc 281.5 291-292 pc 291.5 16.18 7.60 2.01 3.52 0.19 56.25 3.01 0.104 0.85 3.17 301-302 pc 301.5 14.25 6.46 1.88 2.98 0.19 59.59 4.88 0.089 0.90 3.01 309-310 pc 309.5 14.70 6.57 2.16 2.92 0.21 58.82 4.76 0.090 0.93 4.58 217 Table A13. (continued) Minor Element Data for Multicore (mc), Box Core (be) and Piston Core (pc) JT96-06. Sample Depth V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br I (m) (ug/g) (ug/g) (ng/g) (Pg/g) (ug/g) (pg/g) (Pg/g) (Pg/g) (pg/g) (ng/g) (Pg/g) (pg/g) (ng/g) (Pg/g) (Cg/g) (Pg/g) 0-2 mc 1.0 140 127 496 11 48 31 95 3 73 283 16 156 534 17 170 635 2-4 mc 3.0 4-6 mc 5.0 152 141 493 14 51 31 97 7 68 279 18 161 542 8 104 430 6-8 mc 7.0 8-10 mc 9.0 10-12 mc 11.0 148 131 499 10 52 32 97 4 70 279 19 166 552 10 108 326 12-14 mc 13.0 14-16 mc 15.0 151 129 506 9 49 30 92 5 68 287 20 160 547 11 95 312 16-18 mc 17.0 155 142 529 12 47 28 91 6 66 292 18 170 553 4 86 265 Sample Depth V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br I (m) (Pg/g) (pg/g) (Pg/g) (pg/g) (ng/g) (ng/g) (Pg/g) (ng/g) (Pg/g) (ng/g) (ng/g) (Pg/g) (Pg/g) (pg/g) (Pg/g) (ng/g) 0-0.5 be 0.3 146 130 495 20 44 32 99 2 69 289 19 157 498 22 154 484 0.5-1 be 0.8 1-2 be 1.5 2-3 be 2.5 3-4 be 3.5 4-5 be 4.5 155 145 473 20 45 30 96 2 68 278 19 168 495 19 115 313 5-7 be 6.0 7-9 be 8.0 9-11 be 10.0 146 147 472 20 43 29 92 2 66 285 18 176 465 20 100 256 11-13 be 12.0 13-15 be 14.0 155 147 496 20 44 29 92 2 65 287 18 169 487 18 91 215 17-19 be 18.0 21-23 be 22.0 187 127 718 21 54 37 102 2 82 287 23 131 622 23 29 27 25-27 be 26.0 173 123 710 20 51 34 96 5 77 305 24 135 591 23 23 20 29-31 be 30.0 33-35 be 34.0 195 124 791 23 54 40 101 4 71 304 23 129 555 25 22 18 Sample Depth V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br I (m) (pg/g) (ug/g) (pg/g) (pg/g) (Pg/g) (pg/g) (Pg/g) (pg/g) (ng/g) (ng/g) (ng/g) (pg/g) (pg/g) (ng/g) (pg/g) (ng/g) 1-2 pc 0.5 146 132 485 13 47 30 94 4 69 276 19 154 519 13 107 388 ll-12pc 11.5 182 123 743 16 61 40 109 8 89 284 24 137 683 8 16 2 21-22 pc 21.5 170 118 748 20 51 33 98 5 76 307 23 143 616 15 13 -11 31-32 pc 31.5 181 122 807 15 55 36 101 9 82 304 24 137 679 6 13 -10 41-42 pc 41.5 169 116 774 13 56 35 98 6 80 309 24 143 649 14 12 13 51-52 pc 51.5 166 115 790 15 52 34 96 7 77 311 23 136 647 10 11 0 61-62 pc 61.5 181 115 762 18 56 39 102 6 72 286 24 129 586 12 16 -1 71-72 pc 71.5 184 131 805 17 60 37 98 7 66 298 22 132 586 11 14 -7 81-82 pc 81.5 195 123 781 16 59 37 99 7 64 257 23 127 537 10 9 -1 91-92 pc 91.5 195 112 730 15 51 37 96 7 69 252 24 135 537 13 10 -9 101-102 pc 101.5 190 113 714 16 53 38 97 6 70 247 23 133 546 14 11 -12 111-112 pc 111.5 210 115 827 18 56 44 102 4 63 253 23 126 517 14 8 -23 121-122 pc 121.5 202 108 801 21 50 43 100 9 63 261 24 129 505 7 6 -14 131-132 pc 131.5 201 108 783 18 49 43 102 5 67 260 24 131 513 15 8 -17 141-142 pc 141.5 201 114 818 17 53 41 98 6 60 271 24 129 550 10 3 -5 151-152 pc 151.5 197 102 810 21 46 41 101 7 75 257 25 135 536 15 1 -6 161-162 pc 161.5 195 102 793 17 49 41 101 7 75 260 26 136 534 13 2 -1 171-172 pc 171.5 197 109 800 20 51 40 97 8 62 272 24 131 506 11 4 -8 181-182 pc 181.5 203 114 826 17 50 42 98 7 60 273 24 131 524 12 2 -6 191-192 pc 191.5 197 111 809 16 52 40 97 9 58 271 24 127 506 8 1 -11 201-202 pc 201.5 211 118 865 17 52 45 98 8 58 262 23 127 532 9 1 -7 211-212 pc 211.5 199 106 820 16 51 42 103 10 64 264 25 127 531 7 2 -6 221-222 pc 221.5 200 113 830 17 52 40 99 9 59 274 24 129 521 10 3 • .2 231-232 pc 231.5 209 118 859 15 52 45 102 8 62 263 24 126 545 13 2 -16 241-241 pc 241.5 251-252 pc 251.5 205 109 840 22 51 42 104 7 71 251 25 131 527 14 2 -11 261-262 pc 261.5 203 112 836 20 51 42 99 11 59 257 24 125 497 7 3 -15 271-272 pc 271.5 281-282 pc 281.5 291-292 pc 291.5 207 113 846 18 53 44 101 9 62 262 25 128 504 9 0 -8 301-302 pc 301.5 184 105 738 15 44 29 85 6 68 307 26 186 576 10 6 -5 309-310 pc 309.5 181 91 708 16 41 34 92 4 85 284 27 178 598 15 5 2 218 Table A13. (continued) Trace Metal Data for Multicore (mc), Box Core (be) and Piston Core (pc) JT96-06. Sample Deptii Ag Cd Re Mo U (m) (ng/g) (Pg/g) (ng/g) (Pg/g) (Pg/g) 0-2 mc 1.0 111 0.15 2.44 0.62 1.07 2-4 mc 3.0 114 0.25 2.93 0.65 1.34 4-6 mc 5.0 111 0.22 4.26 0.56 1.21 6-8 mc 7.0 111 0.19 5.00 0.60 1.37 8-10 mc 9.0 105 0.18 4.48 0.53 1.33 10-12 mc 11.0 111 0.19 4.87 0.51 1.34 12-14 mc 13.0 97 0.21 4.98 0.56 1.52 14-16 mc 15.0 106 0.23 5.21 0.55 1.26 16-18 mc 17.0 115 0.26 4.73 0.55 1.28 Sample Depth Ag Cd Re Mo U (m) (ng/g) (Pg/g) (ng/g) (Pg/g) (Pg/g) 0-0.5 be 0.3 97 0.16 3.63 0.57 1.49 0.5-1 be 0.8 94 0.16 2.59 0.53 0.96 1-2 be 1.5 84 0.21 3.12 0.54 1.26 2-3 be 2.5 100 0.29 3.64 0.58 1.38 3-4 be 3.5 103 0.31 3.86 0.66 1.40 4-5 be 4.5 117 0.24 3.65 0.60 1.53 5-7 be 6.0 98 0.24 3.91 0.52 1.45 7-9 be 8.0 94 0.17 4.20 0.47 1.38 9-11 be 10.0 106 0.16 3.69 0.48 1.56 11-13 be 12.0 85 0.16 3.90 0.43 0.99 13-15 be 14.0 106 0.20 5.12 0.62 1.64 17-19 be 18.0 91 0.30 5.63 0.51 1.15 21-23 be 22.0 106 0.42 9.10 0.82 2.17 25-27 be 26.0 96 0.22 5.44 1.05 2.44 29-31 be 30.0 106 0.26 3.50 1.35 2.17 33-35 be 34.0 119 0.24 3.23 1.18 2.03 Table A14. Geochemical Data for Multicore and Piston Core JT96-09. Genera! Geochemical Data for Multicore (mc) and Piston Core (pc) JT96-09. Sample Depth Depth » Calendar Age cr 8"N 8"C„. N,„ s,„ CaC02 %Conl C./N Opal Ba,,. (cm) (cm, corr.) (kyrs) (wt. %) (%o vs air) (%o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (Pg/g) 0-1 mc 0.5 0.5 0.11 4.04 6.74 -21.28 0.35 3.20 0.35 1.05 3.07 8.7 7.0 256 1-2 mc 1.5 1.5 0.32 3.47 0.36 3.16 0.35 0.92 3.05 8.4 2-3 mc 2.5 2.5 0.53 3.08 0.37 3.17 0.29 0.85 3.07 8.3 3-4 mc 3.5 3.5 0.74 2.72 0.36 3.07 0.26 0.64 2.99 8.4 4-5 mc 4.5 4.5 0.95 2.46 0.36 3.09 0.25 0.53 3.03 8.4 5-6 mc 5.5 5.5 1.16 2.31 7.03 -21.19 0.35 3.03 0.23 0.58 2.96 8.5 6.7 226 6-8 mc 7.0 7.0 1.48 2.23 0.35 3.01 0.22 0.36 2.97 8.5 8-10 mc 9.0 9.0 1.90 2.16 0.35 3.01 0.22 0.30 2.97 8.5 10-12 mc 11.0 11.0 2.32 2.31 6.66 -21.19 0.31 2.80 0.25 0.18 2.78 8.9 6.8 202 12-14 mc 13.0 13.0 2.74 2.06 0.33 2.84 0.22 0.29 2.80 8.6 14-16 mc 15.0 15.0 3.17 2.17 6.81 -21.17 0.31 2.75 0.21 0.32 2.71 8.7 5.8 230 16-18 mc 17.0 17.0 3.59 2.51 0.32 2.79 0.22 0.33 2.75 8.6 18-20 mc 19.0 19.0 4.01 1.84 0.23 2.11 0.17 0.35 2.07 8.9 20-22 mc 21.0 21.0 4.43 1.92 6.50 -21.39 0.20 1.96 0.18 0.41 1.91 9.4 4.3 219 22-24 mc 23.0 23.0 4.86 1.63 0.19 1.79 0.14 0.72 1.70 9.1 24-26 mc 25.0 25.0 5.28 1.50 7.71 -21.67 0.18 1.79 0.15 1.16 1.65 9.3 4.4 174 26-28 mc 27.0 27.0 5.70 1.61 0.18 1.83 0.15 1.87 1.61 9.0 28-30 mc 29.0 29.0 6.12 1.68 0.18 1.90 0.17 2.10 1.64 9.0 30-32 mc 31.0 31.0 6.54 1.63 7.35 -21.64 0.18 1.91 0.19 2.33 1.63 9.2 4.5 181 32-34 mc 33.0 33.0 6.97 1.80 0.19 1.99 0.20 2.40 1.70 9.1 34-36 mc 35.0 35.0 7.39 1.68 7.93 -21.63 0.18 2.01 0.21 2.37 1.72 9.4 4.9 166 36-38 mc 37.0 37.0 7.81 1.95 0.20 2.08 0.23 2.39 1.79 9.1 38-40 mc 39.0 39.0 8.23 1.77 7.45 -21.85 0.19 2.07 0.27 2.62 1.76 9.2 4.8 181 Sample Depth Depth » Calendar Age cr 8I!N 8"C„ N„ C» s,„ CaCOi "/"Co™ C„/N Opal Ba^ (cm) (cm, corr.) (kyrs) (wt. %) (%. vs air) (%o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (Pg/g) 6-7 pc 6.5 18.5 3.91 2.44 7.33 -21.13 0.30 2.67 0.22 0.83 2.57 8.5 6.6 249 U-12pc 11.5 23.5 4.96 1.93 6.95 -21.55 0.21 2.02 0.20 0.71 1.94 9.2 4.7 168 16-17 pc 16.5 28.5 6.02 2.01 8.17 -21.56 0.20 2.08 0.22 2.15 1.82 9.1 4.5 21-22 pc 21.5 33.5 7.07 2.13 7.77 -21.92 0.21 2.29 0.32 2.34 2.01 9.4 5.7 208 26-27 pc 26.5 38.5 8.13 2.11 8.33 -21.75 0.22 2.26 0.39 1.83 2.04 9.2 4.7 31-32 pc 31.5 43.5 9.18 2.01 8.38 -22.04 0.20 2.23 0.62 2.74 1.90 9.3 5.9 240 36-37 pc 36.5 48.5 10.10 2.01 8.61 -22.02 0.20 2.46 0.64 4.91 1.87 9.1 5.9 41-42 pc 41.5 53.5 10.47 2.01 8.34 -21.96 0.20 2.55 0.42 5.45 1.90 9.6 6.8 231 46-47 pc 46.5 58.5 10.84 2.05 8.89 -21.99 0.19 2.41 0.99 4.43 1.87 9.7 7.1 51-52 pc 51.5 63.5 11.21 1.73 7.10 -22.19 0.12 1.47 0.45 2.50 1.17 9.4 4.8 202 56-57 pc 56.5 68.5 11.57 1.44 6.86 -22.67 0.07 0.93 0.26 1.83 0.71 10.0 4.8 178 61-62 pc 61.5 73.5 11.94 1.03 6.23 -23.11 0.06 0.80 0.21 2.26 0.53 8.7 3.7 169 66-67 pc 66.5 78.5 12.29 1.06 5.89 -23.08 0.05 0.77 0.16 1.87 0.55 11.2 4.1 71-72 pc 71.5 83.5 12.54 0.68 6.21 -23.53 0.04 0.73 0.19 2.77 0.39 10.8 4.3 179 76-77 pc 76.5 88.5 12.74 1.64 7.41 -22.54 0.09 1.20 0.40 2.92 0.85 9.8 4.7 276 81-82 pc 81.5 93.5 12.78 1.48 7.75 -22.75 0.10 1.38 0.34 3.96 0.90 9.3 5.2 285 86-87 pc 86.5 98.5 12.81 1.39 7.20 -22.42 0.09 1.21 0.34 2.67 0.89 9.6 6.1 281 91-92 pc 91.5 103.5 12.87 1.36 7.52 -22.91 0.09 1.30 0.37 3.97 0.82 9.1 4.8 265 96-97 pc 96.5 108.5 13.04 1.35 7.38 -22.44 0.09 1.28 0.40 3.72 0.84 9.8 5.7 265 101-102 pc 101.5 113.5 13.18 1.31 7.36 -22.75 0.10 1.34 0.36 4.04 0.85 8.9 294 106-107 pc 106.5 118.5 13.22 1.31 7.28 -22.69 0.08 1.24 0.31 3.39 0.83 9.9 5.3 237 111-112 pc 111.5 123.5 13.27 1.36 7.10 -23.16 0.08 1.18 0.29 3.47 0.76 9.1 4.6 280 H6-117pc 116.5 128.5 13.31 1.26 6.77 -22.87 0.08 1.15 0.32 3.39 0.74 9.7 5.8 231 121-122 pc 121.5 133.5 13.36 1.38 7.37 -23.01 0.09 1.16 0.38 2.82 0.82 9.0 5.3 250 126-127 pc 126.5 138.5 13.40 1.26 6.75 -23.01 0.08 1.10 0.40 2.99 0.74 9.5 4.7 252 131-132 pc 131.5 143.5 13.44 1.26 7.23 -23.26 0.09 1.12 0.43 2.96 0.77 9.0 5.6 168 136-137 pc 136.5 148.5 13.47 0.64 4.92 -23.75 0.02 0.45 0.30 1.89 0.22 9.1 155 141-142 pc 141.5 153.5 13.47 0.77 7.02 -23.39 0.04 0.76 0.32 2.62 0.45 10.2 180 146-147 pc 146.5 158.5 13.47 0.46 2.93 -23.32 0.01 0.17 0.07 0.87 0.07 6.8 128 151-152 pc 151.5 163.5 13.47 1.01 4.79 -24.24 0.05 0.77 0.33 1.72 0.56 11.0 119 156-157 pc 156.5 152.5 13.49 1.26 4.84 -24.22 0.05 0.74 0.31 1.87 0.52 10.1 4.0 109 161-162 pc 161.5 157.5 13.52 0.93 4.22 -25.20 0.04 0.73 0.16 3.42 0.32 8.3 4.2 204 166-167 pc 166.5 162.5 13.55 0.95 5.00 -24.26 0.05 0.78 0.27 2.36 0.50 9.7 109 171-172 pc 171.5 167.5 13.58 1.06 4.81 -24.64 0.05 0.71 0.29 1.66 0.51 9.7 3.6 76 176-177 pc 176.5 172.5 13.61 0.94 3.98 -24.62 0.05 0.67 0.33 1.54 0.49 10.2 73 181-182 pc 181.5 177.5 13.64 0.95 4.33 -25.05 0.04 0.75 0.14 4.01 0.27 6.9 4.6 226 186-187 pc 186.5 182.5 13.67 0.96 4.57 -24.51 0.05 0.70 0.32 1.77 0.48 10.5 95 191-192 pc 191.5 187.5 13.70 1.00 4.44 -24.86 0.05 0.65 0.38 1.52 0.46 9.7 3.4 69 196-197 pc 196.5 192.5 13.73 0.91 4.21 -24.79 0.05 0.67 0.36 1.27 0.51 11.0 63 201-202 pc 201.5 197.5 13.76 0.91 4.11 -24.68 0.05 0.67 0.34 1.36 0.51 9.8 3.5 65 206-207 pc 206.5 202.5 13.79 0.85 3.72 -24.52 0.04 0.64 0.35 2.02 0.40 10.2 115 211-212 pc 211.5 207.5 13.82 0.99 4.38 -24.69 0.05 0.68 0.37 1.50 0.50 9.5 3.6 37 216-217 pc 216.5 212.5 13.85 0.89 3.50 -24.71 0.05 0.66 0.40 1.21 0.51 10.9 41 221-222 pc 221.5 217.5 13.88 1.17 4.06 -24.81 0.06 0.71 0.36 1.31 0.55 9.9 3.7 33 226-227 pc 226.5 222.5 13.91 0.92 3.81 -24.60 0.05 0.66 0.35 1.27 0.50 10.7 43 231-232 pc 231.5 227.5 13.94 0.81 4.03 -24.68 0.05 0.63 0.37 1.38 0.46 9.5 3.4 42 236-237 pc 236.5 232.5 13.97 0.86 3.85 -24.47 0.05 0.73 0.35 1.32 0.57 11.1 69 241-242 pc 241.5 237.5 14.00 0.91 4.29 -24.62 0.05 0.68 0.35 1.20 0.53 10.1 3.4 34 246-247 pc 246.5 242.5 14.03 1.01 4.31 -24.42 0.05 0.74 0.23 2.07 0.49 10.2 110 251-252 pc 251.5 247.5 14,06 0.83 3.96 -24.59 0.05 0.59 0.38 1.31 0.44 9.7 3.3 11 256-257 pc 256.5 252.5 14.09 0.97 3.85 -24.49 0.05 0.69 0.28 1.42 0.52 10.4 68 261-262 pc 261.5 257.5 14.12 0.97 5.08 -24.49 0.05 0.75 0.24 1.68 0.55 10.3 3.6 68 (continued on next page) 220 Table A14. (continued) General Geochemical Data for Multicore (mc) and Piston Core (pc) JT96-09 (continued). Sample Depth Depth * Calendar Age cr 5"N 8"C„ N» s,„, CaCO; %c„ C./N Opal (cm) (cm, com) (kyrs) (wt.%) (%ovsair) (%ovsPDB) (wt.%) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (ng/fi) 266-267 pc 266.5 262.5 14.15 1.01 4.72 -24.27 0.05 0.76 0.24 1.68 0.56 10.5 82 271-272 pc 271.5 267.5 14.18 0.82 4.91 -24.74 0.05 0.74 0.18 2.77 0.41 8.9 4.1 131 276-277 pc 276.5 272.5 14.21 0.94 4.50 -24.41 0.05 0.78 0.35 2.45 0.49 10.0 129 281-282 pc 281.5 277.5 14.25 1.02 4.85 -24.71 0.06 0.81 0.24 2.27 0.53 9.5 3.8 88 286-287 pc 286.5 282.5 14.28 0.95 5.03 -24.24 0.05 0.92 0.23 3.07 0.55 10.3 119 291-292 pc 291.5 287.5 14.33 0.92 5.44 -24.37 0.06 0.91 0.29 3.33 0.51 9.2 4.4 122 296-297 pc 296.5 292.5 14.43 0.91 5.14 -24.15 0.06 0.96 0.22 3.22 0.57 10.2 101 301-302 pc 301.5 297.5 14.54 0.98 4.93 -24.60 0.05 0.88 0.17 3.12 0.50 9.2 3.5 105 306-307 pc 306.5 302.5 14.60 0.89 4.91 -24.43 0.05 0.84 0.21 2.29 0.57 10.3 88 311-312 pc 311.5 307.5 14.75 0.87 5.58 -24.28 0.05 0.93 0.28 4.14 0.43 8.8 4.1 115 316-317 pc 316.5 312.5 14.85 0.81 5.38 -24.07 0.04 0.83 0.13 3.85 0.36 8.8 199 321-322 pc 321.5 317.5 14.96 0.90 5.31 -24.30 0.04 0.80 0.11 3.45 0.39 9.0 4.9 167 326-327 pc 326.5 322.5 15.06 0.92 5.16 -24.23 0.04 0.78 0.16 3.49 0.36 9.1 211 331-332 pc 331.5 327.5 15.17 0.74 4.45 -24.42 0.04 0.75 0.11 3.86 0.29 7.7 4.8 156 336-337 pc 336.5 332.5 15.27 0.81 5.08 -24.36 0.04 0.75 0.12 3.72 0.31 8.7 227 341-342 pc 341.5 337.5 15.38 0.86 5.59 -24,09 0.04 0.82 0.17 3.66 0.38 8.6 4.4 135 346-347 pc 346.5 342.5 15.48 0.79 5.07 -24.09 0.04 0.89 0.09 4.31 0.37 9.4 351-352 pc 351.5 347.5 15.59 0.86 4.98 -24.41 0.04 0.74 0.12 3.75 0.29 7.7 4.6 173 356-357 pc 356.5 352.5 15.69 0.74 4.58 -24.64 0.03 0.65 0.12 3.37 0.24 8.7 197 361-362 pc 361.5 357.5 15.80 0.92 5.81 -24.10 0.05 0.80 0.10 3.28 0.40 8.7 4.7 201 366-367 pc 366.5 362.5 15.90 0.73 5.65 -24.22 0.04 0.74 0.10 3.52 0.32 9.0 204 371-372 pc 371.5 367.5 16.01 0.78 4.66 -24.52 0.03 0.67 0.14 3.39 0.27 7.7 4.7 151 * Corrected piston core depth (12 cm added to account for loss of surface sediments during piston coring and a 16 cm turbidite was removed from the depth tally). 221 Table A14. (continued) Major Element Data for Multicore (mc) and Piston Core (pc) JT96-09. Sample Depth* AlA FeA K,0 MgO PA SiOj CaO MnO TiO; Na;0 (cm, corr.) (%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 0-1 mc 0.5 13.61 6.88 1.77 3.08 0.25 52.64 2.66 0.054 0.82 1.56 1-2 mc 1.5 2-3 mc 2.5 3-4 mc 3.5 4-5 mc 4.5 5-6 mc 5.5 14.74 6.84 1.53 3.02 0.22 56.54 2.32 0.058 0.82 6.68 6-8 mc 7.0 8-10 mc 9.0 10-12 mc 11.0 15.38 6.85 1.84 3.21 0.20 58.10 2.22 0.058 0.84 2.21 12-14 mc 13.0 14-16 mc 15.0 14.92 6.60 1.63 3.06 0.20 57.55 2.29 0.062 0.84 7.53 16-18 mc 17.0 18-20mc 19.0 20-22 mc 21.0 14.01 6.32 1.97 2.92 0.20 54.56 2.58 0.062 0.85 2.07 22-24 mc 23.0 24-26 mc 25.0 15.64 6.44 1.86 3.09 0.19 58.87 2.95 0.066 0.85 5.57 26-28 mc 27.0 28-30 mc 29.0 30-32 mc 31.0 15.81 6.40 1.96 3.26 0.20 57.76 3.42 0.069 0.86 2.82 32-34 mc 33.0 34-36 mc 35.0 15.72 6.54 1.81 3.16 0.19 57.63 3.56 0.071 0.87 6.14 36-38 mc 37.0 38-40 mc 39.0 16.16 6.83 2.30 3.10 0.22 58.03 3.72 0.072 0.91 3.84 Sample Depth * AlA FeA K,0 MgO PA SA CaO MnO Ti02 Na20 (cm, corr.) (%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 6-7 pc 18.5 14.94 6.68 1.92 2.86 0.20 58.29 2.31 0.063 0.87 7.32 11-12 pc 23.5 15.10 6.23 1.82 3.13 0.18 56.78 2.42 0.058 0.81 2.48 16-17 pc 28.5 21-22 pc 33.5 15.67 6.67 1.80 3.18 0.20 55.62 3.30 0.073 0.87 7.79 26-27 pc 38.5 15.90 6.79 2.09 3.35 0.20 56.26 3.12 0.070 0.86 5.40 31-32 pc 43.5 14.83 6.63 1.94 3.11 0.21 53.04 3.49 0.052 0.83 2.70 36-37 pc 48.5 15.18 6.67 2.01 3.24 0.19 54.57 4.96 0.069 0.80 5.38 41-42 pc 53.5 15.20 6.12 1.84 3.43 0.18 53.33 5.05 0.065 0.79 2.37 46-47 pc 58.5 14.87 7.02 1.91 3.32 0.18 54.58 4.85 0.078 0.78 5.51 51-52 pc 63.5 15.41 6.76 2.18 3.13 0.20 56.98 3.82 0.072 0.83 3.66 56-57 pc 68.5 15.64 6.48 2.11 3.38 0.19 60.43 3.61 0.088 0.78 6.78 61-62 pc 73.5 15.90 6.49 2.13 2.83 0.18 64.47 3.81 0.082 0.77 4.50 66-67 pc 78.5 14.51 5.50 1.99 2.88 0.17 63.98 3.78 0.079 0.71 5.42 71-72 pc 83.5 13.92 5.20 1.65 2.38 0.17 64.89 4.03 0.091 0.74 12.30 76-77 pc 88.5 16.39 6.76 2.41 3.55 0.19 57.88 3.77 0.093 0.78 6.88 81-82 pc 93.5 15.93 6.68 2.37 3.28 0.21 54.55 4.23 0.092 0.78 3.55 86-87 pc 98.5 16.27 6.79 2.45 3.60 0.19 57.05 3.64 0.093 0.79 5.25 91-92 pc 103.5 16.97 6.75 2.35 3.73 0.19 57.11 4.10 0.094 0.76 2.99 96-97 pc 108.5 16.23 7.07 2.28 3.62 0.20 56.50 4.28 0.099 0.78 6.21 101-102 pc 113.5 16.37 6.70 2.33 3.36 0.20 55.68 4.16 0.092 0.79 3.57 106-107 pc 118.5 16.30 6.84 2.34 3.59 0.19 56.94 4.08 0.096 0.79 5.63 111-112 pc 123.5 14.67 6.90 2.18 3.33 0.19 51.27 3.89 0.092 0.79 1.45 U6-117pc 128.5 16.48 6.96 2.36 3.61 0.20 57.16 4.02 0.100 0.79 5.29 121-122 pc 133.5 16.03 6.93 2.34 3.30 0.21 56.20 3.64 0.092 0.80 4.00 126-127 pc 138.5 16.42 7.10 2.31 3.61 0.19 57.76 3.69 0.101 0.79 4.43 131-132 pc 143.5 16.28 6.79 2.17 3.63 0.18 57.84 3.60 0.091 0.76 2.73 136-137 pc 148.5 14.19 5.28 1.47 2.28 0.16 66.23 4.08 0.091 0.71 10.07 141-142 pc 153.5 15.49 6.24 1.61 2.95 0.19 58.87 3.10 0.091 0.73 12.39 146-147 pc 158.5 13.26 4.12 1.19 1.86 0.13 71.17 3.95 0.071 0.61 6.77 151-152 pc 163.5 14.57 5.97 1.66 2.40 0.16 63.94 7.12 0.092 0.71 12.42 156-157 pc 152.5 17.03 7.13 2.10 3.61 0.21 59.32 3.29 0.098 0.85 5.07 161-162 pc 157.5 15.94 6.47 2.79 2.98 0.22 59.93 3.86 0.081 0.92 4.40 166-167 pc 162.5 16.57 7.08 2.04 3.50 0.21 58.78 3.64 0.100 0.86 5.24 171-172 pc 167.5 16.68 7.21 2.05 3.30 0.22 57.98 3.33 0.092 0.87 4.30 176-177 pc 172.5 16.61 7.20 1.89 3.58 0.21 57.78 3.29 0.103 0.84 5.15 181-182 pc 177.5 16.55 6.45 2.97 3.06 0.23 59.14 3.79 0.092 0.94 4.59 186-187 pc 182.5 16.74 7.28 2.03 3.58 0.21 58.81 3.30 0.100 0.85 5.09 191-192 pc 187.5 16.70 7.26 1.94 3.25 0.21 58.35 3.23 0.102 0.85 4.64 196-197 pc 192.5 16.68 7.39 1.93 3.63 0.22 58.41 3.13 0.106 0.84 5.33 201-202 pc 197.5 16.70 7.25 1.93 3.24 0.21 58.31 3.23 0.102 0.84 4.67 206-207 pc 202.5 16.53 6.90 2.13 3.34 0.20 59.95 3.44 0.100 0.84 5.86 211-212 pc 207.5 16.93 7.25 1.81 3.71 0.20 57.82 3.11 0.104 0.84 3.35 216-217 pc 212.5 16.86 7.35 1.91 3.59 0.21 59.05 3.10 0.101 0.85 5.65 221-222 pc 217.5 16.93 7.25 1.94 3.29 0.20 57.69 3.06 0.092 0.86 4.39 226-227 pc 222.5 16.92 7.45 1.88 3.65 0.22 58.65 3.13 0.107 0.84 6.31 231-232 pc 227.5 16.44 7.05 1.87 3.14 0.20 58.73 3.21 0.091 0.84 3.88 236-237 pc 232.5 16.82 7.42 1.98 3.67 0.20 58.00 3.10 0.105 0.84 5.71 241-242 pc 237.5 15.92 7.68 2.08 3.28 0.23 55.99 3.14 0.102 0.87 3.26 246-247 pc 242.5 16.85 7.35 2.45 3.56 0.21 57.70 3.20 0.104 0.88 5.18 251-252 pc 247.5 16.83 6.82 1.75 3.42 0.19 60.18 3.22 0.093 0.80 3.40 256-257 pc 252.5 16.89 7.55 2.02 3.71 0.20 58.20 3.05 0.105 0.84 5.23 261-262 pc 257.5 16.55 7.06 2.11 3.28 0.21 58.61 3.45 0.092 0.87 4.12 (continued on next page) 222 Table A14. (continued) Major Element Data for Multicore (mc) and Piston Core (pc) JT96-09 (continued). Sample Depth * A1A FeA KA MgO PA SiOj CaO MnO TA Na,0 (cm, com) (%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 266-267 pc 262.5 16.93 7.41 1.98 3.71 0.20 58.50 3.31 0.102 0.87 4.80 271-272 pc 267.5 16.34 6.90 2.58 3.12 0.22 60.25 3.88 0.091 0.92 4.54 276-277 pc 272.5 16.39 7.27 2.28 3.56 0.21 58.03 3.55 0.100 0.87 5.19 281-282 pc 277.5 16.21 7.34 2.20 3.28 0.25 57.03 3.60 0.102 0.90 3.63 286-287 pc 282.5 16.20 7.03 2.05 3.45 0.21 58.47 4.18 0.101 0.85 5.30 291-292 pc 287.5 16.27 7.00 2.44 3.13 0.22 58.01 3.95 0.081 0.88 4.54 296-297 pc 292.5 16.38 7.17 1.99 3.52 0.21 58.03 4.10 0.101 0.86 5.15 301-302 pc 297.5 16.37 7.11 2.23 3.21 0.21 57.58 3.80 0.092 0.88 4.56 306-307 pc 302.5 16.58 7.31 2.01 3.63 0.21 57.92 3.60 0.105 0.85 5.79 311-312 pc 307.5 15.12 6.85 2.23 2.88 0.22 56.84 4.58 0.081 0.98 4.36 316-317 pc 312.5 15.54 7.02 2.55 3.13 0.23 59.29 4.31 0.094 1.01 7.53 321-322 pc 317.5 15.85 6.99 2.65 2.95 0.24 57.85 4.10 0.081 1.03 3.94 326-327 pc 322.5 16.04 7.43 2.75 3.18 0.25 57.85 4.00 0.096 1.04 5.01 331-332 pc 327.5 15.05 6.98 2.40 2.80 0.25 58.30 439 0.091 1.09 4.38 336-337 pc 332.5 15.53 6.93 2.76 3.04 0.23 58.82 3.97 0.093 0.99 4.82 341-342 pc 337.5 16.25 7.12 2.55 3.24 0.23 59.55 4.13 0.082 0.99 3.29 346-347 pc 342.5 15.41 6.79 2.34 3.07 0.23 59.89 4.68 0.096 0.96 4.87 351-352 pc 347.5 15.32 6.96 2.50 2.94 0.25 57.15 4.08 0.091 1.07 3.46 356-357 pc 352.5 14.99 7.00 2.56 2.99 0.26 61.09 4.26 0.105 1.10 4.64 361-362 pc 357.5 14.07 7.19 2.44 2.91 0.24 52.65 3.90 0.081 1.00 1.99 366-367 pc 362.5 15.68 6.99 2.61 3.10 0.24 59.88 4.06 0.098 1.03 4.59 371-372 pc 367.5 14.84 6.88 2.41 2.83 0.24 57.32 3.94 0.091 1.00 3.15 * Corrected piston core depth (12 cm added to account for loss of surface sediments during piston coring and a 16 cm turbidite was removed from the depth tally). 223 Table A14. (continued) Minor Element Data for Multicore (mc) and Piston Core (pc) JT96-09. Sample Depth* V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br I (cm, corr.) (pg/g) (pg/g) (pg/g) (Pg/g) (Pg/g) (Pg/g) (Pg/g) (pg/g) (pg/g) (ng/g) (ng/g) (pg/g) (Pg/g) (Pg/g) (Pg/g) (Pg/g) 0-1 mc 0.5 151 131 492 13 65 47 130 5 102 235 19 146 624 21 925 1-2 mc 1.5 2-3 mc 2.5 3-4 mc 3.5 4-5 mc 4.5 5-6 mc 5.5 159 138 456 21 60 45 130 3 89 233 19 139 624 20 167 722 6-8 mc 7.0 8-10 mc 9.0 10-12 mc 11.0 155 145 492 15 63 44 131 2 94 231 21 151 617 15 166 612 12-14 mc 13.0 14-16 mc 15.0 165 139 475 20 59 45 130 2 87 237 20 153 633 18 150 541 16-18 mc 17.0 18-20 mc 19.0 20-22 mc 21.0 173 141 542 13 63 42 126 3 87 250 21 155 598 14 111 247 22-24 mc 23.0 24-26 mc 25.0 187 132 534 20 55 39 115 3 80 258 23 156 596 18 76 114 26-28 mc 27.0 28-30 mc 29.0 30-32 mc 31.0 178 133 570 11 62 40 119 5 87 268 24 155 608 14 81 125 32-34 mc 33.0 34-36 mc 35.0 182 129 533 20 56 41 118 3 82 269 23 147 591 20 73 113 36-38 mc 37.0 38-40 mc 39.0 180 134 574 17 62 43 123 5 92 266 23 151 618 19 91 111 Sample Depth * V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br 1 (cm, corr.) (pg/g) (Pg/g) (Pg/g) (Pg/g) (Pg/g) (ng/g) (Pg/g) (Pg/g) (pg/g) (Pg/g) (Pg/g) (Pg/g) (Pg/g) (Pg/g) (Pg/g) (Pg/g) 6-7 pc 18.5 168 128 492 14 64 47 132 6 91 243 20 160 652 9 120 479 11-12 pc 23.5 168 134 527 15 62 41 126 4 88 249 23 160 576 13 57 191 16-17 pc 28.5 21-22 pc 33.5 186 135 571 16 65 44 129 7 93 257 25 148 631 15 46 90 26-27 pc 38.5 31-32 pc 43.5 176 128 568 15 66 43 121 9 91 261 22 138 640 12 47 79 36-37 pc 48.5 41-42 pc 53.5 183 129 606 15 60 42 115 4 85 280 23 135 642 12 39 55 46-47 pc 58.5 51-52 pc 63.5 194 131 644 22 52 37 106 2 80 280 23 136 618 19 16 77 56-57 pc 68.5 195 130 733 21 48 33 95 74 290 22 134 600 21 61-62 pc 73.5 179 128 682 18 46 28 87 5 70 302 21 140 599 21 2 22 66-67 pc 78.5 71-72 pc 83.5 158 126 654 17 40 20 73 4 58 319 22 212 555 20 19 14 76-77 pc 88.5 205 136 777 22 54 37 110 89 304 23 127 718 25 81-82 pc 93.5 201 128 780 21 54 38 109 5 91 317 22 120 715 24 8 28 86-87 pc 98.5 203 132 774 21 54 37 109 90 291 23 120 720 24 91-92 pc 103.5 200 129 794 22 54 37 107 5 90 305 23 118 724 28 10 15 96-97 pc 108.5 204 131 812 23 56 38 107 87 314 23 121 704 26 101-102 pc 113.5 201 130 807 22 54 38 108 5 90 308 23 122 736 24 9 30 106-107 pc 118.5 201 130 800 22 54 37 105 86 302 22 121 677 25 111-112 pc 123.5 204 132 826 23 54 38 106 5 85 298 23 121 676 23 4 22 116-117 pc 128.5 201 127 811 23 55 38 104 85 300 24 123 676 24 121-122 pc 133.5 203 127 801 23 54 37 105 4 86 290 23 124 683 24 6 24 126-127 pc 138.5 202 131 815 23 56 39 107 89 292 24 125 696 29 131-132 pc 143.5 169 113 724 20 56 37 105 6 91 286 23 128 608 18 8 22 136-137 pc 148.5 145 113 665 51 25 19 64 355 20 172 539 17 141-142 pc 153.5 157 122 669 27 42 26 78 344 20 226 598 14 146-147 pc 158.5 124 103 575 23 28 10 46 377 17 165 486 16 151-152 pc 163.5 179 113 696 23 50 33 90 276 22 131 512 16 156-157 pc 152.5 209 117 791 23 47 38 100 76 259 24 131 569 26 161-162 pc 157.5 188 86 767 20 33 30 100 8 109 248 31 173 635 31 -6 0 166-167 pc 162.5 210 115 799 23 45 37 96 72 268 25 138 556 28 171-172 pc 167.5 214 122 831 24 47 38 97 7 63 266 24 124 526 25 -7 13 176-177 pc 172.5 212 116 821 24 47 40 99 64 266 25 127 522 29 181-182 pc 177.5 187 88 802 20 32 32 109 9 126 241 33 167 673 32 -7 2 186-187 pc 182.5 214 118 827 24 45 39 98 67 265 25 126 547 31 191-192 pc 187.5 216 122 847 24 46 39 96 7 58 273 24 125 520 26 -6 0 196-197 pc 192.5 221 123 852 24 47 41 97 58 263 24 119 513 25 201-202 pc 197.5 221 123 844 25 46 41 95 8 58 268 23 121 516 26 A 11 206-207 pc 202.5 204 112 814 22 42 38 97 74 271 26 137 561 26 211-212 pc 207.5 186 116 809 18 53 41 99 7 66 273 22 126 494 14 -4 20 216-217 pc •212.5 214 122 820 24 46 40 96 57 262 23 120 496 24 221-222 pc 217.5 197 119 827 18 52 42 101 7 66 262 23 125 490 10 -11 4 226-227 pc 222.5 219 122 859 25 47 44 99 58 261 24 119 500 22 231-232 pc 227.5 186 111 789 16 51 39 94 8 61 280 23 133 486 8 -6 5 236-237 pc 232.5 223 124 860 25 47 40 97 60 258 23 119 523 27 241-242 pc 237.5 191 116 816 20 52 42 100 7 64 264 25 124 464 14 •A 7 246-247 pc 242.5 207 109 804 24 44 40 103 85 247 26 131 565 26 251-252 pc 247.5 187 117 800 16 50 38 93 9 60 286 23 135 466 9 -5 18 256-257 pc 252.5 220 125 858 25 49 40 100 62 258 24 122 524 25 261-262 pc 257.5 190 no 781 18 50 39 98 8 73 • 272 25 136 515 9 -5 7 (continued on next page) 224 Table A14. (continued) Minor Element Data for Multicore (mc) and Piston Core (pc) JT96-09 (continued). Sample Depth * V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br I (cm, corr.) (ug/g) (pg/g) (ug/g) (ng/g) (pg/g) (ug/g) (ug/g) (pg/g) (ug/g) (p.g/g) (ug/g) (ug/g) (M-g/g) (ug/g) (ug^g) (ug/g) 266-267 pc 262.5 219 118 840 24 49 39 99 65 259 24 127 539 23 271-272 pc 267.5 184 100 795 19 41 35 101 7 96 260 28 156 572 16 -3 8 276-277 pc 272.5 210 109 816 24 44 38 101 82 254 27 142 572 27 281-282 pc 277.5 194 112 810 16 52 41 101 7 78 268 26 139 525 14 -5 16 286-287 pc 282.5 211 120 825 23 47 36 95 66 279 25 138 556 28 291-292 pc 287.5 184 107 781 16 49 37 100 4 91 264 27 150 561 24 -2 19 296-297 pc 292.5 211 120 824 23 48 37 96 68 275 25 133 544 24 301-302 pc 297.5 186 105 784 17 49 36 100 4 87 258 26 144 547 16 -4 6 306-307 pc 302.5 215 118 848 24 49 39 97 65 261 25 128 536 24 311-312 pc 307.5 174 85 688 19 41 35 98 7 91 277 28 172 523 16 -3 17 316-317 pc 312.5 204 88 726 22 33 31 96 98 249 31 185 618 29 321-322 pc 317.5 178 84 701 17 37 36 105 8 115 248 33 184 595 19 -4 14 326-327 pc 322.5 206 86 758 25 34 35 106 Ul 240 33 178 644 33 331-332 pc 327.5 184 77 799 20 35 33 100 10 100 264 33 203 562 13 _2 7 336-337 pc 332.5 194 81 745 22 31 32 99 111 237 32 187 646 32 341-342 pc 337.5 173 80 696 17 39 35 100 10 110 249 31 186 573 16 -3 15 346-347 pc 342.5 351-352 pc 347.5 178 77 719 18 34 35 103 6 114 244 34 202 587 18 -5 0 356-357 pc 352.5 203 73 803 22 28 29 96 99 246 33 209 602 28 361-362 pc 357.5 176 82 702 19 35 36 102 5 112 243 32 185 581 18 -4 7 366-367 pc 362.5 199 84 764 23 33 33 99 106 242 32 190 628 33 371-372 pc 367.5 171 77 750 17 34 34 98 10 107 248 33 195 551 15 -5 -4 * Corrected piston core depth (12 cm added to account for loss of surface sediments during piston coring and a 16 cm turbidite was removed from the depth tally). 225 Table A14. (continued) Trace Metal Data for Multicore (mc) and Piston Core (pc) JT96-09. Sample Depth * Ag Cd Re Mo U (cm, corr.) (ng/g) (pg/g) (ng/g) (Pg/g) (pg/g) 0-1 mc 0.5 223 0.18 4.15 0.89 1.64 1-2 mc 1.5 213 0.19 4.08 0.64 1.20 2-3 mc 2.5 215 0.29 3.85 0.68 1.35 3-4 mc 3.5 219 0.30 5.44 0.64 1.40 4-5 mc 4.5 247 0.26 5.26 0.67 1.38 5-6 mc 5.5 248 0.26 5.64 0.60 1.40 6-8 mc' 7.0 231 0.20 5.75 0.61 1.34 8-10 mc 9.0 235 0.31 7.15 0.66 1.56 10-12 mc 11.0 187 0.28 8.39 0.63 1.66 12-14 mc 13.0 215 0.25 8.18 0.65 1.59 14-16 mc 15.0 236 0.32 10.32 0.64 1.67 16-18 mc 17.0 236 0.54 11.90 0.69 1.87 18-20 mc 19.0 295 0.86 22.37 0.71 2.27 20-22 mc 21.0 251 0.66 25.07 0.61 2.39 22-24 mc 23.0 271 0.95 29.56 0.81 3.70 24-26 mc 25.0 275 1.03 52.92 0.80 4.31 26-28 mc 27.0 254 0.71 64.45 0.80 5.36 28-30 mc 29.0 245 0.56 50.30 0.97 5.75 30-32 mc 31.0 221 0.41 37.79 0.95 4.98 32-34 mc 33.0 212 0.47 22.36 0.97 4.89 34-36 mc 35.0 236 0.51 16.77 1.04 4.06 36-38 mc 37.0 207 0.68 15.45 0.94 4.22 38-40 mc 39.0 287 0.65 16.62 1.09 3.81 Sample Depth « Ag Cd Re Mo U (cm, corr.) (ng/g) (ps/g) (ng/g) (Pg/g) (Pg/g) 6-7 pc 18.5 255 0.41 18.42 1.41 2.70 U-12pc 23.5 263 0.53 20.70 0.86 3.36 16-17 pc 28.5 220 0.73 22.13 1.13 4.75 21-22 pc 33.5 329 1.35 23.66 1.56 3.59 26-27 pc 38.5 356 1.30 23.46 3.11 4.48 31-32 pc 43.5 395 1.11 23.42 3.46 3.93 36-37 pc 48.5 358 0.67 20.78 2.61 5.04 41-42 pc 53.5 349 0.85 20.08 1.64 3.23 46-47 pc 58.5 349 na 10.44 3.71 2.90 51-52 pc 63.5 167 0.34 8.78 1.32 2.04 56-57 pc 68.5 212 0.21 8.63 0.88 1.84 61-62 pc 73.5 101 0.21 6.38 0.65 1.68 66-67 pc 78.5 90 0.20 4.79 0.73 1.86 71-72 pc 83.5 57 0.16 4.30 0.82 1.58 76-77 pc 88.5 138 na 6.34 3.33 2.52 81-82 pc 93.5 177 0.51 7.38 2.13 3.23 86-87 pc 98.5 164 0.50 5.69 na 2.00 91-92 pc 103.5 159 na 5.23 2.05 2.74 96-97 pc 108.5 145 0.42 4.70 1.84 2.14 101-102 pc 113.5 na na na na na 106-107 pc 118.5 127 0.46 4.66 1.88 2.63 lll-112pc 123.5 142 0.40 4.67 1.81 2.76 116-117 pc 128.5 127 0.39 3.20 1.25 1.83 121-122 pc 133.5 353 0.71 4.00 1.35 1.95 126-127 pc 138.5 125 0.47 3.02 1.42 2.23 131-132 pc 143.5 179 na 5.25 2.08 2.66 136-137 pc 148.5 54 0.18 1.43 0.97 1.24 141-142 pc 153.5 105 0.17 1.98 0.79 1.49 146-147 pc 158.5 na 0.07 na 0.33 na 151-152 pc 163.5 95 0.14 1.50 0.77 1.95 156-157 pc 152.5 98 0.19 1.44 0.79 1.74 161-162 pc 157.5 91 na 1.76 0.93 1.97 166-167 pc 162.5 92 0.15 1.40 0.76 1.42 171-172 pc 167.5 90 0.16 1.90 0.89 1.66 176-177 pc 172.5 80 0.13 1.19 0.79 1.47 181-182 pc 177.5 119 na 0.81 1.00 2.00 186-187 pc 182.5 87 0.13 1.31 0.85 1.26 191-192 pc 187.5 96 0.14 1.10 1.11 1.79 196-197 pc 192.5 71 0.11 1.15 0.67 1.12 201-202 pc 197.5 81 0.13 1.04 0.91 1.21 206-207 pc 202.5 87 0.15 1.09 0.69 1.30 211-212 pc 207.5 90 0.13 1.24 0.61 1.25 216-217 pc 212.5 85 0.12 1.40 0.75 1.23 221-222 pc 217.5 94 0.10 1.52 0.74 1.42 226-227 pc 222.5 82 0.13 1.34 0.96 1.22 231-232 pc 227.5 82 0.12 1.45 0.99 1.35 236-237 pc 232.5 77 0.12 1.26 0.69 1.25 241-242 pc 237.5 104 0.10 1.40 0.53 1.44 246-247 pc 242.5 107 0.13 1.32 0.80 1.54 251-252 pc 247.5 106 0.12 1.55 0.81 1.53 256-257 pc 252.5 81 0.10 1.13 0.64 1.19 261-262 pc 257.5 111 0.57 1.27 0.70 1.44 (continued on next page) Table A14. (continued) Trace Metal Data for Multicore (mc) and Piston Core (pc) JT96-09 (continued). Sample Depth * Ag Cd Re Mo U (cm, corr.) (ng/g) (Pg/g) (ng/g) (Pg/g) (ug/g) 266-267 pc 262.5 99 0.18 1.21 1.01 1.53 271-272 pc 267.5 94 0.29 1.37 1.12 1.58 276-277 pc 272.5 90 0.15 . 0.84 0.85 1.44 281-282 pc 277.5 111 0.13 1.26 0.90 1.87 286-287 pc 282.5 98 0.14 1.69 0.69 1.62 291-292 pc 287.5 394 0.34 1.94 0.95 2.13 296-297 pc 292.5 94 0.20 2.39 0.76 1.88 301-302 pc 297.5 118 0.18 1.58 0.91 1.98 306-307 pc 302.5 92 0.16 2.06 0.72 1.81 311-312 pc 307.5 140 0.46 4.12 1.31 2.18 316-317 pc 312.5 91 0.20 2.17 0.68 2.03 321-322 pc 317.5 102 0.12 1.31 0.68 1.99 326-327 pc 322.5 94 0.16 1.04 0.87 1.97 331-332 pc 327.5 97 0.19 1.00 0.90 1.76 336-337 pc 332.5 115 0.19 1.07 0.79 2.12 341-342 pc 337.5 86 0.08 1.02 1.56 2.35 346-347 pc 342.5 97 0.12 1.14 0.81 1.69 351-352 pc 347.5 86 0.18 0.88 1.13 1.97 356-357 pc 352.5 79 0.13 0.72 0.67 1.64 361-362 pc 357.5 77 0.07 0.50 1.16 2.26 366-367 pc 362.5 94 0.14 0.60 0.80 2.14 371-372 pc 367.5 105 0.15 0.63 1.65 2.10 * Corrected piston core depth (12 cm added to account for loss of surface sediments during piston coring and a 16 cm turbidite was removed from the depth tally). 227 Table A15. Geochemical Data for Trigger Core and Piston Core Tul96-03. General Geochemical Data for Trigger Core (tc) and Piston Core (pc) TuI96-03. Sample Depth Calendar Age Q 5I3N 8"C~ N[0l C[01 S[ol CaCO: %Cora QJS Opal Ba^ (m) (kyrs) (wt. %) (%o vs air) (%<. vs PDB) (wt.%) (wt.%) (wt.%) (wt.%) (wt.%) (wt. %) (ng/g) 6-7 tc 6.5 0.80 0.10 0.91 0.09 0.52 0.85 8.5 11-12 tc 11.5 0.70 0.04 0.82 0.17 2.50 0.52 13.7 16-17tc 16.5 0.79 0.05 0.68 0.24 1.83 0.46 9.9 21-22tc 21.5 0.92 0.05 0.79 0.22 2.50 0.49 10.8 26-27 tc 26.5 0.86 0.04 0.63 0.33 2.17 0.37 8.4 Sample Depth Calendar Age CI' 5"N 8"C„ N,„ C„ S,„ CaCOi %C„ C„/N Opal Baa, (m) (kyrs) (wt. %) (%o vs air) (%o vs PDB) (wt. %) (wt.%) (wt. %) (wt. %) (wt. %) 11-12 pc 11.5 1.19 0.05 0.73 0.38 2.29 0.45 9.0 16-17pc 16.5 1.30 0.05 0.71 0.25 2.50 0.41 8.5 26-27 pc 26.5 1.27 0.05 0.72 0.24 1.67 0.52 10.7 31-32 pc 31.5 1.06 0.04 0.65 0.51 1.67 0.45 10.4 41-42 pc 41.5 1.07 0.04 0.58 0.39 1.67 0.38 10.5 66-67 pc 66.5 1.05 0.05 0.68 0.23 1.83 0.46 9.7 71-72 pc 71.5 1.01 0.05 0.67 0.24 1.67 0.47 10.4 81-82 pc 81.5 1.01 0.05 0.66 0.41 1.67 0.46 9.8 91-92 pc 91.5 1.03 0.05 0.68 0.38 1.67 0.48 10.2 101-102 pc 101.5 0.93 0.05 0.74 0.37 1.67 0.54 10.8 111-112 pc 111.5 1.09 0.05 0.82 0.84 2.62 0.51 9.6 121-122 pc 121.5 0.97 0.05 0.95 0.60 3.33 0.55 10.9 131-132 pc 131.5 1.02 0.05 0.97 0.69 4.17 0.47 9.5 141-142 pc 141.5 0.97 0.04 0.84 0.60 4.17 0.34 7.9 151-152 pc No sample 161-162 pc No sample 166-167 pc 166.5 0.83 0.05 0.98 0.86 4.17 0.48 9.1 171-172 pc 171.5 0.93 0.05 1.02 0.51 4.50 0.48 9.8 181-182 pc 181.5 0.83 0.05 1.65 0.55 10.28 0.41 8.1 191-192 pc 191.5 0.96 0.05 0.97 0.28 4.66 0.41 8.8 201-202 pc 201.5 0.92 0.06 1.00 0.21 4.17 0.50 8.6 211-212 pc 211.5 0.87 0.05 0.76 0.12 3.25 0.37 8.1 221-222 pc 221.5 0.93 0.05 0.78 0.11 3.33 0.38 8.1 231-232pc 231.5 0.88 0.05 1.16 0.08 6.39 0.39 8.2 241-242 pc 241.5 0.91 0.05 0.79 0.13 3.33 0.39 8.1 251-252 pc 251.5 1.09 0.05 0.69 0.15 2.17 0.43 8.8 261-262 pc No sample 271-272 pc 271.5 0.87 0.05 0.67 0.51 2.50 0.37 7.3 281-282 pc 281.5 0.82 0.04 0.72 0.14 2.67 0.40 9.4 291-292 pc 291.5 0.84 0.06 0.78 0.28 2.50 0.48 8.3 301-302 pc 301.5 0.88 0.06 0.78 0.27 2.58 0.47 8.2 311-312 pc 311.5 0.85 0.05 0.78 0.17 2.46 0.49 8.9 321-322 pc 321.5 0.81 0.05 0.76 0.16 2.50 0.46 9.4 331-332 pc 331.5 0.87 0.06 0.75 0.32 2.50 0.45 8.0 341-342 pc 341.5 0.83 0.05 0.76 0.29 2.33 0.48 9.1 351-352 pc 351.5 0.81 0.06 0.77 0.72 2.50 0.47 8.5 361-362 pc 361.5 0.81 0.06 0.75 0.29 2.58 0.44 8.1 371-372 pc 371.5 0.80 0.05 0.76 0.28 2.50 0.46 8.3 381-382 pc 381.5 0.83 0.06 0.76 0.29 2.42 0.47 8.4 391-392 pc 391.5 0.83 0.05 0.78 0.25 2.50 0.48 8.9 401-402 pc 401.5 0.88 0.05 0.88 0.08 1.75 0.67 13.7 411-412 pc 411.5 0.85 0.05 0.62 0.11 2.50 0.32 6.5 421-422 pc 421.5 0.79 0.04 0.53 0.08 1.67 0.33 7.6 431-432 pc 431.5 0.80 0.04 0.54 0.12 1.67 0.34 8.6 228 Table A15. (continued) Minor Element Data for Trigger Core (tc) and Piston Core (pc) Tul96-03. Sample Depth V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br I (m) (Pg/g) (Pg/g) (Pg/g) (pg/g) (Pg/g) (Pg/g) (Pg/g) (pg/g) (pg/g) (pg/g) (pg/g) (Pg/g) (Pg/g) (Pg/g) (pg/g) (Pg/g) 6-7 tc 6.5 143 137 520 11 39 25 78 5 55 308 21 189 464 6 65 162 11-12 tc 11.5 173 120 692 17 45 33 81 7 56 327 22 178 491 9 22 8 16-17 tc 16.5 187 115 836 17 53 48 104 8 67 283 25 133 521 14 18 -4 21-22 tc 21.5 178 121 730 19 49 40 95 6 64 298 23 152 508 16 19 5 26-27 tc 26.5 183 114 853 18 53 47 101 8 66 294 24 132 525 15 15 6 Sample Depth V Cr Mn Co Ni Cu Zn As Rb Sr Y Zr Ba Pb Br I (m) (Pg/g) (Pg/g) (pg/g) (pg/g) (Pg/g) (Pg/g) (ug/g) (pg/g) (Pg/g) (Pg/g) (pg/g) (pg/g) (pg/g) (Pg/g) (Pg/g) (pg/g) 11-12 pc 11.5 191 115 785 21 49 45 106 8 70 272 26 140 546 14 6 12 16-17pc 16.5 186 118 738 17 48 37 93 10 63 304 23 163 520 10 5 14 26-27 pc 26.5 189 114 759 16 50 44 100 6 69 279 24 143 500 12 13 13 31-32pc 31.5 186 119 877 19 55 50 102 12 66 280 24 138 535 8 10 -12 41-42 pc 41.5 174 115 821 18 52 44 101 10 61 302 24 130 512 13 6 0 66-67 pc 66.5 192 120 889 17 56 48 101 5 63 287 25 127 531 15 9 2 71-72 pc 71.5 193 117 883 20 58 49 103 9 62 280 24 128 512 7 11 6 81-82 pc 81.5 197 116 901 19 55 49 101 6 61 277 24 126 509 14 10 -1 91-92 pc 91.5 190 115 859 20 53 49 101 8 59 276 24 123 500 9 11 -11 101-102 pc 101.5 194 115 859 18 56 48 103 8 61 279 23 127 498 11 9 2 111-112 pc 111.5 193 110 818 16 53 46 106 5 75 261 26 138 503 14 13 5 121-122 pc 121.5 196 115 832 20 50 41 99 7 72 291 26 147 529 12 15 5 131-132pc 131.5 184 102 756 18 46 39 98 4 79 292 27 158 539 17 11 21 141-142 pc 141.5 171 86 624 16 36 39 100 4 110 245 33 178 582 19 11 7 151-152 pc No sample 161-162 pc No sample 166-167 pc 166.5 183 113 765 14 49 41 95 5 68 307 27 150 504 15 13 8 171-172 pc 171.5 182 110 757 16 47 41 96 7 79 301 27 157 560 13 14 11 181-182 pc 181.5 159 106 678 17 47 36 92 9 74 363 25 163 529 12 13 18 191-192 pc 191.5 178 116 796 20 55 43 107 4 85 315 26 140 623 15 11 10 201-202 pc 201.5 189 121 876 22 59 49 113 5 83 301 26 130 612 16 13 0 211-212 pc 211.5 180 120 870 20 65 46 109 8 72 315 24 133 586 15 15 1 221-222 pc 221.5 182 118 878 21 62 47 111 6 66 316 25 132 575 13 13 4 231-232 pc 231.5 170 109 825 20 55 46 102 8 65 366 24 135 574 8 15 4 241-242 pc 241.5 167 104 824 22 53 44 107 4 74 336 23 122 575 12 15 18 251-252 pc 251.5 185 109 916 23 56 49 119 8 78 323 24 117 632 13 12 12 261-262 pc No sample 271-272 pc 271.5 168 99 843 18 50 43 105 5 73 348 22 123 648 11 13 22 281-282 pc 281.5 165 95 823 19 45 40 98 6 68 373 22 128 591 11 13 5 291-292 pc 291.5 184 110 886 19 49 45 106 9 72 342 24 125 623 10 19 31 301-302 pc 301.5 184 110 879 21 51 44 107 10 73 345 24 123 599 7 17 27 311-312pc 311.5 180 105 863 17 50 43 106 8 72 345 24 122 627 10 18 29 321-322 pc 321.5 181 107 870 19 49 46 107 8 73 348 23 123 613 10 17 12 331-332 pc 331.5 178 105 856 18 49 45 104 9 72 345 23 125 608 9 17 26 341-342 pc 341.5 184 107 881 20 49 45 107 6 72 346 23 124 619 17 17 16 351-352 pc 351.5 175 106 858 20 49 45 106 7 71 350 23 122 610 15 14 15 361-362 pc 361.5 176 104 853 17 49 45 106 9 71 348 24 121 617 11 17 12 371-372 pc 371.5 178 104 862 20 49 44 104 7 71 347 23 122 605 14 16 10 381-382 pc 381.5 185 104 882 20 49 46 107 9 73 350 22 123 632 9 16 18 391-392 pc 391.5 167 97 824 18 48 44 104 6 70 360 24 124 579 12 16 9 401-402 pc 401.5 188 103 926 18 50 50 110 7 75 348 23 117 642 8 15 19 411-412 pc 411.5 179 99 893 20 49 46 109 5 72 351 23 116 635 14 15 11 421-422 pc 421.5 178 96 912 20 48 48 107 5 71 362 24 118 635 13 14 17 431-432 pc 431.5 177 93 900 20 46 46 107 6 70 364 23 116 632 12 14 11 229 Table A16. Geochemical Data for Trigger Core and Piston Core Tul96-05. General Geochemical Data for Trigger Core (tc) and Piston Core (pc) Tul96-05. Sample Depth Depth* Calendar Age Cl" 81SN Sl3C„ Ntot C™ S10t CaC02 %C„K CoJN Opal Ba^ (cm) (cm, corr.) (kyrs) (wt. %) (%> vs air) (%o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (pg/g) 11-12 tc 11.5 2.58 0.29 2.59 0.46 0.83 2.49 8.6 21-22 tc 21.5 2.62 0.29 2.67 0.61 1.67 2.47 8.6 31-32 tc 31.5 2.44 0.26 2.45 0.53 1.67 2.25 8.6 41-42 tc 41.5 2.50 0.26 2.44 0.51 1.67 2.24 8.8 51-52 tc 51.5 2.32 0.25 2.41 0.65 2.08 2.16 8.7 61-62 tc 61.5 2.24 0.22 2.17 0.61 1.67 1.97 8.9 71-72 tc 71.5 2.15 0.21 2.10 0.63 1.67 1.90 9.0 81-82 tc 81.5 2.65 0.28 2.65 0.41 1.42 2.48 8.7 Sample Depth Depth* Calendar Age cr 8"N 8"C„ Nla C„ S101 CaCO, %C„ C„/N Opal Ba^ (cm) (cm, corr.) (kyrs) (wt. %) (%, vs air) (%<, vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (pg/g) 11-12 pc 11.5 57.5 2.39 2.73 6.52 na 0.31 2.87 0.38 1.00 2.75 8.7 16-17 pc 16.5 62.5 2.60 2.42 7.36 -21.00 0.32 2.75 0.36 1.12 2.62 8.2 21-22 pc 21.5 67.5 2.81 2.42 6.76 -21.10 0.29 2.70 0.40 0.83 2.60 8.9 26-27 pc 26.5 72.5 3.01 2.47 7.18 -21.02 0.29 2.57 0.43 1.04 2.44 8.3 31-32pc 31.5 77.5 3.22 2.46 6.96 -21.18 0.29 2.70 0.39 1.08 2.57 8.8 36-37 pc 36.5 82.5 3.44 2.14 7.14 -21.18 0.28 2.90 0.46 1.36 2.74 9.7 41-42 pc 41.5 87.5 3.66 2.25 6.84 -21.37 0.28 2.69 0.55 1.71 2.49 9.0 46-47 pc No sample 51-52 pc No sample 56-57 pc 56.5 91.5 3.84 1.96 6.87 -21.24 0.26 2.42 0.50 1.05 2.30 9.0 61-62 pc 61.5 96.5 4.06 2.11 7.11 -21.30 0.26 2.48 0.55 0.92 2.37 9.0 66-67 pc 66.5 101.5 4.28 1.98 6.98 -21.29 0.26 2.47 0.63 1.15 2.33 9.0 71-72 pc 71.5 106.5 4.50 1.92 6.83 -21.32 0.26 2.46 0.57 1.61 2.26 8.6 76-77 pc 76.5 111.5 4.72 1.73 6.86 -21.27 0.24 2.33 0.48 1.63 2.13 8.9 81-82 pc 81.5 116.5 4.94 2.11 7.32 -21.45 0.24 2.41 0.61 1.67 2.21 9.2 86-87 pc 86.5 121.5 5.16 1.76 7.25 -21.35 0.22 2.14 0.59 1.42 1.97 9.0 91-92 pc 91.5 126.5 5.38 1.83 6.79 -21.59 0.22 2.21 0.55 1.57 2.02 9.2 96-97 pc 96.5 131.5 5.60 1.64 7.05 -21.40 0.22 2.11 0.60 1.37 1.94 8.9 101-102 pc 101.5 136.5 5.82 1.79 7.01 -21.64 0.21 2.06 0.60 1.08 1.93 9.1 106-107 pc 106.5 141.5 6.04 1.67 7.17 -21.50 0.21 2.06 0.64 1.17 1.92 9.0 111-112 pc 111.5 146.5 6.26 1.72 7.05 -21.69 0.21 2.05 0.47 1.33 1.89 9.2 116-117pc 116.5 151.5 6.48 1.71 7.07 -21.92 0.20 2.01 0.56 1.26 1.86 9.2 121-122 pc No sample 126-127 pc No sample 131-132 pc 131.5 166.5 7.14 1.63 7.27 -21.73 0.20 2.00 0.57 1.67 1.80 9.2 136-137 pc 136.5 171.5 7.36 1.32 7.11 -21.63 0.17 1.81 0.61 2.17 1.55 8.9 141-142 pc 141.5 176.5 7.58 1.65 7.41 -21.82 0.19 2.05 0.48 2.33 1.77 9.1 146-147 pc 146.5 181.5 7.80 1.60 6.66 -21.73 0.19 1.97 0.71 2.30 1.70 9.0 151-152 pc 151.5 186.5 8.02 1.71 7.57 -21.88 0.20 2.15 0.58 2.39 1.86 9.1 156-157 pc 156.5 191.5 8.24 1.65 7.44 -21.65 0.20 2.05 0.50 2.15 1.80 9.0 161-162 pc 161.5 196.5 8.46 1.68 8.02 -21.85 0.20 2.11 0.55 2.17 1.85 9.2 166-167 pc 166.5 201.5 8.68 1.73 7.71 -21.64 0.20 2.16 0.73 3.25 1.77 9.0 171-172 pc 171.5 206.5 8.90 1.91 8.45 -21.71 0.22 2.32 0.55 2.95 1.97 9.1 176-177pc 176.5 211.5 9.12 1.68 7.76 -21.57 0.21 2.30 0.52 3.17 1.92 9.0 181-182 pc 181.5 216.5 9.34 1.75 8.27 -21.79 0.22 2.43 0.63 3.39 2.02 9.2 186-187 pc 186.5 221.5 9.56 1.64 8.26 -21.65 0.22 2.40 0.53 3.27 2.00 9.3 191-192 pc 191.5 226.5 9.78 1.72 7.92 -21.88 0.22 2.50 0.49 3.33 2.10 9.3 196-197 pc 196.5 231.5 10.00 1.64 8.03 -21.75 0.20 2.25 0.47 3.65 1.81 9.1 201-202 pc 201.5 236.5 10.22 2.00 7.91 -21.95 0.22 2.46 0.57 4.04 1.98 9.1 206-207 pc 206.5 241.5 10.44 1.42 6.73 -22.54 0.11 1.41 0.31 2.57 1.10 9.6 211-212 pc 211.5 246.5 10.66 1.39 6.54 -22.89 0.10 1.20 0.38 1.87 0.97 9.7 216-217 pc 216.5 251.5 10.88 1.29 6.47 -22.80 0.09 1.24 0.30 2.70 0.92 10.2 221-222 pc No sample 226-227 pc No sample 231-232 pc Turbidite 1.03 6.79 -22.72 0.08 1.04 0.24 2.43 0.74 9.6 236-237 pc Turbidite 0.78 6.18 -23.19 0.06 0.87 0.19 2.38 0.59 10.1 241-242 pc 241.5 257.5 27.97 0.97 4.93 -23.79 0.04 0.50 0.16 0.90 0.39 10.2 246-247 pc 346.5 362.5 28.38 0.94 4.87 -23.15 0.05 0.61 0.23 1.45 0.43 9.1 251-252 pc 251.5 267.5 28.78 1.07 4.85 -23.01 0.05 0.69 0.14 1.83 0.47 9.4 256-257 pc 256.5 272.5 29.18 1.04 5.30 -22.75 0.06 0.88 0.24 2.72 0.56 9.1 261-262 pc 261.5 277.5 29.58 1.07 4.94 -22.44 0.08 1.25 0.19 3.50 0.83 10.8 266-267 pc 266.5 282.5 29.98 1.05 5.01 -22.12 0.08 1.26 0.27 4.21 0.75 8.9 271-272 pc 271.5 287.5 30.38 1.03 5.78 -22.32 0.08 1.18 0.45 4.17 0.68 8.7 276-277 pc 276.5 292.5 30.79 0.96 4.11 -22.44 0.07 1.07 0.24 3.72 0.62 8.7 281-282 pc 281.5 297.5 31.19 1.13 5.86 -22.30 0.09 1.32 0.31 4.31 0.80 9.1 286-287 pc 286.5 302.5 31.59 1.07 5.87 -22.14 0.09 1.40 0.34 4.75 0.83 9.4 291-292 pc 291.5 307.5 31.99 1.08 6.15 -22.15 0.11 1.67 0.24 5.00 1.07 9.5 296-297 pc 296.5 312.5 32.39 1.02 5.83 -22.28 0.09 1.39 0.19 4.66 0.83 9.1 301-302 pc 301.5 317.5 32.79 0.83 5.57 -22.17 0.09 1.36 0.50 4.33 0.84 9.3 306-307 pc 306.5 322.5 33.20 1.04 5.98 -22.09 0.10 1.46 0.30 5.10 0.84 8.8 311-312 pc 311.5 327.5 33.60 1.11 5.91 -22.14 0.11 1.69 0.38 5.79 1.00 8.9 316-317 pc 316.5 332.5 34.00 1.13 6.35 -22.12 0.11 1.66 0.35 5.11 1.05 9.4 321-322 pc 321.5 337.5 34.40 1.10 6.27 -22.43 0.10 1.50 0.41 4.26 0.99 9.8 326-327 pc No sample 331-332 pc No sample 336-337 pc 336.5 347.5 35.20 1.14 6.67 -22.03 0.12 1.78 0.56 5.52 1.11 9.1 341-342 pc 341.5 352.5 35.61 1.11 5.91 -22.33 0.10 1.56 0.74 5.20 0.94 8.9 (continued on next page) 230 Table A16. (continued) General Geochemical Data for Trigger Core (tc) and Piston Core (pc) Tul96-05 Sample Depth Depth * Calendar Age ei' 5,SN 5I3C„ N„ Ca St« CaC02 0/oC„„ C^/N Opal Ba^ (cm) (cm, corr.) (kyrs) (wt. %) (%o vs air) (%o vs PDB) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (wt. %) (pg/g) 346-347 pc 346.5 357.5 36.01 1.10 5.87 -22.18 0.10 1.60 0.43 5.61 0.93 9.2 351-352 pc 351.5 362.5 36.41 1.06 5.94 -22.16 0.10 1.58 0.58 6.08 0.85 8.3 356-357 pc 356.5 367.5 36.81 1.05 5.79 -22.15 0.11 1.72 0.72 6.48 0.95 9.0 361-362 pc 361.5 372.5 37.19 1.11 6.81 -21.97 0.12 2.09 0.81 7.50 1.19 9.7 366-367 pc 366.5 377.5 37.50 1.17 6.38 -21.94 0.13 2.00 0.72 6.95 1.17 9.3 371-372 pc 371.5 382.5 37.81 1.20 6.94 -22.09 0.12 1.96 1.05 7.25 1.09 9.2 376-377 pc 376.5 387.5 38.12 1.07 5.21 -22.43 0.08 1.36 0.61 5.39 0.71 8.8 381-382 pc 381.5 392.5 38.43 0.98 6.08 -22.75 0.08 1.32 0.69 4.87 0.73 9.0 386-387 pc 386.5 397.5 38.74 1.00 5.41 -22.56 0.08 1.27 0.55 4.70 0.71 9.1 391-392 pc 391.5 402.5 39.04 0.91 5.79 -23.04 0.06 1.02 0.48 3.32 0.62 9.9 396-397 pc 396.5 407.5 39.35 1.08 4.78 -22.65 0.11 1.85 0.56 3.86 1.39 12.2 401-402 pc 401.5 412.5 39.66 0.96 5.29 -22.37 0.09 1.28 0.71 4.17 0.78 9.0 406-407 pc 406.5 417.5 39.97 1.03 5.10 -22.16 0.10 1.51 0.44 5.15 0.89 9.0 411-412 pc 411.5 422.5 40.28 1.05 5.78 -22.11 0.11 1.63 0.64 5.08 1.02 9.4 416-417 pc 416.5 427.5 40.59 421-422 pc No sample 426-427 pc No sample 431-432 pc 431.5 432.5 40.89 1.01 6.21 -22.02 0.12 1.82 0.59 5.83 1.12 9.5 436-437 pc 436.5 437.5 41.20 0.81 5.94 -22.18 0.07 1.23 0.82 6.62 0.44 6.2 441-442 pc 441.5 442.5 41.51 1.17 6.37 -22.11 0.13 1.82 0.46 4.74 1.25 9.5 446-447 pc 446.5 447.5 41.82 1.12 5.82 -21.99 0.11 1.71 0.43 5.55 1.05 9.2 451-452 pc 451.5 452.5 42.13 1.04 5.77 -22.20 0.12 1.81 0.46 6.16 1.07 9.2 456-457 pc 456.5 457.5 42.44 1.01 5.84 -22.11 0.11 1.79 0.42 6.54 1.00 9.3 461-462 pc 461.5 462.5 42.75 1.04 5.69 -22.09 0.12 1.82 0.63 6.25 1.07 9.2 466-467 pc 466.5 467.5 43.05 1.03 6.17 -22.06 0.11 1.72 0.34 6.15 0.98 9.2 471-472 pc 471.5 472.5 43.36 1.05 6.28 -21.95 0.13 1.99 0.52 5.83 1.29 9.6 476-477 pc 476.5 477.5 43.67 1.04 6.84 -22.02 0.14 2.00 0.44 6.12 1.27 9.3 481-482 pc 481.5 482.5 43.98 1.06 5.81 -22.08 0.11 1.88 0.43 6.75 1.07 9.3 486-487 pc 486.5 487.5 44.29 1.01 6.20 -22.10 0.11 1.82 0.44 7.16 0.96 9.1 491-492 pc 491.5 492.5 44.60 1.04 5.66 -22.28 0.12 1.93 0.75 7.44 1.04 8.9 496-497 pc 496.5 497.5 44.90 1.03 6.23 -22.08 0.11 1.95 0.39 7.56 1.04 9.1 501-502 pc 501.5 502.5 45.21 0.99 6.07 -21.97 0.12 2.05 0.60 7.75 1.12 9.3 506-507 pc 506.5 507.5 45.52 0.96 6.59 -22.01 0.12 2.07 0.58 8.30 1.08 9.2 * Corrected piston core depth. 231 Table A16. (continued) Major Element Data for Trigger Core (tc) and Piston Core (pc) Tul96-05. Sample Depth* A120, FeA K20 MgO PA Si02 CaO MnO TiOj Na20 (cm, corr.) (%) (%) (%) (%) (%) (%) (%) (%) (%) (%) 11-12 tc 14.48 6.56 1.89 3.13 0.19 56.39 2.45 0.056 0.83 3.29 21-22 tc 14.67 6.65 1.91 3.15 0.18 55.86 2.77 0.063 0.83 4.11 31-32 tc 14.47 6.45 1.91 3.15 0.18 56.76 2.76 0.058 0.83 3.16 41-42tc 14.64 6.39 1.91 3.14 0.18 57.20 2.84 0.062 0.83 2.28 51-52 tc 14.43 6.42 1.85 3.08 0.18 56.91 2.99 0.058 0.81 2.62 61-62 tc 14.75 6.49 1.90 3.11 0.19 58.07 2.96 0.072 0.84 2.97 71-72 tc 14.63 6.43 1.94 3.08 0.18 57.87 2.88 0.071 0.83 2.85 81-82 tc 14.36 6.33 1.78 3.05 0.19 56.80 2.67 0.065 0.83 3.41 Sample Depth* A1A FeA K20 MgO PA Si02 CaO MnO Ti02 Na20 (cm, corr.) (%) (%) <%) (%) (%) (%) (%) (%) (%) (%) ll-12pc 57.5 14.07 6.47 1.85 3.03 0.19 55.95 2.43 0.067 0.82 3.04 16-17 pc 62.5 14.66 6.08 1.83 3.10 0.19 57.28 2.55 0.061 0.82 5.75 21-22 pc 67.5 14.14 6.45 1.84 3.04 0.18 55.05 2.35 0.062 0.82 5.49 26-27 pc 72.5 31-32pc 77.5 14.13 6.23 1.79 3.04 0.19 55.96 2.65 0.066 0.82 3.06 36-37 pc 82.5 41-42 pc 87.5 14.24 6.58 1.84 3.09 0.19 54.90 2.81 0.059 0.82 2.98 46-47 pc 51-52 pc 56-57 pc 91.5 61-62 pc 96.5 14.28 6.52 1.96 3.07 0.18 55.04 2.41 0.066 0.83 7.30 66-67 pc 101.5 14.91 6.55 2.01 3.18 0.18 57.19 2.57 0.062 0.83 5.02 71-72 pc 106.5 14.26 6.36 1.90 3.05 0.18 56.28 2.85 0.066 0.82 7.47 76-77 pc 111.5 15.02 6.13 1.96 3.06 0.18 58.25 2.94 0.062 0.82 5.29 81-82 pc 116.5 13.81 6.46 1.90 2.99 0.18 54.72 2.96 0.064 0.81 9.13 86-87 pc 121.5 14.96 6.35 2.06 3.09 0.18 58.23 2.81 0.063 0.82 5.26 91-92 pc 126.5 14.54 6.43 1.97 3.09 0.19 57.21 2.80 0.068 0.83 5.61 96-97 pc 131.5 15.02 6.42 2.10 3.14 0.18 58.46 2.84 0.065 0.82 8.36 101-102 pc 136.5 14.40 6.41 1.96 3.07 0.19 57.41 2.71 0.069 0.82 6.32 106-107 pc 141.5 15.18 6.52 2.10 3.17 0.18 58.23 2.56 0.064 0.83 5.35 111-112 pc 146.5 14.66 6.39 2.03 3.10 0.19 57.35 2.72 0.070 0.84 4.33 116-117 pc 151.5 15.20 6.33 2.07 3.13 0.19 58.48 2.74 0.066 0.83 5.63 121-122 pc 126-127 pc 131-132 pc 166.5 14.57 6.43 1.97 3.06 0.19 57.44 2.94 0.065 0.83 4.10 136-137 pc 171.5 14.89 6.20 1.97 2.98 0.18 58.85 3.33 0.069 0.81 6.49 141-142 pc 176.5 14.57 6.55 1.94 3.10 0.19 56.92 3.30 0.070 0.85 2.90 146-147 pc 181.5 14.84 6.68 2.00 3.12 0.19 57.78 3.42 0.067 0.81 5.03 151-152 pc 186.5 14.81 6.90 2.01 3.27 0.19 55.93 3.24 0.066 0.85 2.70 156-157 pc 191.5 15.54 6.71 2.06 3.31 0.19 57.63 3.15 0.067 0.86 5.30 161-162 pc 196.5 14.70 6.77 2.02 3.22 0.20 - 55.65 3.27 0.067 0.84 3.26 166-167 pc 201.5 15.84 7.38 2.19 3.38 0.20 60.17 4.24 0.072 0.88 5.17 171-172 pc 206.5 14.63 6.82 2.02 3.24 0.19 55.09 3.59 0.066 0.84 2.35 176-177pc 211.5 15.36 6.70 2.03 3.30 0.19 55.97 3.83 0.067