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The late Quaternary history of primary productivity in the subarctic east Pacific McDonald, Darcy 1997

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THE LATE QUATERNARY HISTORY OF PRIMARY PRODUCTIVITY IN THE SUBARCTIC EAST PACIFIC by Darcy McDonald B.Sc, The University of British Columbia, 1993 A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF MASTER OF SCIENCE in THE FACULTY OF GRADUATE STUDIES (Department of Earth and Ocean Sciences) We accept this thes^asj conforming to the required standard The University of British Columbia March 1997 © Darcy W. McDonald, 1997 In p resen t ing this thesis in partial fu l f i lment of the requ i remen ts fo r an advanced degree at the Universi ty o f Brit ish C o l u m b i a , I agree that t h e Library shall make if/, f reely available f o r re ference and s tudy. I fu r ther agree that permiss ion fo r ex tens ive c o p y i n g of this thesis fo r scholar ly pu rposes may be g ran ted by the head o f m y d e p a r t m e n t or by his o r her representat ives. It is u n d e r s t o o d that c o p y i n g o r pub l i ca t i on o f this thesis for f inancial gain shall n o t be a l l o w e d w i t h o u t my w r i t t e n permiss ion . V ' D e p a r t m e n t o f t h e Univers i ty of British C o l u m b i a Vancouver , Canada Date DE-6 (2/88) Abstract Palaeoceanographic work in the subarctic northeast Pacific has previously been limited by the common presence of turbidites and by the relatively shallow CCD. ODP Site 887 is on the Patton Murray Rise, an elevated plateau which is largely isolated from these effects. This area is centered under the Alaska Gyre, a region characterized by the domal upwelling of nutrient-rich waters. Sections from Holes 887A and 887C were spliced into the 887B profile to obtain a continuous composite record. Although calcite is not continuously distributed throughout the drilled section, the benthic foram 6 l s O profile from this record closely reflects the global (SPECMAP) signal. Hence, based on 6 l s O data and on radiocarbon dates from site survey cores,* a detailed chronology spanning the past 750 ky was developed. The controls on productivity in this region remain unclear since nitrate is perennially present in excess and does not limit primary production. However, past increases in productivity and rapid settling of biogenic matter are suggested by the episodic accumulation of diatomaceous oozes, and by high-resolution Si/Al, Ba/Al, C 0 r g , and CaC03 profiles. Moreover, episodic shoaling of the redox boundary produced by increased organic matter flux to the seafloor, as during the deposition of the diatomaceous oozes, is suggested by trace metal and Mn enrichments. Significant 613Corg maxima in the major diatomaceous bands suggest that productivity events have occurred which were sufficient to draw down mixed-layer PCO2. Thus, in contrast to the near-neutral ocean-atmosphere PCO2 gradient which exists today, an episodic sink for atmospheric CO 2 may have existed in the Gulf of Alaska in the past. Surface water temperature and salinity effects are evident in negative regional surface water 6 l s O variations. The episodes of enhanced productivity at Site 887 occur synchronously with such 6 1 8Qocal excursions, suggesting a direct link to low salinity and/or warming events in the Gulf of Alaska which occur on sub-Milankovitch timescales. This suggests that significant climate variability in the Gulf of Alaska has been superposed on glacial-interglacial cycles through the Brunhes Chron. Temporal variations in the availability of iron may also have played a key role in constraining export production through time. Episodic input of iron, which may have stimulated productivity, may have occurred from either Asian sources via the atmosphere or more local sources via meltwater input. Table of Contents Abstract , ii Table of Contents iv List of Tables , vi List of Figures viii Acknowledgements xi 1. Introduction 1 2. Regional Setting 6 2.1 Hydrography 6 2.2 Dynamics of Primary Production 10 2.2.1 Iron Limitation 13 3. Background 18 3.1 Barium as a Productivity Tracer 18 3.1.1 Ba Distributions 18 3.1.2 Alternate Ba Sources 22 3.1.3 Diagenesis 23 3.2 Biogenic Silica 25 3.3 Carbonate 29 3.3.1 Dissolution vs. Productivity 29 3.3.2 Carbonate Cycles in the Pacific 31 3.3.2 North Pacific Carbonate Records 34 3.4 Carbon Isotopes 37 3.4.1 Some Constraints on the 6 1 3 C Record 40 3.4.1.1 Gas Exchange 40 3.4.1.2 Diagenesis, Vital Effects 41 3.4.2 Carbon Isotopes in Organic Matter 43 3.4.2.1 Organic Carbon in Marine Sediments 43 3.4.2.2 6 1 3 Gorg and C0 2 44 3.4.2.3 Caveats 45 3.4.2.4 Metabolic Effects 47 3.5 Oxygen Isotopes 49 4. Chronostratigraphy 52 4.1 Stratigraphy 52 4.2 Composite Depth Model 53 4.3 Chronology 54 4.3.1 Comparison with Biostratigraphic Data 63 5. Results and Discussion 68 5.1 Sedimentary Geochemistry 68 5.1.1 Si/Al 68 5.1.2 Mn/Al 74 5.1.3 Ba/Al 79 5.1.4 Organic Carbon 81 5.2 Carbonate Records 85 5.2.1 Carbonate Presence 85 5.2.2 Foraminiferal Abundance 88 5.2.2.1 Planktonic 89 5.2.2.2 Benthic 92 5.3 Foraminiferal Isotopes 94 5.3.1 Oxygen 94 iv 5.3.2 Carbon 101 S^S^Cbrg 107 5.5 Environmental Changes and Palaeoproductivity I l l 5.5.1 Carbon Dioxide 114 5.5.1.1 Calculation of [C02(aq)] from the 6 1 3 C o r g Record 114 5.5.1.2 PC0 2 117 5.5.2 Evidence of Temperature and Salinity Variations 120 5.5.2.1 Relationships Between 6 1 8 0 and Transfer Function Data 120 5.5.2.2 Alkenone-Derived Temperatures 124 5.5.2.3 C. davisiana Records 127 5.5.2.4 Mechanisms of Change 128 5.5.2.5 High Frequency Climate Variability 130 5.5.3 Factors Affecting Surface Ocean Productivity 132 5.5.3.1 Upwelling 132 5.5.3.2 Surface Water Properties 134 5.5.3.3 Iron 137 6. Summary and Conclusions 140 Bibliography 144 Appendix 1: Methods 157 A l . l Sampling Strategy and Sample Preparation 158 A1.2 X-ray Fluorescence Spectrometry 159 Al.2.1 Sample Preparation 159 Al.2.1.1 Minor Elements 159 Al.2.1.2 Major Elements 159 Al.2.2 Analysis 160 A 1.3 Total Carbon and Nitrogen 170 A 1.4 Coulometry 174 A1.5 Mass Spectrometry 176 A 1.5.1 Calcite Analysis 176 Al.5.2 Organic Carbon 177 A1.6 Foraminifera Counts 178 Appendix 2: Data 179 v List of Tables Table 4.1. a) sections missing from 887B relative to 887A and 887C, based on correlation of GRAPE records; b) actual core depths + missing sections in (a) 53 Table 4.2. Age control points for Hole 887B. Sedimentation rates represent the intervals following; e.g., sed. rate for 0.3 to 0.6 mbsf = 4.6 cm/kyr 57 Table 4.3. 887B age control points based on C. davisiana patterns and last occurence datums of radiolarians 64 Table 5.1. Glacial-interglacial 6 1 8 0 shifts hi core 887B 97 Table 5.2. Glacial-interglacial 6 1 3 C shifts in core 887B 105 Table A l . l . Criteria for selection of foraminiferal samples used in counts and in stable isotope analyses 158 Table A1.2. XRF instrument parameters for elemental analysis. Detection by gas flow counter using 90% Ar and 10% CH4; kv = 60, ma = 40 160 Table A1.3. Measured and recommended major element concentrations for standards used in XRF analysis 162 Table A1.4. Measured and recommended minor element concentrations for standards used in XRF analysis 164 Table A1.5. Measured and recommended Ba concentrations for standards used in specific Ba XRF run 168 Table A1.6. XRF analytical precision for major elements determined by seven runs of an individual disc and by analysing six separate discs of a homogenized sample 169 Table A 1.7. XRF analytical precision for minor elements determined by 6 runs of each of two separate discs of a single homogenised sample 169 Table A 1.8. Elemental analysis standards for C and N determination 171 Table A1.9. Measured and recommended values for standards used in C and N analysis 172 Table ALIO. Analytical precision for C and N analysis estimated by running six separate splits of a homogenized sample 173 v i Table A2.1. Chemical data from Hole 887B with general lithology (legend at bottom) 180 Table A2.2. Composite depth model for Hole 887B with spliced sections from Holes 887A and 887C 189 Table A2.3. Composite foraininiferal abundance and isotopic data for Hole 887B 194 Table A2.4. Measured and calculated (transfer function-based) 6 l s O (%o) based on data from cores 887B and PAR87A-1, 2, and 10 201 List of Figures Figure 1.1 (a) North Pacific map showing the location of Site 887 and Station P; (b) bathymetry of the Patton-Murray Seamount Group 2 Figure 2.1. Temperature isotherms from the Vertex Alaska transect 7 Figure 2.2. Nitrate, phosphate, silicate, and temperature transects at 5 m depth vs. latitude 7 Figure 2.3. Profiles of (a) salinity, (b) temperature, and (c) oxygen for the upper 200 m of the water column at Station P 8 Figure 2.4. Dissolved Fe profile at Station P, shown with oxygen and nitrate distributions 14 Figure 2.5. A dissolved Fe section from the Vertex Alaska transect illustrating an offshore-inshore trend to higher Fe levels 14 Figure 3.1. Seawater Ba vs. alkalinity, normalized to a salinity of 35 over 9 widely spaced GEOSECS sites 19 Figure 3.2. C o rg/Ba in sediment trap samples vs. depth for three ocean areas (best fit curves) 19 Figure 3.3. Bathymetric variations in equatorial Pacific %CaCC>3 through time, contoured in six isopleths at the 90, 80, 60, 40, 20, and 10% CaCC>3 (deepest water depth) levels 33 Figure 3.4. Mass accumulation rates and weight percentages of biogenic opal,TOC, andCaC03 over the last 1 my. at ODP Site 882 36 Figure 4.1. SPECMAP composite S 1 8 0 from Imbrie et al., as modified by Shackleton et al. (1990) 56 Figure 4.2. 887B time series of benthic composite 6 l s O (see Section 5.3.1) and G. bulloides 6 1 8 0 with isotopic stage boundaries noted 58 Figure 4.3. Age-depth relationship in 887B. Sedimentation rates between control points average 6.6 cm/ky, with a range of 1.7 to 15.1 60 Figure 4.4. Excess 2 3 0 T h ( 2 3 0Th 0 ex),%CaCO3, and Si/Al values for Core PAR 87-1 62 Figure 4.5. Distributions of %C. davisiana, benthic 6 1 8 0, and local A 6 l s O with depth in 887B 66 Figure 5.1. Ice-rafted debris (IRD) abundance, Si/Al weight ratio, %CaC03,andCaCC>3 accumulation rate (g/m2a) vs. depth at Site 887 69 Figure 5.2. 887B time series of Mn/Al, Ba/Al, 6 1 3 Co r g (%o), %Corg, Si/Al, and%CaC03 72 Figure 5.3. Schematic diagram of dissolved and solid phase Mn profiles in a hypothetical steady-state system (redrawn from Froelich et al., 1979) 74 Figure 5.4. 887B time series of Mn/Al, Fe/Al, Mn and Fe (ppm), and ice-rafted debris (IRD) abundance 76 Figure 5.5. Schematic representation of Mn carbonate formation 78 Figure 5.6. 887B time series of S 1 3 Co r g (%0), Corg/N, %C0rg, and C o r g accumulation rates (g/m2a) 82 Figure 5.7. Results of cross-spectral analysis of the 887B CaC03 record with the SPECMAP stack 86 Figure 5.8. 887B foraminiferal abundance shown with Si/Al and %CaCC>3 profiles for reference 90 Figure 5.9. Benthic (Gyroidinoides, Uvigerina, and Cibicides spp.) and planktonic (G. bulloides) 6180(%o) vs. depth at Site 887, shown with replicate values 95 Figure 5.10. Paired benthic 6180(%o) data from 887B. Cibicides mean offset from Uvigerina is 0.67 ±0.21 (Is) 94 Figure 5.11. 887B time series of benthic composite 6180(%o) and "local" 6180(%o) shown with isotopic boundaries and the SPECMAP stack (modified from Imbrie et al., 1984) 99 Figure 5.12. Time series of 887B benthic and planktonic 6 1 3 C (%o) 103 Figure 5.13. Scatter plot of 887B 6 1 3 Co r g vs. C o r g / N . Mean 6 1 3 Co r g = -22.8 %o,s=1.2 108 Figure 5.14. 887B geochemical, isototopic, and preliminary alkenone temperature records to 200 ka 112 Figure 5.15. 887B 6 1 8 Q O C a l , D. seminae (%), Si/Al, CaC0 3, and GRIP 6 l s O records 113 Figure 5.16. 887B 613Cbrg and estimated dissolved CO2 concentrations ([C02(aq)]) using the relationship derived by Rau (1994) 115 ix Figure 5.17. Estimated surface water [CC»2(aq)] andPCC>2 records derived from the 6 1 3 C o r g values in Hole 887B 118 Figure 5.18 a) PAR87A-10 time series of measured and calculated Ddc relative to modern values. Measured and calculated ASc are based on transfer function salinity and temperature data i n (b) 122 Figure 5.19. Coherency estimates for measured vs. calculated 6 l s O , core PAR87A-10 123 Figure 5.20. Transfer function (TF) and alkenone-derived sea surface temperatures through the past 30 ky 125 Figure A l . l . % C o r g vs. % total N i n Site 887 samples. Linear regression is shown and is indicated by the solid line 173 x Acknowledgements The path through graduate school has for me been filled with both discovery and difficulty. Were it not for the help and guidance of those dedicated souls listed here, this work could not have been completed. Firstly, thanks to Bente Nielsen for her many hours spent at spent extracting dependable isotope data from limited amounts of sample, and to Maureen Soon for her help, instruction, and immutable good humor in the lab. Thanks are also due to Ian Wattson for many long (though efficient) hours spent in the lab manipulating 887 samples. In progressing through this study, my bewilderment at strange palaeoceanographic goings on led me to draw on the sagacity of the following scientists who in addition provided indispensable data: John Crusius ( 2 3 0 T h data), Greg Cowie (alkenone data), Connie Sancetta (diatom data), Anne deVernal (transfer function work), and Joe Morley (radiolarian data). I must also thank John Farrell and Steve Calvert for their obliging and instructive discussions. My deepest appreciation goes to Tom Pedersen to whom I am grateful for invaluable advice, direction, and support through the course of this project. Lastly, but certainly not least, thanks to Shelley Jo and Andrea Jean for proving the alternate, and previously unconsidered hypothesis of meaningful life outside of oceanography. "Strange and beautiful things were brought to us from time to time which seemed to give us a glimpse of some unfamiliar world." C. Wyville Thomson: The Challenger Expedition, 1895 xi 1. Introduction Although the Pacific is the largest of the major ocean basins, the subarctic region remains poorly understood in terms of Quaternary history. Palaeoceanographic work in this area has so far been limited by a scarcity of cores. Moreover, continuous downcore records remain an enigma, owing primarily to the dissolution of biogenic calcite due to the relatively shallow calcite compensation depth (CCD; ~ 2400 m in the northwest, deepening by at least 1 km in the east; Rea et al., 1995) and to common turbidite sequences south of the eastern Aleutian arc and west of the continental shelf. As a result, little information is available about the history of surface productivity and its relationship to North Pacific sedimentary properties. A small number of cores have been taken from plateau areas in an attempt to minimize regional turbidite and CCD effects; from these it is clear that variations in sediment properties reflect profound changes in climate, productivity, and water column chemistry in the past (e.g., Zahn et al., 1991a; Keigwin et al., 1992). This thesis investigates such changes in the subarctic northeast Pacific Ocean through the study of a sedimentary section from ODP Site 887, collected during Leg 145 from the Patton-Murray Seamount Group in the central Gulf of Alaska (Figures la and lb). Site 887 lies beneath the Alaska Dome upwelling cell and thus chronicles past changes in this unique sub-polar system. The section (887B) spans the past 750,000 years and, via the application of geochemical and stable isotopic data, provides a glimpse of environmental processes active through most of the Brunhes Chron. 1 a ) . 180° 150° 60° 40° - ' i . f V i „ , Gulf of Alaska Bering Sea Site 887 Patton-Murray Seamounts 180° Alaska Gyre i Station P I , y itinn P <r Subarctic Current N. Pacific Current California Current i i 2 5 0 km 150° 120°W 60°N 50° 40° 120°W Figure 1.1 a) North Pacific map showing the location of Site 887 (54° 21.92'N, 148° 26.78'W: water depth 3647 m), Station P located at 50°N, 145°W; b) bathymetry of the Patton-Murray Seamount Group with positions of Sites 887, PAR87-01, and PAR87-10 (same locale as 887; bathymetry courtesy of ODP). 2 3 Specifically, this work attempts to determine: 1) the nature of temporal variations in the deposition of biogenic phases. Commensurate increases in Si/Al and CaC03 with Ba/Al and 6 1 3Co rg values indicate that throughout the Brunhes Chron episodic increases in primary production paralleled by rapid settling of biogenic matter are manifest in diatom ooze and possibly carbonate-rich strata. 2) whether primary production influences the exchange of CO2 between ocean and atmosphere in this area over glacial-interglacial time scales. Significant ° 1 3Corg maxima in the major diatomaceous bands suggest that productivity events have occurred which were sufficient to draw down mixed-layer PCO2. 3) the geochemical response of the sediments to increased flux of organic matter to the seafloor. Authigenic Cd, Mo, and Re enrichments in site survey cores suggest that shoaling of the redox boundary is produced by enhanced organic matter flux to the seafloor, as during the rapid deposition of diatom oozes (Bedard, 1992; McDonald, 1993; J. Crusius, unpublished data). This is confirmed by the occurrence of downcore Mn enrichments which are coeval with intervals of increased biogenic deposition. 4) the effect of long-term environmental changes on primary productivity. Isotopic data provide evidence of abrupt changes in surface salinity and/or temperature which occur synchronously with increased fluxes of biogenic matter to the seafloor. Such changes may be influenced by local processes and occur independently of sub-stage glacial cycling. Though each of the tracers examined in this study provides a unique record of change, the diversity of physical and chemical effects at work in ocean systems limits the description of present and past ocean processes to the sensitivity of the tracer which is used. Although each proxy is sensitive to different controls and thus offers a unique record of change, the distribution 4 of an individual proxy may contradict evidence seen in other tracer patterns. Hence, in order to constrain ancillary influences on the signal or process in question, the interpretation of data from a suite of tracers is required. This is illustrated in the estimation of palaeo-CC>2 levels in the surface ocean which are influenced by a number of parameters including the 6 1 3 C o f both primary producers and seawater, temperature, and salinity. Moreover, understanding the controls on these parameters calls for a knowledge of other factors such as nutrient distributions and utilization, and fluxes of particulate and dissolved matter. Each factor is linked by larger physical and chemical cycles which are at present infeasible to describe based on the limited palaeoceanographic data at hand from the northeast Pacific. Comprehension of the nature of biogeochemical cycles becomes increasingly important as we enter an era where the burning of fossil fuels and changes in land use may affect climate change. Understanding ocean processes is crucial to this awareness since the record of long term change is contained in marine sediments. To this end, further expansion and interpretation of a "palaeo" database from high-latitude sediment cores is requisite to constraining models of global change. 5 2. Regional Setting 2.1 Hydrogra phy The subarctic Pacific region stretches across the Pacific Ocean north of 45°N and is bounded in the east by the Alaskan Coast. Cyclonic circulation of the Alaska Gyre (Figure 1.1), induced by the easterly flow of the Subarctic Current and the westward flowing Alaska Current, sustains regional upwelling of cold, nutrient-replete waters in the area known as the Alaska Dome, centered near 54°N (Figures 2.1 and 2.2). In addition to the upward movement of deep water, upwelling near the sea surface is forced by the wind-stress curl. Strong westerly winds are induced by a low pressure system off the Aleutians in winter while in summer, northward migration of a high pressure zone off southern California produces weaker southwesterlies. Upward Ekman velocities associated with divergent southern Ekman transport thus occur south of the northern boundary created by the Aleutian chain through the entire subarctic Pacific, where wind-stress curl (wE)>0(Gargett, 1991). Precipitation exceeds evaporation in the subarctic Pacific and this maintains a low average surface salinity in the Gulf of Alaska of 32.6 on a year-round basis. Moreover, stratification is seasonally reinforced by summer heating and runoff from eastern and northern coastal boundaries. Convection is restricted by a near-surface halocline between 100 and 150 m such that no deep water mass is formed today (Warren, 1983). Even despite frequent storms, cooling, and low river discharge during winter, mixing is restricted to relatively shallow depths (Luick et al., 1987; Figure 2.3). Low evaporation rates relative to the subarctic Atlantic are effected by lower subarctic Pacific temperatures, principally in summer, which are in turn produced by a 6 Sta t ion N Lat i tude Figure 2.1. Temperature isotherms from the Vertex Alaska transect. Note the ascent of the 4° isotherm from 950 m at 35°N to 120 m at 54°N (after Martin et al., 1989). 30 40 50 60 N Lat i tude 30 ,0 l • • . 1 o 30 40 SO 60 N Lat i tude Figure 2.2. Nitrate, phosphate, silicate, and temperature transects at 5 m depth vs. latitude. Upwelling of the coldest (11.74° C) and most nutrient-replete waters occurs at 54°N; [ P O 4 ] =0.93, [ N O 3 ] = 9.0, and [S1O3] = 25.1 pmol kg"1 (after Martin et al., 1989). (a) Salinity (%<>; 32.5 33 33.5 ° [ tg 1 50 £ 100 0) 150 200 (b) Temp (°C) 34 4 6 8 10 12 14 100 150 200 ~° February * Augus t Figure 2.3. Profiles of (a) salinity, (b) temperature, and (c) oxygen for th< upper 200m of the water column at Station P (50°N, 145°W); d) oxygei profile expanded to 4200 m depth. 8 relatively low influx of subtropical waters. This is due in part to differences in basin morphology, i.e. an analogue of the Norwegian Sea does not exist in the Pacific. In the Atlantic, vertical circulation driven by arctic cooling allows the poleward transport of subtropical water and thus, heat. Deep-water convection in the subarctic Pacific during glacial-interglacial transitions has been postulated (e.g., Dean, 1989), though most recent studies suggest that historical ventilation of the deep Pacific during the Pleistocene is unlikely and that stratification probably increased in glacial times, thereby reducing deep convection (e.g., Keigwin et al., 1992; Zahn et al., 1991a). Nevertheless, Zahn et al. (1991a) indicate that convection may have occurred to intermediate depths during glacials. The North Pacific is at present the terminus of global deep ocean circulation; deep Pacific waters are largely derived from the North Atlantic via the Southern Ocean and upon reaching the North Pacific, these waters are "old", significantly depleted in oxygen and replete in 2XD2 and nutrients (Kroopnick, 1985). Bottom waters are somewhat higher in oxygen, reflecting the advection of Antarctic Bottom Water (Figure 2.3). The deep North Pacific is known to comprise the most homogeneous water mass in the oceans. Properties such as temperature, salinity, and oxygen concentration show little variation throughout much of the deep subarctic Pacific, though deep water appears "oldest" in northernmost latitudes, consistent with circulation models. Indeed, the Alaska Gyre appears generally to maintain its expression to the ocean floor. Warren & Owens (1988) suggest that two nearly zonal deep currents stretch across the subarctic Pacific; the southern current flows northeastward from the Emperor Seamounts to the Gulf of Alaska where most of the jet bends north and returns westward forming a deep extension of the Alaska Current along the Aleutian Arc. 9 2.2 Dynamics of Primary Production The subarctic Pacific is one of a few high nutrient, low chlorophyll (HNLC) ocean regions, where substantial concentrations of the major nutrients nitrate, phosphate, and silicic acid are continuously present. Ecological understanding of this region remains unclear, and varying views are held about the factors that control primary production in such areas. The present consensus on processes controlling productivity in the subarctic east Pacific (SEP) is summarized briefly in this section in order to provide some rationale for past changes in new production at Site 887. Current estimates of primary production rates in the Alaska Gyre (Station P; 0-80 m) for spring (March-May), summer (June-August), autumn (September-November), and winter (December-February) are 415, 466, 366, and 283 mgC n r 2 d" 1, respectively (Wong et al., 1995b). Mean annual primary production, estimated at 140 g C n r 2 y r 1 (383 mgCm'^d" 1), is close to that of equatorial Pacific waters. Regenerated production appears to dominate the SEP system, with reduced nitrogen produced by excretion and remineralization being the prominent nitrogen source for primary producers. Based on a mean annual / -ratio of 0.1 to 0.4 at Station P, new production is 10 to 40% of total production1. Uptake of ammonium and urea occurs more rapidly than nitrate in the surface layer, though it is presently unclear whether high nitrate concentrations reflect a preference for ammonium or suppression of nitrate uptake (D. Varela, unpublished data; Wheeler & Kokkinakis, 1990). Although surface nutrient levels in the SEP are seasonally elevated, they remain below concentrations found in the subhalocline layers (> 100-200 /-ratio is defined as the ratio of new to new plus regenerated (ammonium and urea) nitrogen uptake (Eppley & Peterson, 1979). 10 m, nitrate ~45 \iM) due to a lack of deep winter mix ing . However, some mix ing of the surface layer w i t h waters between the seasonal thermocl ine at ~ 35 m a n d the ha loc l ine does occur i n late win te r d u r i n g February a n d M a r c h as thermal strat if icat ion breaks down. Thus, spr ing ni trate concentrat ions i n the mixed layer at Station P can reach 17 n M , though w i th substant ial i n t e rannua l va r i a t ion ; as uptake by p h y t o p l a n k t o n occurs f r o m A p r i l to October, ni t rate typ ica l ly falls to 6 M M (Mi l le r et al . , 1991b). Near the center of the Alaska Dome, major n u t r i e n t levels m a y r e m a i n h igher , owing to enhanced u p w e l l i n g . D u r i n g a September 1992 cruise , for example , m i x e d l aye r N O 3 a n d S 1 O 4 concent ra t ions were measured at 11 u M a n d 22 ^ M for Sta t ion A G (55°N, 145°W), and 8 »M and 16 j i M for Station P (50°N, 145°W; D . Vare la , unpub l i shed data). Seasonal s t ra t i f icat ion along w i t h the pe renn ia l presence of NO 3 , P O 4 , and Si04 i n the r e g i o n are p re requ i s i t e to s p r i n g b l o o m c o n d i t i o n s , yet substantial b looms as seen i n the subarctic At lan t i c have not been observed. The absence of seasonal va r ia t ion i n the phy top lank ton stock is we l l k n o w n i n the subarctic Pacific. This is i l lus t ra ted by cumulat ive data col lected at Station P since 1959 wh ich document year r o u n d c h l o r o p h y l l values of close to 0.3 mg C h i a m ~ 3 . In contrast, N o r t h At lan t ic phy top lank ton blooms i n excess of 1 mg C h i a m ~ 3 are observed y e a r l y even i n remote oceanic sites ( summary i n M i l l e r et al . , 1991). A n u m b e r of explana t ions have been presen ted to account for the balance w h i c h is c o n t i n u a l l y sus ta ined between p h y t o p l a n k t o n g rowth a n d loss i n the subarct ic east Pacific (SEP). These are based o n a suppos i t ion by H e i n r i c h (1962) w h i c h involves con t ro l of p h y t o p l a n k t o n growth a n d hence, l i m i t a t i o n of b looms by z o o p l a n k t o n g raz ing . A general s u m m a t i o n of the cur ren t ly accepted "mix ing a n d micrograzer" hypothesis fol lows. 11 Mesozooplankton (0.2-20 mm) are omnivorous and feed on both phytoplankton and microzooplankton (20-200 um; Frost, 1987). Recent studies have shown that mesozooplankton cannot graze efficiently on small (0.2-20 urn; pico and nano) plankton of the area which include Synechococcus spp., flagellates, dinoflagellates, crytptomonads, coccoids, and diatoms. Hence, mesozooplankton grazing is insufficient to maintain balance since 90% of plankton biomass is typically in cells < lOum in diameter. Alternatively, microheterotrophs (protozoans) have been cited as the primary control on phytoplankton stock in the subarctic Pacific (see summary in Miller, 1993). Microzooplankton, e.g. heterotrophic flagellates and ciliates, have been observed at Station P in large numbers; they can feed day and night and increase at up to 5 doublings d" 1 while phytoplankton are capable of about 2 doublings d"1 (e.g., Booth, 1988). Thus, their population growth in response to increased phytoplankton abundance can limit any increase in phytoplankton stock. However, since microheterotrophs are smaller than the larger phytoplankton species, their control on floral abundance depends on dominance of the phytoplankton by very small forms. Thus, a phytoplankton population of small-sized cells is requisite to maintaining a balanced system in the SEP where stocks fluctuate within a narrow range, controlled by unicellular grazers. Large phytoplankton are present, albeit only as a small component, and their growth seems to be matched by mesozooplankton grazing (Horner & Booth, 1990). Still, the enigma of which mechanism supports the consistent dominance of phytoplankton in the subarctic Pacific by flora of very small size remains unresolved. SEP waters remain stable year-round above the permanent halocline at 110 m (Section 2.1). However, since no such barrier to mixing exists in the North Atlantic, winter mixing occurs to depths of at least 250 m (Levitus, 1982). 12 Following the vernal North Atlantic bloom and subsequent nutrient depletion, winter mixing clears phytoplankton stocks from the surface layer in this basin. In contrast, since deeper Pacific waters which are devoid of phytoplankton do not inundate the euphotic zone, SEP phytoplankton and micrograzer stocks are not much reduced in the winter months. Consequently, the SEP phytoplankton-grazer relationship remains extant throughout the year (Miller et al., 1991a). 2.2.1 Iron Limitation In the marine environment, iron is required for the synthesis of chlorophyll and for the reduction of C02,SC>4, and NO3 as organic compounds are produced during photosynthesis. In spite of this importance as a micro nutrient, the analytical techniques required to measure open ocean Fe levels (< 1 nmolkg'l) have only recently been developed (Martin, 1988; Martin et al., 1989). In the Gulf of Alaska, dissolved Fe profiles show surface depletion and deep enrichment, characteristic of nutrient distributions (Figure 2.4). Studies indicate that low Fe levels in the subarctic Pacific are responsible for reduced phytoplankton growth rates in waters that are otherwise apparently nutrient-replete (Martin, 1988; Martin et al., 1989). For example, Coale(1991) states that "...dramatic increases in phytoplankton productivity, chlorophyll-a, and cell densities occurred after the addition of 0.89 nM Fe..."; a similar effect due to Cu addition is attributed to a negative effect on grazers. Thus Fe appears to play a crucial role in phytoplankton production in HNLC regions, though other trace 13 n m o l Fe kg 0.0 0.2 0.4 0.6 0.8 1.0 1 2 1.4 i . 1 . 1 . 1 . — T , 1—7 1 , ,— u m o l 0 2 kg "1 4.0 1 • 1 ' 1 • ' ' 1 ' " i 1 i 1 0 10 20 30 40 50 60 70 u m o l N 0 3 k g Figure 2.4. Dissolved Fe profile at Station P, shown with oxygen and nitrate distributions (after Martin et al., 1989). Sta t ion 30 40 50 60 N Lat i tude Figure 2.5. (a) A dissolved Fe section from the Vertex Alaska transect illustrating an offshore-inshore trend to higher Fe levels. Note also the general surface depletion and enrichment with depth. The maximum at T4 delineates westward flow from the California Margin (-1500 km to the east), while the minima at T6 represents eastward flow from the western Pacific (>8000km to the west). Fe-enrichment is also coincident with oxygen minima at T2-T6andT8. 14 metals such as Mn, Co, Ni, Cu, and Zn may also influence production and ecology in such areas (Bruland et al., 1991). Relative to offshore Fe-limited waters, Fe concentrations increase with proximity to the Alaskan continental margin to > 1.5 nM near the Alaskan shelf break (Figure 2.5). Such high dissolved Fe concentrations are associated with particulates rich in Fe; for example, particulate Fe values typically exceed 1.0 nmol kg"1 north of the Alaska Dome. Since Fe is introduced into the water column at continental margins, dissolved Fe enrichment generally decreases with distance offshore. Coincident with Fe enrichment, nitrate depletion occurs in near shore surface waters, supporting the contention that excess nutrients are maintained in the Gulf of Alaska and other upwelling regions since Fe is the limiting agent (Martin et al., 1989). To illustrate, in the center of the Alaska Dome (54°N) the surface water Fe concentration is about 0.07 nM Fe and there is 11 L1MNO3. With a phytoplankton C:Fe ratio of 10,000 to 100,000 this would be sufficient to produce up to 7 u,mol of C; assuming a Redfield ratio of 6.6 C:N, enough N would be present to produce up to 73 pmol of C. Hence, based on Fe concentration, the phytoplankton use only about 10% of the available NO3 (see Martin, 1992). Such observations would seem to indicate that Gulf of Alaska phytoplankton are not able to consume excess surface N and P, the use of which could substantially increase new production in the SEP. Martin's supposition is that open ocean waters are not capable of supporting substantial phytoplankton growth without Fe from sources other than waters mixing up into the euphotic zone. Sunda et al. (1991) challenge this hypothesis, suggesting that the coastal species used in Martin's experiments (Thalassiosira weissflogii) require higher Fe concentrations for growth than comparable oceanic species (e.g., Thalassiosira oceanica). In fact, the oceanic species was found to have extremely low iron requirements with 15 C:Fe ratios of 500,000:1. In a subsequent paper, Martin (1992) argues that Fe concentrations in upwelled waters are too low for maximal growth even for species with low Fe requirements. He maintains that below a threshold level (~ 0.3 - 0.5 nM), low diffusion rates may limit the ability of plants to concentrate the Fe required for maximum growth rates. Thus, "a very sharp cut-off exists between organisms that can grow maximally at oceanic Fe levels ... and those that must wait for periodic elevated Fe levels in order to grow". As a result, productivity rates increase as Fe becomes more available to species with higher Fe requirements. Small-sized phytoplankton, dominant in the Gulf of Alaska, appear to be growing as fast as temperature and illumination allow, and are limited in abundance by grazing, yet they are not severely Fe-limited (Miller et al., 1991b). Although the role of Fe limitation remains unclear, it may foster the scarcity of large cells with low surface-area to volume ratios. According to Morel et al. (1991), in order to double once a day at ambient [NH4+] = 0.3 uM, cells > ~ 7 um in size must supplement the available ammonium with nitrate. Since growth on NO3- requires additional cellular iron, larger cells (r > 10 um) should be Fe-limited owing to their larger size and increased need for NO3". These larger cells likely constitute a large component of the organic matter exiting the euphotic zone; thus, the supply of available Fe may limit NO3" utilization and hence, new production (Dugdale &Goering, 1967). In terms of bioavailability, the speciation of Fe plays a key role. Insoluble Fe(III) oxide minerals may be reduced via photoreduction in the atmosphere and in seawater to the soluble Fe(II) state. Fe is rapidly oxidized from the soluble +2 to insoluble +3 state; thus in the presence of oxygen, dissolved Fe(II) would be rapidly reoxidized (reoxidation rate = ~ 20 h _ 1 ) . Organic compounds enhance photoreduction in seawater by providing an electron source for photoreduction and by complexing Fe(II) in solution, 16 thereby slowing the reoxidation rate. However, it is probable that the main effect of seawater photoreduction is the formation of Fe(III)-oxide coatings on mineral surfaces, since Fe(II) oxidizes more quickly at Fe(III) oxide surfaces than in solution. These coatings are labile and more soluble than the original oxides; hence, they should generate increased dissolved Fe(III) concentrations, thereby enhancing Fe bioavailability to ligands in phytoplankton outer membranes (summary in Morel et al., 1991). 17 3. Background 3.1 Barium as a Productivity Tracer 3.1.1 Ba Distributions The dilemma of isolating productivity from dissolution signals is inherent in the use of productivity indicators such as the opal and carbonate content of marine sediments. Each of these proxies is recycled to some degree in the water column and in sediments such that the productivity signal is diminished. Si and alkalinity are removed in the upper water column via the formation of biogenic silica and carbonate, and are regenerated at depth. Although the mechanisms of barium cycling differ from those of Si and carbonate, recent mapping by GEOSECS and other programs indicates that all three covary in their distributions (Figure 3.1). As the more refractive chemical species, Ba preservation patterns may more accurately record a productivity signal in sediments. The association between Ba and productivity has been known for some time. Goldberg & Arrhenius (1958) linked high sedimentary barium contents with high productivity below the equatorial divergence in the Pacific and Indian Oceans and in the years since, a number of studies have affirmed this connection (refs. in Shimmield et al., 1994). Schmitz (1987), for example, used increases in sedimentary barium content to trace the northward shift of the Indian Plate during the Cenozoic as it passed beneath the equatorial divergence. In the pelagic realm, microcrystalline barite (BaS04; ~ 1 ^m) precipitates concurrently with the degradation of organic matter (Dehairs et al., 1980). The mechanism of barite precipitation in the water column remains 18 160 2 2 5 0 2 3 0 0 2 3 5 0 2 4 0 0 2 4 5 0 2 5 0 0 Alka l in i ty no rma l i zed to S - 3 5 par ts per 1 0 l 2 ( i j e q u i v k g - 1 ) Figure 3.1. Seawater Ba vs. alkalinity, normalized to a salinity of 35%o over 9 widely spaced GEOSECS sites. Ocean regions are indicated by the following symbols: circles = Atlantic, crosses = Pacific, diamonds = Indian (from [Lea, 1989#182]. Corg/Ba 0 40 80 120 160 200 0 -j ! 1 1 1 (_ Figure 3.2. C o rg/Ba in sediment trap samples vs. depth for three ocean areas (best fit curves). The rate of decrease with depth systematically increases from the Atlantic to the North Pacific (from Dymond et al., 1992). 19 equivocal, though a clue may be found in the similarities between dissolved Si and Ba distributions. This relationship may be the result of barite formation in microenvironments (diatom frustules, fecal material, etc.) enriched in opal and in sulphate from decaying organic matter during transport to the seafloor. Such environments, composed of aggregates of biogenic detritus, could become supersaturated with respect to barite, resulting in the formation of barite crystals. This scenario is evinced by water column profiles in which maximum particulate Ba concentrations occur just below the euphotic zone and decrease slightly with depth, suggesting that barite formation occurs after the suspended load sinks below the surface layer. Conversely, dissolved Ba is depleted in the euphotic zone and increases steadily with depth (Dehairs et al., 1980; Bishop, 1988). Thus enhanced new production may augment barite precipitation through the provision of substrate which is conducive to barite formation. Studies at mesopelagic depths (100 to 600m) suggest a link between new production and barite accumulation, with between 50 and 100% of total Ba in suspended material associated with organic matter degradation (e.g., Dehairs et al., 1992). Consequently, a decrease in C o r g flux with depth due to C o r g degradation should be associated with an increased Ba flux. This is supported by sediment trap data in intermediate and deep waters which indicate a strong correlation between Ba and particulate organic carbon fluxes. The preservation factor of Ba is the highest of biogenic materials; ~ 30% of Ba flux to the seafloor is preserved (Dymond, 1992). Dymond et al. (1992) observe systematic decreases in C o r g / B a ratios of settling particles as a function of water depth for a given site, suggesting that both a decrease in organic matter flux via degradation and an increase in Ba flux control the C 0 r g / B a decrease. Moreover, seawater Ba concentrations may 20 influence Ba uptake via the process of organic matter degradation, producing site to site changes in the decrease rate of the C 0 r g / B a ratio with depth. An empirical power function describes the C o r g :Ba relationship, using different constants to account for changes in dissolved barium between water masses (Figure 3.2). Based on this power function, Dymond et al. (1992) propose a quantitative algorithm for the assessment of variations in new production, which incorporates dissolved barium content, water depth, and particulate barium flux. Their estimates of palaeoproductivity in the northern California Current during the last 18 kyr show that new production was significantly lower during the last glacial maximum, consistent with foraminiferal Ba/Ca data which indicate that glacial deepwater Ba concentrations were lower in the equatorial Pacific and higher in the Atlantic relative to today (Lea & Boyle, 1990). Nonetheless, it should be noted that the processes reflected in regional California Current palaeodata could be completely independent of the very broad scale processes implied by Lea and Boyle. At present, barium displays a nutrient-like distribution, increasing in concentration with depth from surface waters (35 nM) and along the path of bottom water flow through the Atlantic (70 nM) and into the Pacific (150 nM) (Collier & Edmond, 1984). Thus for a given amount of organic matter degradation, barite formation may be higher in Ba-enriched Pacific waters relative to the Atlantic (Figure 3.2). If this is the case, then local productivity signals will be superimposed over inter-ocean nutrient and alkalinity variations. Thus, temporal variations in water mass chemistry should produce similar changes in sedimentary biogenic Ba contents within ocean basins. 21 3.1J? Alternate Ba Sources Though the sedimentary distribution of barium in pelagic sediments largely reflects "biogenic" barium flux, aluminosilicates contain 300 to 1000 ppm Ba and are an important Ba source in nearshore environments (Dymond, 1992). Since the preservation of aluminosilicate Ba is likely, this source could obscure the biogenic Ba signal. Based on the assumption that the Ba/Al ratio of terrigenous material remains constant, Dymond et al. (1992) apply a correction using a normative relationship which assumes a 'typical' Ba/Al value of 0.0075 for detrital aluminosilicates. This is claimed to account for aluminosilicate contributions up to 50% of total Ba: Excess (bio) Ba ppm = Total barium ppm - (Al wt%x 75 ppm/wt%) Nevertheless, the average Ba/Al ratio changes with sediment provenance, hence a knowledge of aluminosihcate sources is necessary in order to make an accurate correction. The excess value is independent of dilution if one assumes that Ba and Al are diluted equally by other phases such as opal, calcite, and quartz. Another approach, developed by Von Breymann et al. (1992), is based on the premise that thorium abundance reflects the lithogenic component of sediments. They used Ba/Th ratios in Peru Margin sediments to estimate excess barium relative to the aluminosilicate fraction and found Ba/Th ratios to be significantly higher in deep ocean sediments relative to the Peru Shelf, consistent with a lack of biogenic barium accumulation in shallow-water sediments. As yet no clear solution exists to account for terrigenous input and as a result, reconstruction of productivity changes based on barium data is limited to deep pelagic environments. 22 Xenophyophores are a group of barite-secreting benthic protozoans related to foraminifera. These organisms generate barite crystals 2-5 i^m in diameter which are indistinguishable from barite associated with suspended material (summary in Gingele & Dahmke, 1994). Intensified xenophyophore production in response to an increased supply of organic matter to the seafloor may augment the barium flux from the euphotic zone, thereby amplifying the surface productivity signal. However, the significance of this barium source is unknown, since the range and population density of xenophyophores remains unclear. Hydrothermal barite is another potential source of sedimentary barium since barite commonly precipitates from hydrothermal solutions. Nevertheless, hydrothermal Ba is likely to be of importance only near venting sites, as shown by results from the East Pacific Rise (Dymond, 1981). 3.1.3 Diagenesis Barite preservation in sediments is highly influenced by interstitial water chemistry. Oxic and suboxic sediment pore waters are generally saturated with respect to barite (Church & Wohlgemuth, 1972), hence postdepositional mobilization of barium in such environments is unlikely. However, barite becomes highly soluble as sulphate is depleted during sulphate reduction. Under such conditions, sulphate concentrations in pore waters can approach zero: BaS04 — Ba 2 + + SOj" 2CH20+ SOj- — H2S + 2HCO-As dissolution of barite proceeds, downward sulphate diffusion and upward Ba diffusion produce a Ba 'front' at the base of the sulphate reducing zone in which reprecipitation occurs. This diagenetic layer is characterized by an 23 enrichment of particulate barite. Overprinting of the productivity signal by this process has been documented in regions of anoxic diagenesis such as the Peru margin (Von Breymann et al., 1992). Pore water analysis may help to identify diagenetically-produced barite, though it is unclear if corrections can be made for diagenetic overprinting. The concentration of Ba in sediments could be affected by the presence of manganese oxyhydroxides which contain 1000 to 2000 ppm Ba (Dymond et al., 1984). As these are reduced and dissolve at the sediment redoxcline, Mn is remobilized from underlying reduced sediments and reprecipitated in the oxic zone. Froelich et al. (1979) suggest that the depth of the Mn spike is determined by the balance between upwardly diffusing Mn 2 + and02 diffusing downward, hence a solid Mn02 peak occurs near the base of the oxic layer. Changing carbon accumulation rates can produce redox changes such that variable Mn (and associated Ba) enrichments occur which may obscure the biogenic barium signal. Still, the Mn spike can migrate upward as sediment is deposited; in this case, as a steady state feature the Mn enrichment would not influence the downcore barium signal. 24 3.2 Biogenic Silica In this study, Si/Al ratios are used as a proxy for opal, or biogenic (hydrated) silica, in the absence of % opal data. The utility of sedimentary opal records in hindcasting productivity is controversial. Hence, rather than make a case for the quantitative interpretation of primary production on the basis of Si/Al measurements, a rationale for the qualitative application of these data is presented in the following paragraphs. The link between opal presence in sediments and productivity in overlying surface waters has been established for sometime. Bramlette (1946) suggested that the formation of diatomaceous sediments is associated with "...the drifting of this microplankton by currents from the open ocean into areas of deeper water...". Such deposits were later linked to regions of upwelling (Calvert, 1966), though difficulties in the quantitative analysis of sedimentary opal have limited the mapping of its distribution in marine sediments until recently. Unlike calcite, opal is not affected by differential dissolution with depth. Seawater is undersaturated with respect to opal and dissolution of tests occurs both as they sink and at the sediment-water interface. As a result, opal production rates in surface waters must transcend dissolution rates in seawater and in sediment pore waters in order for diatomaceous ooze to accumulate. After burial by additional opal and/or sediment, further dissolution and subsequent saturation of pore waters enhances preservation. Modeling studies on opal preservation in tropical and subtropical sediments by Archer et al. (1993) indicate that simple interpretation of productivity changes cannot be made on the basis of variations in sedimentary opal. Archer et al. (1993) suggest that: "present-day %Opaltot data... show little 25 or no systematic trend with the rain rate of opal" and "when burial represents only a small fraction of the opal rain...the rain of opal must be balanced by dissolution". However, where the burial flux represents a larger proportion of total biogenic silica production (i.e.'burial > 10% of opal rain), faster burial enhances preservation of the productivity signal in sediments through kinetic competition with dissolution. Thus, sediments which reflect high opal flux, as in the Southern Ocean, may be interpreted qualitatively in terms of opal productivity. Though opal data ostensibly provide useful palaeoceanographic information, the relationship between productivity and sedimentary opal could be influenced by the availability of dissolved silica due to shifts in the global distribution of dissolved silicate (Si(OH)4) brought about by changes in circulation. This is exemplified in the contrast between the equatorial zones of the Atlantic and Pacific; at a given level of productivity, opaline sediments are more enriched in the Pacific than in the Atlantic due to weaker silicification of Atlantic frustules. Differences in the concentrations of Si(OH)4 and PO43" between the two basins are responsible for this contrast: sub-surface waters in the Atlantic are more undersaturated relative to phosphate than those at equivalent depths in the Pacific (see summary in Berger & Herguera, 1992). As a result, the Si/P ratio in waters that well up to the surface in the Atlantic, being lower than that in the Pacific, supports growth of relatively weakly silicified diatoms. Only a small fraction of biogenic silica survives the journey from surface waters to burial in sediments, ranging from 10 to 30% in the Southern Ocean. Still, patterns of opal deposition in the Southern Ocean over the past 450 ka suggest that dissolution effects in seawater and in sediment pore waters are not the primary control on sedimentary opal presence (Charles et al., 1991). 26 The fundamental control on the degree of abundance in sediments is the flux of opal to the seafloor, itself governed by variations in the productivity of overlying surface waters. This is not to say that opal abundance provides an exact "mass flux" or quantitative record of palaeoproductivity, but simply that changes in opal content echo the processes in surface waters that control opal production. Production rates sufficient to support opal accumulation in sediments are at present sustained only in areas of strong upwelling; e.g., equatorial productivity zones, off southwest Africa, and the Kuroshio-Oyashio confluence in the northwest Pacific (Leinen et al., 1986). During the last glacial period, dissolved silicate concentrations in Pacific intermediate waters may have been reduced, akin to modern Atlantic conditions. If so, it is conceivable that weak silicification of tests in the glacial Pacific caused by a reduction in sub-surface Si(OH)4 produced a decrease in opal flux to the seafloor, thereby decreasing the presence of opal in glacial sediments. Such 'Atlantic type' behavior in regions of the glacial Pacific has been noted in carbonate records (Section 3.3). Nevertheless, Haug et al. (1995) have noted a shift in the western subarctic Pacific toward enhanced CaCC>3 accumulation during interglacial periods of the Brunhes Chron. This suggests that historical oceanographic conditions in regions of the subarctic Pacific, including dissolved silicate concentrations, may have been decoupled from those in the lower latitude Pacific Ocean. In such a case, a shallower lysocline and associated nutricline may have sustained high upwelled Si(OH)4 concentrations in the subarctic Pacific during glacial times. For the purposes of this study it is assumed that the opal dissolution rate is constant and represents a consistent proportion of the opal burial flux. Hence, the fraction of opal flux which is preserved and buried remains temporally constant. Though this is no doubt a simplistic view, the 27 undersaturation of opal in seawater seen globally implies that circulation changes do not markedly affect the dissolution rate. Thus, it seems reasonable to assume that, unlike CaC03, no systematic changes have occurred in opal dissolution over glacial-interglacial timescales. This is in agreement with a study by Shemesh et al. (1989) who showed that the preservation of diatom assemblages in the Pacific and Atlantic sectors of the Southern Ocean between Last Glacial Maximum and Holocene times was similar. 28 3.3 Carbonate The distribution of carbonate in marine sediments is largely controlled by the balance between the production of biogenic CaCC>3 produced in the upper ocean and dissolution at depth. Of particular importance is the increase in dissolution with water depth as a result of increased calcite solubility due to pressure and temperature changes (Hawley & Pytkowicz, 1969). As a result, carbonate preservation is generally enhanced in shallow sediments relative to deeper ones. Quaternary changes in temporal and regional preservation patterns have been attributed to changes in productivity, varying ratios of calcareous to opaline material, dilution by inorganic sediment, and variable dissolution caused by fluctuations in lysocline and CCD depths in response to changing deep water chemistry and circulation (refs. in Snoeckx & Rea, 1994). Debate centers around whether variations in the presence of carbonate reflect enhanced carbonate production with constant bottom water "corrosiveness" or changing bottom water chemistry with constant production. 3.3.1 Dissolution vs. Productivity The fraction of calcite production which is buried in deep-sea sediments depends on the area of seafloor that is above the calcite saturation depth or sedimentary lysocline, defined as the water depth where: [C032"]jj] situ =[C032"]saturation-Calcite dissolution occurs within a transition zone of variable thickness between the lysocline and the CCD (0% sedimentary CaC03; see summary in Farrell & Prell, 1991). An imbalance between sources such as continental weathering and biogenic calcite input and sinks (deep and shallow water 29 deposition) results in a shift in [CO3"] or 'CaC03 compensation', according to the reaction: C0 2 + (Ca)CO^~ + H 2 0 2HCO" + (Ca 2 +) The subsequent changes in lysocline depth are manifest as carbonate preservation/dissolution events in deep-sea cores. Evidence that carbonate flux to the seafloor increases with productivity was first recognized by Arrhenius (1952) who noted a depression in the CCD below the regional level in the equatorial Pacific region. He argued for higher glacial productivity, based on the observation that the CCD was further depressed during this time. Since then, a number of studies have shown that although productivity controls the carbonate input to the deep sea, carbonate flux is nonlinear with productivity and quickly flattens out at the high end (Berger & Herguera, 1992). This non-linearity may be the result of a change in the C0rg/CaC03 ratio which generally increases with productivity. Specifically, CaC03 dissolution is enhanced via the release of CO 2 and a concurrent reduction in the pH of interstitial waters, produced by the oxidation of an increased supply of organic matter (e.g., Berger & Keir, 1984). Thus, the carbonate burial flux is likely not a reliable quantitative productivity indicator, except within limited conditions and productivity ranges (Berger & Herguera, 1992). However, in focusing on the amplitude of variation rather than on absolute values, changes in carbonate abundance can be interpreted as a function of carbonate production. In order to avoid dilution effects and to reflect more accurately the flux to the seafloor, carbonate input can be expressed as a Mass Accumulation Rate (MAR), calculated as the product of sedimentation rate, %CaC03, and dry bulk density (p). 30 Carbonate Cycles in the Pacific One assumption common to both proponents of productivity and dissolution is that carbonate abundance provides a climostratigraphic record. Increased carbonate abundances in glacial relative to interglacial periods have been observed in a number of cores from the central equatorial and north Pacific (Arrhenius, 1952; Hays et al., 1969; Karlin et al., 1992). Nine carbonate maxima occur during the Brunhes and comparison with oxygen isotope records indicates that heavier 6 l s O events are generally associated with enhanced preservation (e.g., Farrell & Prell, 1989; Figure 3.3), though this classic carbonate pattern is not well defined in the eastern equatorial Pacific (Snoeckx & Rea, 1994). Carbonate preservation maxima and minima lag respective glacial and interglacial transitions by an observed range of 3 to 20 kyr, attributed to the response time of carbonate ion in the ocean to changes in the material balance between input and loss of CO3 2 " (Broecker & Peng, 1982). During the mid-Brunhes (200-450 kyr), both Atlantic and Pacific records show an increase in CaC03 dissolution and large glacial-interglacial amplitudes relative to the early and late Brunhes. Such a long-term change may indicate global changes in the ocean carbonate reservoir (Crowley, 1985). In the central equatorial Pacific, Lower Brunhes cycles are characterized by high CaC03 values and relatively low amplitude variation; conversely, lower Ca303 values and higher glacial-interglacial amplitudes are seen in younger cycles (Snoeckx & Rea, 1994). Throughout most of the Late Quaternary, glacial/interglacial carbonate preservation cycles appear to have remained largely out of phase between the North Pacific and the North Atlantic, possibly the result of changes in the 31 formation rate of NADW (e.g., Crowley, 1985; Farrell & Prell, 1991). As thermohaline flow progresses south from the North Atlantic to the circumpolar current and subsequently north into the Pacific, the oxidation of a continual influx of organic matter causes a gradual increase in 2GZ>2- Thus, deep waters acquire a lower carbonate concentration with age and become increasingly corrosive to CaCC>3. Since the North Pacific is at present the terminus of global deep ocean circulation, carbonate preservation is reduced there relative to the North Atlantic. At present, deepwater Atlantic and Pacific (D3 2" contents are about 108 and 92 imiol/kg, respectively. During glacial times, a reduction in the supply of nutrient-depleted North Atlantic Deep Water (NADW) could produce a rise in the Atlantic lysocline, making the deep Atlantic more corrosive due to increased 2CD2 and thus lower carbonate ion concentrations. A similar pattern of glacial dissolution is also found in Southern Ocean sediments, consistent with a reduction in NADW and high nutrient levels in deep water (Howard & Prell, 1994). A concurrent 400 to 800 m deepening of the lysocline in the Pacific would have made the deep Pacific less corrosive (Farrell & Prell, 1991), reflecting either balancing of reduced NADW flow by increased input from a southern source or the formation of less corrosive North Pacific deep water masses (Duplessy & Shackleton, 1985). However, increased productivity rather than enhanced preservation as the dominant control has not been ruled out (Pedersen, 1983), and there is as yet no general agreement on the causes of Pacific carbonate cycles in the Quaternary. 32 uojjniossjp 6U|SB9JOU! 3.3.2 North Pacific Carbonate Records Deviations from the Pacific-type carbonate patterns are apparent at a number of sites. Hovan et al. (1991) found shifts in CaCC»3 accumulation patterns in northwest Pacific cores, switching from the Atlantic-type pattern of enhanced accumulation during interglacials in the Lower Brunhes to the Pacific pattern of greater accumulation during glacial periods following 350 kyr BP. In contrast, other northeast Pacific cores generally follow the central equatorial Pacific pattern (Karlin et al., 1992; Zahn et al., 1991b). Carbonate profiles at two subarctic sites in the Emperor Seamount Chain investigated by Keigwin et al. (1992) and Haug et al. (1995) at ~ 53° and 50°N respectively also show deviations from 'classic' Pacific CaCC>3 stratigraphy. Keigwin et al. (1992) observed lower carbonate accumulation and a slight increase in CDrg and opal during the last glaciation relative to the present. They suggested that enhanced carbonate, opal, and C o r g fluxes on the last glacial-interglacial transition represent a temporary increase in productivity in the northwest Pacific. This productivity spike appears to be coincident with a sea surface salinity drop of 1, possibly the result of the northward withdrawal of a seasonally fluctuating ice margin. Russian authors have also noted productivity maxima associated with deglaciation at similar open ocean latitudes in the Pacific and in the Okhotsk Sea (refs. in Keigwin et al., 1992). Haug et al. (1995) report an increased presence of carbonate in Quaternary interglacials which they interpret as enhanced preservation, suggesting an "Atlantic-type" pattern. They attribute these changes to fluctuating deepwater ventilation in the Northwest Pacific at the terminus of the thermohaline conveyor belt, suggesting that increased upwelling of nutrient-enriched deepwater produced by intensified interglacial NADW 34 formation led to enhanced surface productivity. Such unusually high carbonate production and accumulation may have exceeded dissolution during early-peak interglacial times, resulting in CaC03 maxima. This hypothesis is supported by coincident variations in opal and CDrg records which suggest enhanced interglacial productivity. However, carbonate preservation in the upper portion of the record appears discontinuous with significant spikes near terminations 1 and 2. In addition, many inconsistencies exist between the three profiles; for example, though CaCC>3 maxima are generally parallel with opal, a number of opal maxima occur during periods of low carbonate accumulation (Figure 3.4). If true, this scenario has important implications for palaeoclimates. Reduced upwelling in the northwest Pacific could, for example, produce a decrease in the outgassing of CO 2 and subsequently a lowering of atmospheric POO2. Furthermore, a decrease in nutrient supply could explain, at least in part, the glacial Cd-depletion (20-30% relative to today) observed in northwest Pacific foraminifera tests (Boyle, 1992), thereby weakening the hypothesis of a glacial North Pacific source of deep water2. Nonetheless, other factors may have been responsible for changes recorded in the sediments cored at Site 882. For example, interglacial productivity could have increased due to: (1) higher sea surface temperature, conducive to greater carbonate production, (2) changes in sea ice cover, and (3) increased insolation during deglaciation (Keigwin et al., 1992). 2 A lack of good core-top data, partially due to enhanced dissolution, and a scarcity of cores limits the interpretation of Boyle's North Pacific Cd record. Furthermore, little difference between glacial and modern nutrient contents in deep northwest Pacific waters is suggested by SA 3C data. 35 Brunhes/Matuyama Jaram. A g e (Ma) Figure 3.4. Mass accumulation rates and weight percentages of biogenic opal, TOC, andCaCC>3 over the last 1 m.y. at ODP Site 882. Glacial Stages are extrapolated from the benthic oxygen isotope record of Tiedemann et al. (1994) from Site 659 (from Haug et al., 1995). 36 3.4 Carbon Isotopes Complex equilibria between dissolved carbon species, atmospheric CO2, CaG03, organic matter, and seawater serve to maintain most of the carbon in the deep ocean in the dissolved form (2(D2; Kroopnick, 1985). The synthesis of biogenic particles in the surface layer and their subsequent sinking and degradation as well as vertical and horizontal mixing, influence the distribution not only of 2CD2 but also oxygen and nutrients in the sea. Because oxygen and nutrient concentrations are closely linked to the carbon cycle, temporal and geographic variations in ocean chemistry and productivity can be traced through the use of carbon isotope ratios. Two stable isotopes, 1 2 C and 1 3 C , make up about 98.89 and 1.11 % of the earth's carbon respectively (Faure, 1986). The 1 3 C / 1 2 C ratio is normally expressed in terms of per mil (%o) deviation (S 1 3C) from the PDB belemnite standard, according to the following equation: 6 1 3 C values in calcareous marine shells reflect principally the 5 1 3 C of dissolvedCO2 in ambient seawater (6 1 3 C2C02); t n u s changes in the 6 1 3 C of fossil shells can provide a record of past changes in ocean chemistry. However, a number of factors including air-sea exchange, the addition of CO2 via the decomposition of both organic and inorganic matter, and circulation patterns influence the oceanic distribution of 513Q>;c02 a n d serve to complicate this record. 6 1 3C = 10001 \(13c/12c\ 37 Because 1 2 C is fixed preferentially over 1 3 C during photosynthesis, phytoplankton are depleted in 1 3 C by about 20%o relative to seawater ZCD2; marine phytoplankton 6 1 3Co rg values typically range between -20 and -30%o (Sackett et al., 1965). Carbon isotopes are fractionated to a much lesser extent during the formation of CaC03; the difference between CaCC>3 carbon and dissolved carbon in seawater is only about l%o. Hence, the primary control on 6 1 3 C is the photosynthesis-respiration cycle, while carbonate formation/dissolution and associated alkalinity changes are of secondary importance (Broecker & Peng, 1982). As organic matter is oxidized after export from the euphotic zone, 1 2 C is preferentially transferred into deeper water, resulting in an enrichment of about2%o in surface 1 3 C relative to deeper waters (Broecker, 1982). Thus, 613C2C02 values are largely determined by the input and subsequent degradation of organic matter as it sinks through the water column. This is evinced by the covariance between S 1 3 C and nutrient distributions, for example, the inverse correlation seen globally between S 1 3 C and phosphorus. Accordingly, S 1 3 C is used as a proxy for nutrient concentrations, as summarized by Broecker and Peng (1982). The use of S 1 3 C as an indicator of carbon system properties has important implications for understanding atmospheric CO 2 variability. Two general explanations have been proposed for changes in atmospheric CO2 concentrations. One involves the transport of photosynthetically-produced carbon from the surface to the deep ocean via the 'biological pump', whereby a decrease in 2 © 2 in surface waters produced by an increase in the efficiency of the biological pump could cause a drawdown in atmospheric CO2. Glacial to interglacial CO 2 shifts recorded in ice cores are characterized by an increase of 0.4 to 0.7%o in the 6 1 3 C of atmospheric CO 2 concurrent with 38 an increase of ~0.35%c in the deep ocean. These "heavy" 6 1 3 C values may intimate a decrease in high latitude surface water productivity and the extraction of organic carbon from the ocean via growth of the terrestrial biosphere and storage of carbon in soils. Low atmospheric CO 2 concentrations during the last glacial may be due to increased nutrient supply and/or productivity in the low-latitude surface ocean (summary in Marino et al., 1992), although recent carbon and nitrogen measurements argue against this (Farrell et al., 1995; Pedersen et al., 1991). Another proposal argues for lower pC02 in the glacial ocean produced by higher surface alkalinity (CO32"). Alkalinity changes caused by variable carbonate dissolution due to the rearrangement of water masses and/or changes in CaC03 sedimentation may have produced a glacial deepening of the lysocline in the Pacific and a concurrent modest rise in the Atlantic (see Section 3.3). Although these processes are not thought to be apparent the carbon isotope record, there is evidence that atmospheric CO2 variations may have been produced by a combination of changes in biological pump strength and oceanic alkalinity (Marino et al., 1992; Leuenberger et al., 1992). Preliminary work by Spero et al. (1995) indicates that there may indeed be a link between [CO32"] and the planktonic foraminiferid G. bulloides 6 1 3 C. They suggest that during the last glacial maximum, an increase in the carbonate ion concentration of about 80 fxmol/kg would have produced a 6 1 3 C decrease of 0.3 to 0.4%o. This coincides with an increase in surface ocean pH of 0.15 units along with 30% lower atmospheric C0 2 . Applying a "pH correction" to planktonic glacial S 1 3 C values could increase benthic-planktonic A S 1 3 C differences thus enhancing the perceived strength of the biological pump. 39 3.4.1 Some Constraints on the 613C Record 3.4.1.1 Gas Exchange In seawater, 5 1 3C2C02 varies from ~0to2%0, averaging 9%o higher than atmospheric CO2 (Kroopnick, 1985). Because the dissolution of C 0 2 in seawater requires the kinetically-controlled exchange of carbon atoms among HCO3" andCC>32" ions in addition to C 0 2 , the replacement time for carbon isotopes is about 15 years. Although near-equilibrium between atmospheric and surface water 2XD2 may be attained, surface waters are in general replaced too quickly in the sea for isotopic equilibrium to occur. Nevertheless, air-sea exchange can influence the relationship between 6 1 3 C and nutrients, leading to some degree of decoupling (Broecker & Maier-Reimer, 1992). Isotopic fractionation due to air-sea exchange increases with lower temperatures, ranging from 10.6%o at 0°C to 7.6%o at 30°C (Broecker & Peng, 1982). This occurs as a result of enhanced gas exchange at low temperatures. On the basis of nutrient-normalized 6 1 3 C values, Charles (1993) suggested that gas exchange effects an increase of up to l%o in Antarctic surface water 130>:co2 relative to low latitude surface waters. The sinking of these waters yields a slight "thermodynamic enrichment" in deeper Antarctic water masses. Conversely, the thermodynamic effect creates a relative depletion in NADW 1 3 C , presumably since warmer temperatures prevail during air-sea exchange prior to cooling and sinking of this water mass. Deep-water thermodynamic effects are not likely to overprint the glacial-interglacial 6 1 3 C shift since, for example, a change of 0.5 %o calls for a 5°C change in equilibration temperature which requires profound changes in deep water formation (Broecker & Maier-Reimer, 1992). Moreover, such effects 40 are not at present readily identified in the North Pacific, since deep water is not formed there. 3.4.1.2 Diagenesis, Vital Effects Two processes work in sediments to alter the 6 1 3 C signal: 1) the decay of organic matter adds isotopically light CO 2 into the pore water and 2) the dissolution of CaC03 releases slightly heavy CO2. Where contributions from each source are equal, organic carbon decay dominates and the net effect is that pore waters a few centimetres below the seafloor yield 8 1 3 C values up to 1.5%o lighter than the overlying bottom water (McCorkle et al., 1985). This diagenetic influence is evident in a number of studies summarized by McCorkle et al. (1985) which show that porewaters with varying 8 1 3 C values can be found beneath the same water mass, with higher carbon rain rates producing lighter 6 1 3 c values. Thus, the 6 1 3 C values of infaunal foraminiferal species including Uvigerina and Gyroidinoides may be influenced by porewater 6 1 3 C which is sensitive to the settling flux of carbon and reflects the in-situ degradation of isotopically light organic matter (Zahn et al., 1986). For this reason, epifaunal species such as Cibicides are more likely to provide an accurate estimation of bottom water 6 1 3 C (Zahn et al., 1986; McCorkle et al., 1990). Variations in the size of foraminiferid tests may also affect the 8 1 3 C signal, since the size of a shell is proportional to its carbon isotope ratio. Small (juvenile) shell chambers may be reduced by > 2%o relative to large (mature) shells due to the incorporation of carbon from respiration CO2 into juvenile shell material and/or kinetic fractionation during calcification (Oppo & Fairbanks, 1989; Spero, 1995). Associated correction factors may be larger than the amplitude of the downcore signal; however, the selection of a number of 41 shells within a specific size range serves to average out and thus minimize this effect (Curry & Crowley, 1987). Other factors such as symbiont photosynthesis can produce species-specific variations in 6 1 3 C values. Nonetheless, it has been demonstrated that in some species such vital effects do not significantly affect the relationship between foraminiferal and seawater 513C(Spero, 1992). In the interpretation of 5 1 3 C records it is important to understand the depth habitat of the signal producers. For example, deeper thermocline dwellers (e.g., N. dutertrei) typically record lower 6 i 3 Crjo2 than "true" surface dwellers such as G. bulloides (Kroon, 1988). Such attenuation of the 6 1 3Crjo2 signal could significantly affect reconstructions of surface CO2 which assume atmospheric equilibration, as noted by Jasper & Hayes (1993). 42 3.4.2 Carbon Isotopes in Organic Matter 3.4.2.1 Organic Carbon in Marine Sediments The rate of new (export) production is enhanced in areas where vertical mixing provides a nutrient source. As a result, upwelling regions are more likely to support high new production (and high export flux) than areas of vertical stability (Dugdale & Wilkerson, 1992). Because the settling flux of organic matter decreases with depth via degradation (Suess, 1980), at a given production rate more organic carbon arrives at the seafloor in continental marginal areas than in the open ocean. Furthermore, high rates of sedimentation in such settings increase carbon burial rates and enhance carbon preservation (Muller & Suess, 1979; Hedges & Keil, 1995).. Hence, the flux of carbon arriving at the seafloor is controlled regionally by the depth of the water column. However, variations in the accumulation of carbon in sediments are fundamentally determined by changing rates of new production. Thus, down-core increases in C o r g appear to reflect enhanced palaeo-productivity in the surface ocean (Calvert, 1987; Pedersen & Calvert, 1990). On the supposition that sedimentary C o r g content is controlled by variations in marine primary productivity, down-core isotopic variations in organic matter should reflect isotopic changes in plankton communities. Consequently, the sedimentary 6 1 3Co rg profile provides a proxy for plankton 6 1 3 C which, as described in the following section, varies in response to surface water [CCtyaq)] (Popp et al., 1989; Rau et al., 1991). 43 3.4.2.2 613Corg and C 0 2 Sackett et al. (1965) first noted a correlation between temperature and 6 1 3 C in plankton, and Degens et al. (1968) later suggested that [C02(aq)] concentrations actually control the observed 6 1 3 C: temperature relationship. A number of studies have since shown that 6 1 3 C o r g values reflect the degree to which photosynthetic carbon isotope fractionation varies in response to G32(aq) fluctuations (e.g., Rau et al., 1991; Jasper & Hayes, 1990). © 2 solubility increases with lower temperatures in accordance with Henry's Law. Based on this relationship, Rau et al. (1992) calculated that up to 2.5 times more C02(aq) should be present in colder high-latitude waters relative to warmer equatorial waters. They also found that the latitudinal change in [C02(aq)] corresponds to a range of 12 %o in plankton 6 1 3 C , with lower values in cold, high [C02(aq)] waters than in warmer, low [CC>2(aq)] waters. A later model based on a compilation of data by Rau (1994) indicates that ~ 90% of global bulk plankton 6 1 3 C variation is due to a negative linear response to [CC>2(aq)]-[C°2(aq)]±2.0>M = " V ? ' 9 ~ 1 A I x Rau suggests that the 80 ^atm glacial-interglacial change in pC02( atm) documented in ice cores corresponds to a 1 to 2 %o decrease in plankton 6 1 3C, equivalent to a 2 to 3 raM higher surface ocean [CC»2(aq)]- Thus, given 6 1 3 Corg and surface ocean temperature estimates, calculation of [C0 2( aq)] and PCO2 is possible on the basis of sedimentary data. This approach has been applied to Quaternary sediments in a variety of settings, including the Gulf of Mexico (Jasper & Hayes, 1990), the Panama Basin (Pedersen et al., 1991), and the Mediterranean (Fontugne & Calvert, 1992). However, changes through time in the 6 1 3 C of surface water Z(D2 can significantly affect 6 1 3 Corg values and a 44 correction must be applied; this involves subtraction of deviations in the down-core foraminiferal 6 1 3 C record relative to core top values (e.g., (Pedersen et al., 1991; Fontugne & Calvert, 1992). 3.4.2.3 Caveats The 6 1 3 C signal could be altered by the input of isotopically light terrigenous organic matter; low-latitude end member values are given by Jasper & Gagosian (1989) at 6 1 3Co rg terrigenous = -26.6%o and marine = -20.6%o. A measure of terrestrial input is provided by C:N ratios which are commonly 6 to 9 in sedimentary marine organic matter and 20 to 200 in terrestrial organic matter (Emerson & Hedges, 1988; Hedges & Parker, 1976). Sedimentary C/N ratios are higher than plankton values of 6 to 8 (Redfield et al., 1963) due to the preferential loss of nitrogen over carbon during diagenesis (Rosenfeld, 1981). Thus, concurrent increases in C/N values with lighter 6 1 3 C values would imply that the presence of terrestrial organic matter has influenced the 61 3Co rg results. Isotopic fractionation during diagenesis may also influence the 613Corg record, though it appears that a considerable fraction of sedimentary organic matter can be oxidized without fractionation (McArtliur et al., 1992). Fontugne & Calvert (1992) point out that changes of up to -2.8 %o due to diagenetic enrichment of 1 2 C are possible, though diagenetic trends vary with depositional environment. Such changes should be manifest in the sedimentary record as systematic differences between sedimentary 6 1 3Cor gand plankton 61 3Co rg. 45 Although the negative correlation between plankton 6 1 3 Corg and [G02(aq)i is well established, metabolic effects may serve to complicate the relationship, as described in the following section. 46 3.4.2.4 Metabolic Effects Expressed as ep, overall isotopic discrimination during carbon fixation is represented by the subtraction of phytoplankton 6l3Corg from inorganic source carbon, 6 1 3 Cco2 (Francois et al., 1994): E p = 6 1 3 Cco2 - 61 3Cpoc Variability in photosynthetic fractionation is to some extent contingent on the mode of transport of inorganic carbon through the cell membrane. Carbon transport is facilitated by either passive diffusion or by active uptake of CO2 or HCO3'. Francois et al. (1994) suggest that phytoplankton 6 1 3 C varies inversely with [CO 2(aq)] only if carbon transport across the cell membrane is driven by passive diffusion and is not actively transported. In the case of active transport, cells are induced to concentrate inorganic carbon where diffusion of molecular CO2 is limited, as in regions of high productivity. Hence, ep does not depend directly on [C02(aq)L The 1 3 C which accumulates in the cell in response to the preferential fixation of 1 2 C results in reduced 13C-discrimination, the net result being that fixed-carbon 6 i 3 C approaches that of the source carbon. Other considerations include rates of carbon fixation and intracellular leakage, as well as the speciation of the carbon taken up. Where passive diffusion controls uptake, ep is dependent on [C02(aq)J-Specifically, an e p decrease concurrent with a 8 1 3Co rg increase will occur with lower [C02(aq)] since any C02(aq) entering the cell will likely be fixed. In contrast, increased [C02(aq)] would allow greater expression of isotopic discrimination associated with enzyme-mediated carbon fixation in the cell. However, other factors such as high growth rate, larger cell size, reduced cell wall permeability, and variations in the degree of fractionation can produce a similar effect. 47 Francois et al. (1994) predict that ep should decrease in cells of large size and high growth rate. This is supported by culture studies which suggest that isotopic fractionation declines with increasing growth rates (Fry & Wainright, 1991). Moreover, large and fast-growing diatoms were shown by Fry & Wainright (1991) to be 1 3C-enriched relative to smaller phytoplankton in the North Atlantic (Georges Bank); 13C-replete diatoms were linked most closely with the latter stages of blooms and with periods of rapid growth. For example, in August large (40-60 diameter) Coscinodiscus with values of -15.5 %o were common in net samples, though they did not dominate 5 1 3 C P O M which ranged from -21.9 to -21.2. The same study indicates that in high-productivity areas where diatoms are important, heavy S 1 3 C P O M values of -21.5 to -18.5 are common. For example, off the Washington coast (~ 48°N), high-productivity areas (> 1 ng l " 1 chlorophyll) yielded S 1 3 C P O M values of ~ -20 to -17, whereas values for oligotrophic areas ranged from ~ -26 to -20. Thus, in marine ecosystems diatoms provide a potentially large source carbon replete in 1 3 C . The influence of 1 3 C-rich diatoms must be carefully considered when interpretations of historical oceanographic processes are based on the sedimentary 6 1 3 Co r g record. Certainly, changes in diatom abundance may influence phytoplankton 613Corg independently of [CC»2(aq)]-48 3.5 Oxygen Isotopes In the years since Emiliani's pioneering work (e.g., Emiliani, 1955), the oxygen isotope record from foraminiferal carbonate in deep sea sediments has been used extensively as a stratigraphic tool. Records are available from most oceanic regions, and the relationships between 6 1 8 0, temperature, and global ice volume have become a major focus of palaeoclimatic research. The relative abundances of 1 6 Oand 1 8 0 are 99.63 and 0.20 % respectively (Faure, 1986), though the 1 8 0 / 1 6 0 ratio of about 500 varies over a few percent between seawater, foraminiferal calcite, and ice-cap snow. This ratio is expressed as 6 1 8 0, representing a deviation from PDB or SMOW standards, analogous to 6l 3 C notation (Section 3.4). Average seawater 8 1 80(6W) values increase with global ice volume due to the accumulation of 160-replete snow. This is effected by thermodynamic fractionation; since higher energy is required to maintain the heavier isotope in the gas phase, H 2 1 6 0 evaporates preferentially from the oceans relative to Et21 80. Thus, the evaporation of seawater produces vapor that is depleted in 1 8 0 . As the vapor is transported in the form of clouds poleward, declining temperatures cause progressive condensation and the remaining water vapor becomes progressively enriched in 1 6 0 . As a result, modern Greenland and Antarctic ice caps are, respectively, about 3 and 4.5 % (i.e., 6 1 8 0 = -35 and-45 %o) lighter isotopically than seawater. By supposition, the glacial snow which formed much larger ice caps was isotopically depleted relative to modern snow. Foraminiferal 6 1 8 0 is a function of ambient seawater 6 1 8 0 which is influenced not only by continental ice volume, but also by local temperature and salinity changes. Along with changes in the 6 1 8 0 of surface waters, the preferential loss of 1 6 0 due to evaporation increases surface salinity, resulting 49 in a positive correlation between salinity and 6 1 8 0 . Thus, regions of high evaporation/precipitation ratios, and therefore high salinity, will be enriched in 1 8 0 . At high latitudes, seawater 6 1 8 0 and salinity are also influenced by meltwater derived from ice and snow with low 6 1 8 0 values. This is a linear relationship, estimated by Zahn et al. (1991a) for the North Pacific (40°-60°N) as: 6 W = 0 .405xS -14.014 The influence on 6 W of salinity variations resulting from sea-ice formation can be effectively ignored since the fractionation factor for ice in equilibrium with water is very small (1.002; Faure, 1986). The fractionation between oxygen isotopes in water and precipitated carbonate is strongly influenced by temperature; an increase in carbonate 6 1 8 0 of about 0.23%o occurs per degree (°C) of cooling. The specific temperature dependence of calcite 6 1 8 0(6 C ) is given by Craig & Gordon (1965): 16.9-4.2(s c - S w ) + 0.13(sc - 5 W ) Regional 1 8 0 variations due to temperature and salinity changes are primarily confined to surface waters; hence the deep ocean 1 8 0 distribution is more homogeneous. Consequently, benthic foraminiferal 6 1 8 0 (particularly in the Pacific) largely reflects changes in continental ice Volume, though the amplitude of change over glacial-interglacial cycles is lower in the Pacific than in the Atlantic Ocean due to enhanced glacial cooling in the deep Atlantic (Labeyrie et al., 1987). Correction of the planktonic 6 1 8 0 signal for ice volume effects thus allows an estimation to be made of regional hydrographic influences on the surface-water isotope signal. As with 6 1 3 C, preliminary work by Spero et al. (1995) indicates that G. bulloides 6 1 8 0 may be influenced by changes in seawater pH, with higher ambient pH producing lower 6 1 8 0 values. Assuming a universal dependence of 50 foraminiferal 6 1 8 0 on pH, Broecker (1995) suggests that a 0.2%o correction would be required to correct glacial 6 1 8 0 values for lower pH. This is equivalent to ~ 1°C cooler 6 1 8Oderived glacial temperatures. For example, ice-volume-corrected 6 1 8 0 measurements in planktonic forams would yield isotopic palaeotemperatures too cold by 1°C in the absence of corrections for the postulated pH effect. 51 4. Chronostratigraphv 4.1 Stratigraphy The sediments in Hole 887B are primarily composed of olive grey fine silt and silty clay with intercalated layers of ice-rafted debris (IRD; Table A2.1). Thin turbidite sections (1 to 4 cm thick), bioturbated sections, and mottled layers are present throughout the hole. Diatom-rich bands of pale olive colour and varying thickness (to > 1 m) characterized by very abundant diatoms > 150 um in size are common, as are intervals containing visible foraminifera. Homogeneous ash layers are prevalent with thicknesses up to 29 cm, notably in the following core intervals: 1H2- 28-36 2H5 - 0-29 2H5 - 46-60 2H6 - 42-58 3H3-47-54 For the sake of brevity, an extensive description of core lithology is omitted from this text. However, a detailed visual log is available from the author; as well, shipboard core descriptions and photos are to be found in Rea et al. (1993). 52 4.2 Composite Depth Model Based on correlation between high-resolution GRAPE and magnetic susceptibility records, missing intervals relative to holes 887A and 887C were spliced into the 887B profile to obtain a composite depth model (Table 4.1, Appendix 2). The estimated total of the missing sections is 3.37 m, yielding a corrected (composite) core length of 43.71 m. Spliced sections from 887A and 887C were overlapped where possible and extended to include disturbed sections of 887B in core 3H (upper 50 cm) and 4H (upper 17 cm). Overlaps were checked by examination of CaC03 and isotope profiles, and by aligning core photographs. In developing an age model, intervals representing "instantaneous" events (ash layers, turbidites) totaling 1.29 m were excluded (Table A2.1). Table 4.1. a) sections missing from 887B relative to 887A and 887C, based on correlation of GRAPE records; b) actual core depths + missing sections in (a). CORING GAPS Relative to 887A interval (mbsf) |missing (m) Relative interval (mbsf) to 887C missing (m) Between 1 and 2 1.90-3.66 1.76 Between 2 and 3 13.37-15.78 0.53 Between 3 and 4 23.95-25.6 0.92 Between 4 and 5 31.09-31.25 0.16 b). CORE Actual co core length (m) re depth bottom (mbsf) + Missing core length (m) sections bottom (mbsf) 1 • 1.88 1.88 3.64 3.64 2 9.72 11.60 10.25 13.89 3 9.68 21.28 10.60 24.49 4 9.40 30.68 9.56 34.05 5 9.66 40.34 9.66 43.71 53 4.3 Chronology Age control for the upper 887 section is based on accelerator mass spectrometry (AMS) 1 4 C dating of benthic and planktonic foraminifera in site survey cores PAR87-10 and PAR87-1 (Zahn et al., 1991a; T.F. Pedersen, unpublished data). In order to develop a radiometric age model, stratigraphic and geochemical data were used to correlate PAR87-10, PAR87-1, and 887B records. T.F. Pedersen provided extensive geochemical data for PAR87-10 and PAR87-1; as well, additional CaC03 analyses were done on trigger-core samples (PAR87-10) provided by Trudie Forbes (P.G.C., Sidney, B.C.) as a means of increasing resolution in the upper core. Together with 6 1 3 C o r g data from PAR87-1 (McDonald, 1993), this information facilitated the development of a common depth scale which indicates that sections of 7 cm and 45 cm are missing from the tops of PAR87-10 and PAR87-1, respectively. Hole 887B and PAR87-10 trigger core profiles appear similar and extrapolation using the sedimentation rate from the upper two age control points yields a core-top age for 887B of ~ 200 years, which is less than the error of 1 4 C dates. Hence, for the purposes of this study, a core-top age of zero is assumed for 887B. Radiocarbon dates were corrected by subtracting a reservoir age of 717 + 47 years from the initial 1 4 C results, as suggested by the average of five Northeast Pacific reservoir ages given by Southon et al. (1990); their reservoir age estimate at the 6360 ka time horizon was excluded. The 717 year correction is applied to all 1 4 C dates, though changes in reservoir age prior to 11 ka are unconstrained. In order to convert radiocarbon to calendar years, palaeoceanographic studies commonly use a linear calibration extending to 22 ka (calendar; Bard et al., 1993). However, since PAR-1 is dated to ~39 ka (cal.), a 54 tentative non-linear calibration by E. Bard is used here; this is based on two pairs of U/Th- 1 4 C dates older than 22ka (-30 ka; Keigwin & Jones, 1994): A § C a l e n d a r ) = "5 .85x 10"6(14C age c o r r.) 2 +1.39(14C agecon-.)-1807 Beyond the limit of 1 4 C dating, age control points were assigned based on correlation of 887 6 1 8 0 records with the SPECMAP stacked 6 ] 8 0 curve (Imbrie et al., 1984; Figure 4.1, Table 4.2). In the interval below Stage 16 (620 ka), the modification proposed by Shackleton et al. (1990) is applied. This accommodates a change between 620 ka and 800 ka from obhquity-dominated forcing (41 ka) to a climate system dominated by the eccentricity cycle (100 ka). This calibration places the magnetic reversal of the Matuyama-Brunhes boundary at 780 ka, consistent with recent radiometric constraints (Izett & Obradovich, 1994). Development of the time scale for 887B (age model 1) was based on linear interpolation between age control points (Paillard, 1994), the placement of which requires a certain degree of subjectivity given the discontinuous distribution of foraminifera throughout the core. Despite these limitations, the data suggest that the composite benthic-foram 6 1 8 0 profile reflects the global (SPECMAP) 6 1 8 0 signal (Figure 4.2). Identification of major isotopic boundaries is relatively straightforward and in agreement with previous work (e.g., Zahn et al., 1991). The planktonic 6 1 8 0 curve also reflects the global pattern generally, but is influenced to some extent by higher-frequency local variations (Section 5.3.1). Sedimentation rates average 6.6 cm/kyr, ranging from 1.7 to >15 cm/kyr (Table 4.2; Figure 4.3.), thus time resolution (calculated as the Nyquist period: N = 2[sample interval/sed. rate]) at 10 cm sampling intervals averages 3.03 kyr, with a range between 1.3 kyr and 11.7 ky. Resolution is further limited by bioturbation, seen to extend over irregular core intervals of up to 5 cm. 55 6 1 8 0 (per mil) 0 2.2/Tt -1 100 (1.2/135 f T 6.4H51 200 300 ro 500 600 700 18 800 3*0/24' 4.2/65 ^ ^ L w , , , , , , 6 . 6 / 1 8 3 1^4^5.3/99 5 . p i 094 w.5 ,.„.,„. ^> 6.5/171 sao« 5.5/122 - JSS»»7.1/194 8.2/249 7.4/228 «aiSS „.„«,., .8.0/245 *">8.3/257 7 . 3 / 2 1 $ S*» 7.5/238 8.4/269 <*£T. 8.6/299 10 . 2 / 341 8.5/287 9:0/303, 9.2/320 .«&M*rt«^'ww'»'-,;'wB!<fTu=M*,1'O?0'/a3'3'9, 11.2/375 ;«CT 400 1 2 . 2 / 4 3 4 MSST" .1-3,0/423^ 1 2 . 4 / 4 7 1 «*C ™ t > 9-1/310 SS» 9.3/331 3/405 J 2 . 3 1 / 4 4 3 I > 1 2 . 3 3 / 4 6 1 "T3':0/4"7'8«"»~»«»„Ig. 1:3 T 1 / 4 8 1 13.12/491 < > : . ' / S 0 2 4.0/524 14.2/53B 14^3 /552 1 4 " 4 / 5 6 p • < : ^i»5- .d/ -56 5- , „ I 15.2/585 <*£XZZZ i 1 5 . 4 / 6 0 7 5 » j15.1/574 j;..- 1 5.^ /596 _ .....v-.*.•*"~-''-"*^ 6.0/620' 16.23/6$7V^,_ ; b 2/702 ^ " : > 1 7 - 1 / 6 9 3 2 / 7 2 0 ^J4o/7ir—^1 7 -3/7io 18-4/750 %, 20.2/793 «y »20:O/-788' 19.0/.7.7.0 3*19.1/783 5.5/617 - 2 -1 Figure 4.1. SPECMAP composite S 1 8 0 from Imbrie et al. (1984), as modified by Shackleton et al. (1990). Numbers denote Stage/Age (ka) Table 4.2. Age control points for Hole 887B. Sedimentation rates* represent the intervals following, e.g., sed. rate for 0.3 - 0.6 mbsf = 4.6 cm/kyr. . Composite depth, m Comp depth - ash layers Isotopic Event Foraminiferal Species 14C Age**, ka Calendar Age, ka Sed. Rate cm/ka 0.30 0.30 N. pachyderma 5.923 ± 304 6.221 4.6 0.60 0.60 Mixed Planks. 10.923 ± 1 1 0 12.678 3.4 0.63 0.63 Mixed Planks. 11.628 ± 1 1 5 13.565 14.3 0.75 0.75 Mixed Planks. 12.298 ± 130 14.402 10.3 1.24 1.24 Mixed Planks. 16.168 ± 165 19.137 7.1 1.38 1.35 N. pachyderma 17.463 ± 170 20.683 5.3 1.57 1.53 N. pachyderma 20.383 ± 220 24.095 12.0 1.90 1.78 N. pachyderma 22.203 ± 230 26.171 6.2 2.56 2.44 N. pachyderma 32.183 ± 700 36.868 4.7 2.67 2.55 Mixed Planks. 34.523 ± 920 39.208 7.2 4.15 3.98 4.0 59 5.7 4.83 4.66 5.0 71 11.0 6.17 5.98 5.2-5.1 83 7.1 6.95 6.76 5.3-5.2 94 7.7 7.72 7.53 5.4-5.3 104 15.1 9.38 9.19 5.5-5.4 115 5.0 10.38 9.84 6.0 128 6.4 12.06 11.31 6.4 151 6.1 14.05 13.26 6.6 183 5.5 17.49 16.65 8.0 245 5.2 18.11 17.27 8.3 257 6.5 20.91 20.01 8.6 299 4.9 21.98 21.08 9.3 331 8.0 23.40 22.49 10 350 8.2 27.99 27.01 11.3 405 1.7 28.30 27.32 12.0 423 4.5 32.41 31.34 13.2 513 5.0 33.71 32.58 14.2 538 5.1 35.10 33.97 15.0 565 4.5 37.64 36.47 16.0 620 5.2 38.89 37.72 16.3-16.2 644 4.5 41.97 40.68 17.3 710 5.5 43.17 41.88 18.3 732 5.5 45.86 44.52 B/M 780 Based on calendar age and comp. depth, minus ash layers and turbidites Reservoir corrected 717 ± 47 years. 57 887 benthic comp. 5 1 8 0 (%o) -o- 887 G. bulloides S 1 8 Q (%») Revised S P E C M A P 6 1 8 Q (%») Figure 4.2. 887B time series of benthic composite (see Section 5.3.1) 6 1 8 0 (%o) and G. bulloides 6 1 8 0 (%o) with isotopic stage boundaries noted. The SPECMAP stack, represented by the solid curve, is from Imbrie et al. (1984), as modified by Shackleton et al. (1990). (continued) 887 benthic comp. S 1 8 0 (%o) 887 G. bulloides S 1 8 0 (%o) 5.0 4.5 4.0 3.5 3.0 4.0 3.5 3.0 2.5 2.0 Revised SPECMAP 6 1 8 Q (%») Figure 4.2 (completed) 59 0 100 200 300 400 500 600 700 800 Age (ka) Figure 4 .3 . Age -dep th r e l a t i onsh ip i n 887B. Sed imen ta t i on rates between contro l points average 6.6 c m / k y , wi th a range of 1.7 to 15.1. A broad association exists between h i g h sedimentat ion rates (calculated based on chronologic control points) and the deposi t ion of diatomaceous bands; for example, a h igh sedimentat ion rate of 11 c m / k y r occurs dur ing a pe r iod of d ia tom ooze deposit ion f rom 83 to 71 ka . This supposi t ion is supported by 2 3 0 T h data w h i c h indicate that ve ry h igh sedimentat ion rates are associated wi th the d e p o s i t i o n o f some d i a tomaceous l aye r s . T h e r a t i o n a l e for e v a l u a t i n g palaeoflux to the seafloor based o n 2 3 0 T h data rests on the p r inc ip l e that the flux of scavenged 2 3 0 T h to sediments is balanced by its local p roduc t ion rate i n the water c o l u m n v i a U decay (Bacon, 1984). This is due to the short residence time of 2 3 0 T h relative to timescales of la teral ocean mix ing . As a result, excess 2 3 0 T h ac t iv i ty i n settling part icles is inverse ly re la ted to the total mass f lux; 60 thus the flux of sedimentary components can be estimated using decay-corrected 2 3 0 T h activity ( 2 3 0 T h o ) as a reference. To illustrate, 2 3 0 T h values in core PAR 87-1 indicate a doubling of sedimentation rates during deposition of the aforementioned diatomaceous bands during late Stage 5 and on the Stage 5/4 transition (Figure 4.4). This suggests that the diatomaceous strata were deposited at approximately 10 to 12 cm/ka. Although quantitative interpretation of sedimentation rates may be complicated by irregularities in the 887 age model, it is evident that mass flux to the seafloor controlled by export production was substantially increased during deposition of diatomaceous strata at Site 887. 61 Si/Al (wt. ratio), %CaC03 0 1 0 2 0 3 0 0 2 4 6 8 1 0 Excess 2 3 0Th° (dpm/g) Figure 4.4. Excess 2 3 0 T h , %CaCOa, and Si/Al values for core PAR 87-1. 4.3.1 Comparison with Biostratigraphic Data At some high-latitude locations, relative abundance changes in C. davisiana appear to vary coherently with the global isotope curve, increasing in abundance during glacial times relative to interglacial periods (Morley & Hays, 1983). Based on a composite C. davisiana curve and last occurrence datums (LODs) of radiolarians, Morley et al. (1995) have developed a chronology for three Leg 145 sites, including Site 887 (Table 4.3). Morley states that "the exact causes of such variations (in C. davisiana) are still unknown, but sea-surface temperature does not appear to be a dominant controlling factor". A number of discrepancies are apparent between the 887 C. davisiana chronology of Morley et al. (1995) and the 14C/isotope-derived chronology (age model 1), a partial list of which follows: 1) The two age models differ in their positioning of the 6/7 boundary by 1.15 m; the C. davisiana model fixes the 6/7 boundary (186 ka) at 13.06 mbsf, a depth close to the 6.4/6.5 boundary (167 ka) according to age model 1. 2) Morley et al. (1995) suggest that the interval from ~ 180 ka through 300 ka (oxygen isotope stages 7-8) was not recovered, though age model 1 indicates that only the interval from ~ 172 ka to 190 ka is absent from this section (i.e. the gap between cores 2 and 3 including the disturbed upper 50 cm of core 3). As a result, Morley's placement of the 7/8 boundary is ~ 3 m shallower than the equivalent depth marked by age model 1. Presumably due to this and to the disparity in (1), substantial differences occur between the two chronologies in stages 7 through 8. 3) Morley et al. (1995) establish the LOD of S. universus on the 12/11 boundary which they place at 450 ka. If shifted to conform with the SPECMAP 63 Table 4 .3 . 887B age control points based on C. davisiana patterns and last occurrence datums of radiolarians; some points were correlated by the author to 887A data (unpublished data from Joe Morley)- Bracketed numbers indicate alternate placement of the 11/12 boundary (see text). Comp dept l -ash Composite depth, m Age, ka radiolarian sed rate cm/ka Species Marker* 0.65 0.65 11 8.54 C. davisiana 2.0 1.76 1.88 24 5.27 C. davisiana 3.0 3.13 3.30 50 8.40 L. nipponica LOD 4.39 4 .56 65 10.09 C. davisiana 4.2 7.82 8.01 99 8.26 C. davisiana 5.3 9.72 10.26 122 5.00 C. davisiana 5.5 10.02 10.56 128 8.27 C. davisiana 6.0 11.84 12.61 150 1.25 C. davisiana 6.? 12.29 13.06 186 3.22 C. davisiana 7.0 12.58 13.36 195 1.90 C. davisiana 7.1? 12.77 13.56 205 3.64 C. davisiana 7.2? 13.17 13.96 216 1.72 C. davisiana 7.3? 13.67 14.46 245 7.45 C. davisiana 8.0 21.49 22.39 350 6.02 (8 .25) S. aquilonium LOD 27.51 28.49 4 5 0 ( 4 2 3 ) S. universus LOD Age, ka * * age model 1 Rad. age -age model 1 13.7 -2.7 26.0 -2.0 47.2 2.8 66.2 -1.2 105.9 -6.9 125.6 -3.6 130.8 -2.8 159.7 -9.7 167.1 18.9 171.8 23.2 175.0 30.0 181.5 34.5 190.5 54.5 336.5 13.5 4 2 7 . 2 5 4 22.7 (-4.3) * numbers indicate isotopic stages; LOD * * f rom the model in Table 4.2 last occurence datum 64 chronology, this event would occur at 423 ka and both age models would agree within 4 ka at that depth in the core (28.49 mbsf). Although trends in the C. davisiana data are broadly consistent with the benthic 6 l s O profile of 887B, C. davisiana variations appear significantly coherent with the local A 6 1 8 0 (planktonic-benthic) curve (Figure 4.5; Section 5.3.1). This suggests that C. davisiana abundance is influenced to some extent by regional changes which affect both species' (C. davisiana and G. bulloides) environments, yet do not alter the benthic 6 1 8 0 signal. 65 887 benthic comp. S 1 8 0 (%o) 887 Local S 1 8 0 (%o) 887 % C. davisiana — i — 887 % C. davisiana — i — Figure 4.5. Distributions of %C. davisiana, benthic 6^80, and local (planktonic -benthic) A6 1 8 0 with depth in 887B; note reversed C. davisiana scale. Stratigraphic legend appears on following page. Isotopic stages according to age model 1 (see text). (continued) 66 8 8 7 benthic comp. 6 1 8 0 (%o) 8 8 7 Local 6 1 8 0 (%o) — ^ 5.0 4.5 4.0 3.5 3.0 0 -0.5 -1 -1.5 -2 Stage 40 30 20 10 0 40 30 20 10 0 8 8 7 % C . davisiana — i — 8 8 7 % C . davisiana — I — Figure 4.5 (completed) diatom ooze clay/silt clay/silt (mottled) ash 67 5. Results and Discussion 5.1 Sedimentary Geochemistry Assuming that sedimentary aluminum is exclusively of aluminosilicate origin, elemental concentrations in this study are normalized to Al content in order to illustrate more clearly compositional variations which are unrelated to variable inputs of lithogenic detritus. Recently, Murray et al. (1993) cast some doubt on this assumption, suggesting that in the low-latitude Pacific, the Al content of settling particles includes a component adsorbed from seawater. They noted Al/Ti ratios on the order of 60 to 100 in highly biogenic oozes. Down-hole Al/Ti ratios (excluding ash bands) at Site 887 average 17 (a= 1), varying only from 13 to 23. These values do not vary significantly from the average crustal ratio of 16 (Krauskopf, 1979), suggesting that scavenging of Al from the water column has had an insignificant effect on the aluminum content of bulk sediments at this location. 5.1.1 Si/Al Si/Al weight ratios range from 2.95 to 30.45, the higher values occurring in diatomaceous strata3, for example at ~ 4.69-5.19, 5.69-6.79, and 7.79-8.19 mbsf (Figure 5.1; Table A2.1). The high Si/Al ratios in such intervals reflect the %opal content: they are substantially higher than in other strata, reaching in Stage 5, for example, ten times the average crustal value ( S i / A l c r u s t a l = 3.48; Krauskopf, 1979). Hence, Si/Al ratios are used here as a proxy for sedimentary opal content. Nevertheless, diatomaceous layers that are visually defined in split core do not correspond to a uniform degree of silica 3Diatoms of > 150 \im in size are consistently abundant in such intervals. 68 Figure 5.1. Ice-rafted debris (IRD) abundance, Si/Al weight ratio, %CaCO 3, and CaCO: accumulation rate (g/m 2a) vs. depth at Site 887 (legend appears on following page); GRAPE density composite is shown at left for comparison. Italicized numbers indicate light S 1 8 0 excursions. (continued) 69 Si /A l — CaC0 3 (%) — 0 5 10 15 20 25 30 35 0 5 10 15 20 25 30 35 40 stage 2 2 9 n vr r c a IRD abundance va very abundant a abundant c common r rare vr very rare n none 0 0.5 1.0 1.5 CaC0 3 acc. rate (g/cm2ka) -diatom ooze clay/silt clay/silt (mottled) Figure 5.1 (completed) enrichment; to illustrate, Si/Al ratios in the uppermost diatom band recorded in the log at ~ 0.50-0.65 mbsf reach only 5.26. Such variability is likely due to dilution by detrital clay and carbonate phases, since observations at the microscope show the diatomaceous strata in 887B are largely devoid of ice-rafted debris (IRD; Figure 5.1). Furthermore, the diatom-rich bands are coeval with both GRAPE and dry bulk density minima (Table A2.1). Changes in opal accumulation at Site 887 do not appear to have been significantly affected by post-depositional sediment redistribution. This is supported by observed similarities in Si/Al, %CaC03, Ba/Al, and 613Corg records between cores 887B, PAR87-10 and PAR87-1 (Figure 5.2 and data in McDonald, 1993). It is unlikely that redistribution would affect all three cores similarly, since PAR87-1 was collected 65 km distant from PAR87-10 and 887B sites. Opal enrichment does not generally occur during full glacial times, as can be seen in Stages 2.2, 4.2 , 6.4 to 6.2, 8.2, 10.2, 12.2, 16.2, and 18.2 (Figure 5.2)4. However, Si/Al maxima do occur in association with subglacial conditions during Stages 6.5, 8.5, 12.3, 14.3, and 18.3 (Figure 5.2). Hence, contrary to patterns observed in the western subarctic Pacific (Haug et al., 1995), high opal values are not constrained to interglacial periods through the Brunhes. Moreover, though no consistent cyclicity is apparent at the substage level, opal is commonly enriched in cooler interglacial substages such as 5.2, 5.4, 7.4, 9.2, 11.2, 13.12, 13.2, and 15.2, and on transitions (e.g., 2/1, 5/4, and 13/12) though peaks also occur rarely during full interglacial times (e.g., Stage 9.3). glacial and interglacial stages with corresponding ages are illustrated in Figure 4.1. 71 M n / A l - ° - 5 1 3 C o m (%o) ^ S i /A l — 0.05 0.1 0.15 0.2 -24 -23 -22 -21 -20 5 10 15 20 25 30 35 numbers correspond with O j o c a j excusions; numbered sections at the right represent carbonate stage boundaries from Farrell and Prell (1989). Revised SPECMAP 6 1 8 0 curve (see Section 4.3) shown for reference; dotted circles indicate the presence of pteropod fragments. (continued) 5.13 MnlAl Mn concentrations in 887B range from 192 to ~ 9200 ppm (Table A2.1), and Mn/Al varies from 0.008 to 0.221; for comparison, crustal values are 1000 ppm and 0.012, respectively (Krauskopf, 1979; Figure 5.2). The surficial Mn spike in 887B probably represents M n 2 + remobilized from underlying reduced sediments and reprecipitated as Mn (IV) oxyhydroxides in the near-surface aerobic zone. Froelich et al. (1979) suggest that the depth of this diagenetic "trap" is determined by the balance between upwardly diffusing M n 2 + and O2 diffusing downward, and that both species are exhausted within the spike. Hence, a solid MnC»2 peak occurs near the top inflection in the dissolved M n 2 + profile as it increases downwards (Figure 5.3). Mn concentrat ion Figure 5.3. Schematic diagram of dissolved and solid phase Mn profiles in a hypothetical steady-state system (redrawn from Froelich et al., 1979). 74 Downcore Mn variations can reflect burial of surficial Mn enrichments produced by changes in the flux of organic matter to the seafloor (Finney & Lyle, 1988). This occurs as an increased flux of organic matter causes a shoaling of the Mn redox boundary "stranding" a deeper Mn oxide peak which may survive subsequently due to incomplete reductive remobilization. Such a mechanism may to some degree be responsible for Mn enrichments in 887B which are coincident with increased biogenic deposition, e.g. the diatom bands at ~ 72, 88, 324, and 514 ka The similarities between downcore Fe and Mn profiles suggest that diagenetic transport of Fe may have to some degree occurred in association with Mn remobilization (Figure 5.4). However, it is likely that most of the downcore Mn and Fe is aluminosilicate bound; thus, the correspondence in Figure 5.4 may be a reflection of the varying mineralogy of detritus, rather than an indication of authigenic minerals. An alternative explanation for significant downcore Mn enrichments may be involve the formation of a manganese carbonate phase (kutnahorite; e.g., Pedersen & Price, 1982) formed under reducing conditions below an oxic layer. Degradation of an increased flux of organic matter to the seafloor during episodes of high export productivity could result in O2 depletion at the sediment-water interface if the rate of O2 consumption exceeds the rate of supply. Under such conditions, subsequent utilization of pore water oxygen promotes a shallower oxic boundary below which degradation processes proceed with oxidants reduced in the order of decreasing energy yield: NO 3 - » M n 4 + » Fe 3 + (suboxic) » S 0 4 2 " (anoxic) (Froelich et al., 1979). Such a shoaling of the redoxcline would result in the concentration of diagenetically cycled M n 0 2 into a thinner zone. Burial of this particularly Mn-rich horizon would support the remobilization of Mn at relatively shallow 75 M n / A l — Fe/AI — I R D — 0.05 0.1 0.15 0.2 0.4 0.6 0.8 n vr r c a va 0 1000 2000 3000 4000 5000 1000 2000 3000 4000 n vr r c a M n ( p p m ) — Fe ( p p m ) — Figure 5.4. 887B time series of Mn/Al, Fe/AI, Mn and Fe (ppm), and ice-rafted debris (IRD). Revised SPECMAP 6 1 8 0 curve (see Section 4.3)is shown for reference; legend appears on following page. (continued) Mn/A I — 0.1 0.15 i . . . . i . . . i . . . . i . . . i . . . . i . f 1000 2000 3000 4000 5000 M n ( p p m ) — 1000 2000 3000 4000 F e ( p p m ) — Figure 5.4 (completed) va very abundant a abundant c common r rare vr very rare n none 77 depths, producing high concentrations in pore water. This "manganese pump" will cause precipitation of Mn carbonate if the solubility product is exceeded (Calvert & Pedersen, 1996; Figure 5.5). Mn02 oxic Mn^ 4 " released, M1-1CO3 precipitates sulphate reduced t alkalinity, | Mn Figure 5.5. Schematic representation of Mn carbonate formation. Although Mn carbonates have been primarily found in anoxic environments, production of alkalinity due to sulphate reduction does not appear to be requisite for Mn carbonate production in hemipelagic sediments. Instead, Mn oxyhydroxide supply and the "catalytic" effect of coarser grain size on carbonate nucleation may control Mn carbonate formation in such environments (Pedersen & Price, 1982). Indeed, porewater so4 2 " concentrations decrease only a small amount in 887B from 28.6 mmol/L at 1.45 mbsf to 25.7 mmol/L near the bottom of the hole (Pedersen & Ingram, 1994), indicating that intense anoxic diagenesis does not occur through the length of the hole. Sti l l , sulphate reduction has occurred, presumably in microenvironments, evinced by rare framboidal pyrite noted during foraminifera counts of 887B samples. In the northwest Pacific, authigenic Mn carbonate has been noted previously in diatomaceous sediments which are slightly reducing and contain only moderate amounts of carbon (0.26-0.74% Corg! refs. in Pedersen and Price, 1982). Similar strata in 887B show Mn enrichments (e.g. at 86, 420, 463, 514, and 78 708 ka.) which may reflect Mn carbonate formation induced by a strengthened manganese pump in response to redoxcline shoaling. The occurrence of such redox changes is also evinced by trace metal enrichments concurrent with diatom ooze strata through Stage 5 in core PAR87-10 (McDonald, 1993). Moreover, Mn/Al is positively correlated with %CaC03 (R = 0.65), and cross-spectral analysis intimates consistent coherence at the 80% confidence level. Significant M n / A l maxima are coeval with CaCC>3 enrichments and independent of increased opal at ~ 130,440, 540, and 630 ka (Figure 5.2). Mn carbonate was not directly observed in preliminary scanning electron microscopy of 887B samples at 10, 86, 420, and 708 ka. Likewise, EDS (X-ray) techniques were inconclusive in the separation Mn oxyhydroxide and Mn carbonate signatures. However, the presence of Mn carbonate in 887B cannot be ruled out without further analysis of these and other Mn-rich samples. In any case, the occurrence of high Mn concentrations whether in oxyhydroxide or Mn carbonate form, indicates that deposition occurred under oxygenated conditions (Calvert & Pedersen, 1996). 5.1.3 Ba/Al Given the variable dilution downhole by non-biogenic components, the Ba/Al ratio rather than excess (bio) Ba (Dymond, 1992) will be used here as a qualitative measure of palaeoproductivity. Ba/Al values in 887B vary from 0.006 to 0.074 (o = 0.01; Table A2.1). Peak Ba concentrations are closely linked with biogenic silica deposition, as illustrated by the strong association between Ba/Al and Si/Al maxima (Figure 5.2). Ba/Al excursions are also concurrent with %CaCC>3 and 613Corg maxima and are broadly associated with C o r g enrichments. 79 Ba/Al ratios vary widely amongst different rock types, the provenance of which is not well known at Site 887. Nevertheless, the range of Ba/Al ratios in common rocks 5 is much less than the variations seen in 887B; moreover, high Ba/Al values are associated with biogenic intervals such as the diatom oozes in Stage 5 (e.g., @ 72ka, Ba/Al=0.057) which are virtually devoid of IRD. This suggests that significant downcore Ba/Al excursions are not the result of changing aluminosilicate mineralogy or sediment provenance. The association between Mn enrichment and organic matter deposition may alter and to some extent amplify the productivity signal recorded by the Ba/Al profile (Section 3.1.3). Mn/Al and Ba/Al distributions do covary in 887B, yet no consistent relationship exists between Mn and Ba enrichment, nor is it clear which Mn phase (e.g., oxyhydroxides, MnCC>3) is dominant. As a result, it is difficult to quantify authigenic Ba contributions. Dymond et al. (1992) indicate that ferromanganese oxyhydroxides may account for up to 30% of Ba in surface sediments from the Equatorial Pacific. As a steady state feature, Ba enrichment would not significantly affect the downcore Ba signal, except at depths where increased Ba occurs in association with Mn oxyhydroxide remobilization (~ 40-50 cm depth in 887B). However, some barium may be associated with buried, relict Mn enrichments. This process may be in part responsible for Ba maxima associated with Mn peaks, for example at 86 and 514 ka. Ba/Al end members: basalt=0.004, granite=0.01 (Krauskopf, 1979). 80 5.1.4 Organic Carbon The mean organic carbon (C o r g ) content (wt%) in 887B is 0.43%, varying from 0.03 to 1.03%; a = 0.14 (Table A2.1). The C o r g mass accumulation rate 6 (MAR) profile is grossly similar to that of % C o r g , though significant discrepancies do exist between the two signals (Figures 5.2 and 5.6). Notably, significant carbon enrichment does not generally occur in biogenic silica-rich (diatom ooze) strata (e.g., at ~ 69, 88, 308, 464 ka), and minima in C o r g MAR are for the most part more pronounced such intervals. If high opal abundance reflects an increased supply of biogenic matter to the sediments in response to surface productivity, then an increase in C o r g MAR would be expected during periods of deposition of the diatomaceous bands. The resolution of the 887B timescale does not permit accurate determination of sedimentation rates specifically for such layers. However, calculation of mass accumulation rates from a more detailed chronology may reveal that the rate of carbon accumulation during such times was rapid; this is supported by 2 3 0 T h dta from core PAR87-1 (Section 4.3). The C o r g MAR profile indicates that C o r g burial was enhanced during glacial periods in the Upper Brunhes (Stages 2, 6, 8, and 10), yet increased carbon burial is not invariably constrained to glacial episodes, as evinced by increased carbon accumulation in early Stage 5 and in Stage 11. Although no regular or cyclic variation occurs in the %C 0rgandC 0rg MAR profiles through the Brunhes, significant minima consistently occur during interglacial inceptions. Across glacial stages 2, 6, and 10 which precede such minima, carbon content increases by up to a factor of three to peak values at glacial termination, increased C o r g presence infrequently occurs in association with 6MAR = (sedimentation rate) x (concentration) x (dry bulk density) 81 o 1 3 C o r g (%o) -o— C o r q acc. rate (g/cm 2ka) — -25 -23 -21 -19 0 0.02 0.04 0.06 0.08 0.1 Figure 5.6. 887B time series of 6'-'C o rg (%o), C o r g / N , % C o r g , and C o r g accumulation rate (g/m2a). Italicized numbers represent 8 1 8 0 | o c a j excursions; ash indicates intervals of probable ash dilution. (continued) Figure 5.6 (completed) 83 CaGC»3 maxima which are unaccompanied by increased opa l content, as at ~140-130, 260-240, and 546 ka. Profound C o r g enr ichments also occur independen t ly of increases i n S i / A l , B a / A l , and CaCC>3 (e.g., at ~ 17, 79, 109, 187, and 600 ka), suggest ing that c a r b o n p re se rva t ion i n these in te rva l s was enhanced b y factors independent of organic carbon flux to the seafloor, such as increased bu lk sediment accumula t ion rate or changes i n sediment texture (see reviews by Calvert, 1987; Hedges & Ke i l , 1995). Al te rna t ive ly , increased carbon accumula t ion may occur i n response to increased p r o d u c t i o n of phy top l ank ton w h i c h have no sil iceous or carbonate h a r d parts such as the p rymnes iophy te Phaeocystis pouchetti, though this is not consistently supported by B a / A l and S 1 3 C o r g evidence (e.g. Stage 2 carbon enr ichment ) . Phaeocystis can r a p i d l y export large amounts of organic matter to deep water as mar ine snow (Wassman et al. , 1990) and its dominance w o u l d leave l i t t le evidence of h i g h p roduc t iv i ty i n the fo rm of skeletal remains. This species has been noted by Booth et a l . (1993) as an impor tan t a n d re la t ive ly abundant p r i m a r y producer at Station P; m a x i m u m concentrations ( > 1 0 6 / L ) of over an o rder of magni tude greater than those seen i n the An ta r c t i c have been observed i n the mode rn subarctic Pacific (refs. i n Booth e ta l . , 1993). 84 5.2 Carbonate Records 5.2.1 Carbonate Presence Carbonate concentrations range from 0 to 39% throughout the hole. Abrupt changes in content are common and little or no CaCC»3 is present in many intervals (Figure 5.1). %CaCC»3 andCaCC>3 mass accumulation rates (MAR) covary, suggesting that %CaC03 is not controlled by non-carbonate dilution. A close relationship is evident between %CaCC>3 maxima and climatic cycles, characterized by higher carbonate content in glacial sediments. The largest %CaC03 maxima generally appear during the latter half of glacial intervals and on glacial-interglacial transitions. An anomalous CaCC>3 enrichment with respect to this pattern occurs during Stage 3 which is a relatively minor interglacial stage. The highest enrichments are associated with glacial-interglacial transitions; the spike centered on the 2/1 boundary has been observed in a number of North Pacific cores (e.g., Karlin et al., 1992; Keigwin et al., 1992). A long-term change in CaCC>3 preservation is apparent in the 887B record, manifest by a minimum in CaCC>3 presence at mid-depths in the hole. This corresponds to the mid-Brunhes dissolution event, which extends from about 200-450 ka and has been reported in deep sea records from the North Atlantic and Pacific (Crowley, 1985) and the Indian Ocean (Peterson & Prell, 1985). The cause of this event is unknown, and may be the result of a long-term excursion of the deep sea carbonate reservoir (Crowley, 1985). The average interval between carbonate events (excluding the mid-Brunhes gap during Stages 9-10 and taking the peaks at 44, 132, 240, 423, 535, 626, and 711 ka) is 111 kyr, suggesting a 100 kyr cyclicity. This is supported by cross-spectral analysis with the SPECMAP stack which shows greatest power in 85 the 100 kyr band; significant peaks also occur in the obliquity and precession bands (Fig 5.7). Similar 100 kyr periodicities have been documented in a number of equatorial Pacific studies and in NE Pacific records extending to 300 ka (Karlin et al., 1992). Moreover, 887B carbonate maxima generally occur in intervals of enhanced carbonate preservation described by Farrell & Prell (1989); pteropod fragments are also observed in the same maxima (Figures 3.3 and5.2)7. 125-95 50-40 23 19 1500 t j 1000 c o 500 1 1 ' I ' 1 " 1 1 '•• • " 1 i 1 " 1 1 1 1 1 " • i i — LO -. 1 . . . . . . i . . . . i . : . . . i : . . . i . . . . Frequency 0 0.01 0.02 0.03 0.04 0.05 0.06 0.07 Period 100 50 33 25 20 17 14 Figure 5.7. Results of cross-spectral analysis of the 887B CaCC>3 record with the SPECMAP stack. 7 The aragonitic pteropod shell is much less resistant to dissolution than the calcite tests of coccoliths and foraminifera, thus the presence of pteropod remains intimates enhanced carbonate preservation. 86 The coherence between equatorial and NE Pacific records suggests that a similar process or processes has controlled the CaCC>3 content of the deposits in both regions. Other NE Pacific records have indicated a similar relationship; Karlin et al. (1992) proposed that a migration of the CCD from 4500 m to less than 2700 m has occurred off the Oregon coast over the past 45 kyr. The magnitude of this CCD shift (>1800 m) is much greater than the 400-800 m variations seen in the equatorial Pacific lysocline, and is in relative agreement with the suggestion by Zahn et al. (1991b) of a major (almost 2500 m) glacial-interglacial lysocline shift. Such amplification of CCD variations indicates that some regional effect(s) in addition to changes in NADW production may influence NE Pacific carbonate cycles. High surface production could provide a mechanism for a shallow interglacial CCD via the release of CO2 from the oxidation of organic matter. Archer (1991) has shown that under oxic conditions when COrg/CaC03 mass flux ratios exceed 1.0, increased productivity enhances CaC03 dissolution. At lower ratios, dissolution is dominated by the CaC03 flux resulting in CaC03 preservation. From a compilation of sediment trap data, Karlin et al. (1992) indicated that relatively high Corg/CaC03 ratios (>1) characterize the NE Pacific, while lower ratios (<1) are found in the equatorial Pacific. Thus, changes in productivity may have different effects on CaC03 dissolution in each region. Haug et al. (1995) have suggested that increased surface productivity driven by enhanced upwelling may have occurred during interglacial periods through the Brunhes Chron in the subarctic Pacific. During such times, the Corg/CaC03 ratio of particulate organic matter (POM) was probably also high (increasing with C o r g flux), thereby increasing CaC03 dissolution. By this process, the CCD could have been driven upward, shoaling to lesser depths 87 than in the equatorial Pacific. Nevertheless, in the case of unusually high carbonate production and settling flux, local dissolution may be surpassed resulting in the occurrence of sedimentary CaCC>3 maxima as suggested by Haug et al. (1995) for the northwest Pacific. A similar mechanism may explain carbonate maxima at Site 887, particularly during times of enhanced carbonate dissolution elsewhere in the Pacific (Farrell & Prell, 1989). 5.2.2 Foraminiferal Abundance %CaCC>3 carbonate in 887B is controlled to a large extent by foram abundance with a few exceptions, notably from about 404 to 420 ka (Figure 5.8). In the absence of planktonic forams this increase in CaCC>3 is likely due to other carbonate producers, in particular coccolithophorids. Echinoderm spines are common in this interval; pteropod fragments and limited numbers of benthic forams including Pyrgo, Melonis, Triloculina, Gyroidinoides, Cibicides, and Uvigerina are also present. Increased coccolith numbers have been noted elsewhere in the section, particularly on the Stage 2/1 transition in core PAR 87A-10 (A. deVernal, unpublished data). Carbonate from authigenic or non-biogenic sources such as IRD could also contribute to the high CaCC>3 content in the 404 to 420 ka interval; however, no carbonate-bearing IRD phases were observed. 88 5.2.2.1 Planktonic G. bulloides is believed to live in the near-surface layer above the thermocline in the Gulf of Alaska. This species appears from April through August with greatest abundances coinciding with high productivity in June and July. At this time, surface waters are warm and thermal stratification increases. However, abundant G. bulloides have been noted during winter months in years of anomalously high fall/winter sea surface temperatures. N. pachyderma, considered to live at or below the thermocline, dominates the foraminiferal population throughout the year, comprising up to 40% of the annual population. Peak (sinistral) shell production occurs in spring, when the upper water column is coldest and nearly isothermal. During the fall, N. pachyderma abundance increases to a lesser degree, coincident with a shift from left to right-coiling forms (Sautter & Thunell, 1989). In ocean regions of lower summer sea surface temperatures (< 5°C), the relative abundance of N. pachyderma increases to about 95% of the fauna (Be, 1977). Surface sediment assemblages at Station P reflect spring conditions. N. pachyderma comprises about 90% of the planktonic foraminifera in surface sediment samples, with the remaining fraction dominated by G. bulloides (Sautter & Thunell, 1989). In 887B, abundance maxima of both species are greatest during glacial and minor interglacial periods, and are focused within intervals of intensified carbonate preservation described previously (Figure 5.8). Virtually no planktonic forams are present in the mid-Brunhes. G. bulloides dominates during minor interglacial stages 3, 5.2, 5.4, and 15.2 as well as on glacial-interglacial transitions 2/1, 13/12, and 18/17; peaks also occur in glacial stages 2 and 12. N. pachyderma is dominant during glacial stages 6, 8, and 14; dominance at ~ 7 and 56 ka may be due to an increase in dissolution of G. bulloides which is the more easily dissolved species (Thunell & Honjo, 1981). 89 Planktonic forams/g 0 2000 4000 Benthic forams/g 0 4 8 j stage^ 2 Figure 5.8. 887B foraminiferal abundance shown with Si/Al and %CaCO 3 profiles for reference. Italicized numbers correspond with 8 ^ 8 O j o c a j excursions; carbonate stage boundaries from Farrell et al. (1989) are indicated at the right. Foram counts are listed in A2.3. (continued) Figure 5.8 (completed) The observed shifts in foram distributions suggest that past variations have occurred in the regional hydrographic regime. Episodes of G. bulloides dominance indicate warmer and stratified surface ocean conditions, whereas increased N. pachyderma abundance may be due to a colder, more isothermal environment. Foram abundance peaks generally correspond with 6 1 8Q Dcal events, suggesting that changes in local hydrographic conditions affect foram abundance independently of longer-term dissolution cycles. 5.2.2.2 Benthic As with planktonic forams, abundance maxima of benthic forams are focused on intervals of intensified carbonate preservation, with none present in the mid-Brunhes. Gyroidinoides and Uvigerina are thought to inhabit similar endobenthic environments (Zahn et al., 1991a), though their distributions vary in 887B. Uvigerina peaks occur near the 2/1, 4/3, and 7/6 boundaries and in stages 11 and 14, and are coincident in the upper Brunhes with planktonic foram maxima. Gyroidinoides quantities more closely follow planktonic abundance throughout the Brunhes. It is speculative to draw conclusions from the relationships between the two genera here, since benthic foram numbers are small and the addition of one or two forams may significantly affect relative abundance. However, if Gyroidinoides flourishes as does Uvigerina in response to a high rain rate of organic matter (e.g., Pedersen et al., 1988), then its abundance profile would reflect a response to food supply from the surface. The more frequent occurrence of Gyroidinoides downcore may reflect higher resistance to dissolution relative to Uvigerina. Cibicides is epifaunal, inhabiting substrates protruding above the sediment surface (Corliss, 1991). Its abundance in 887B follows that of other benthic taxa, yet a number of Cibicides maxima occur independently. These are 92 associated with biogenic silica enrichments, for example at 82, 105, 280, 302, 510, and 605 ka. During these times, it is possible that increased oxygen demand in response to high organic matter flux causes a reduction in surface sediment oxygen levels. Once below a threshold O2 concentration, infaunal taxa such as Uvigerina and Gyroidinoides may become oxygen-limited, whereas the epifaunal environment remains oxygen-replete; this could account for the greater observed abundance of Cibicides. 93 5.3 Foraminiferal Isotopes 5.3.1 Oxygen The inconsistent distribution of foraminiferal species downcore required the development of a composite benthic record based on mean 5 l s O values (including replicates) of benthic species Gyroidinoides, Uvigerina, and Cibicides spp. from a given sample (Table A2.3, Figure 5.9). All Cibicides data were corrected for specific fractionation by adding 0.64%o (Shackleton & Opdyke, 1973). Gyroidinoides 6 1 8 0 values remain uncorrected since no significant offset is apparent between Gyroidinoides and Uvigerina 6 1 8 0 , as observed previously by Zahn et al., 1991 (Figure 5.10). 5 4.5 C3 3.5 2.5 y = 0.71 + 0.82x, r= 0.92 A Gyroid y = 0.00 + 0.84x, r= 0.90 0 Cibicid I 1 1 6 1 80 o O o o o ° o 0 1:1 _L • • I . . . I I 2.5 3.5 4 U v i g e r i n a 4.5 Figure 5.10. Paired benthic 6 l s O (%o) data from 887B. Cibicides mean offset from Uvigerina is 0.67 ± 0.21 (la); Gyroidinoides is offset from Uvigerina by 0.09 ± 0.14 (tstatistics:T=0.576, P=0.573). Regression parameters and correlation coefficients (r) are shown for both distributions. 9 4 Figure 5.9. Benthic {Gyroidinoides, Uvigerina, and Cibicides spp. (+0.64 %o)) and planktonic (G. bulloides) 8 1 8 0 (%o) vs. depth at Site 887, shown with replicate values. Both benthic and planktonic composites represent mean values of all data at a given depth; gaps in the profiles reflect intervals with insufficient numbers of forams for analysis. Core lithology legend appears on the following page. (continued) 887 benthic 6 1 8 0 (%o) 4.5 4.0 3.5 3.0 Gyroid • Gyroid rep -f Uviger X Uviger rep o Cibicid • Cibicid rep - Composite. 887 G. bulloides 6 1 8 0 (%») 4.5 4 3.5 3 2.5 diatom ooze clay/silt clay/silt (mottled) Figure 5.9 (completed) 96 6 1 8 0 records (Table A2.3; Figures 4.1 and 5.9) indicate shifts for the last glacial-interglacial transition (Aci-int) of 1.24%o and 1.56%o for the planktonic (G. bulloides) and benthic composite records, respectively. The two maximum glacial values were averaged to obtain the planktonic Aci-int value; this would increase to 1.4 l%o if only the heaviest glacial value was selected. Similarly, the benthic Aci-int would increase to 1.62%o. In assessing the benthic amplitude, the average (n=8) "modern" benthic 61 8C» value given by Zahn et al. (1991) for core PAR 87A-10 (3.33%o) is used. 887B data (n=4) yield a similar modern estimate of 3.27%o; however, if the heaviest glacial and lightest Holocene values from 887B are chosen, the benthic AQI-inteaches 2.02%o. Aci-fnt estimates at Site 887 through the Brunhes Chron (Table 5.1) are similar to amplitudes that have been observed at other North Pacific sites (Keigwin, 1995; Zahn et al., 1991a; Zahn et al., 1986) and are lower than concurrent shifts in the Atlantic. This contrast is apparently due to larger temperature fluctuations in the Atlantic relative to the Pacific over glacial-interglacial timescales (Labeyrie et al., 1987; Shackleton et al., 1983). Table 5.1. Glacial-interglacial 6 1 8 0 shifts in core 887B. Transition AG|-mt 6 1 8 0 Benthic Planktonic 2/1 1.56 1.24 6/5 1.66 1.36 8/7 1.16 0.62 10/9 0.92 1.34 12/11 1.93 1.40 16/15 0.91 1.12 97 Major isotopic boundaries are clearly discernible to Stage 18 in the 887B composite benthic 6 1 8 0 profile, which generally reflects the global (SPECMAP) signal (Figures 4.1 and 5.11). Though the G. bulloides S 1 8 0 profile emulates the global pattern to some extent, regional hydrographic effects are evident in the planktonic data through contrasts with the benthic record. Since benthic foraminiferal 6 1 8 0 in the deep Pacific Ocean largely reflects ice volume changes (see Labeyrie et al., 1987), normalization of planktonic data to the benthic signal provides a "local" surface-water isotopic history in the subarctic east Pacific. Contemporary benthic data are not available at the same sample depths as a number of G. bulloides 6 1 8 0 values; thus for the purposes of calculating local 6 1 8 0 (S 1 8 Qocal) 8 , benthic 6 1 8 0 at these points was interpolated by comparing a spline-curve (Paillard, 1994) fit to the 887 benthic composite data with the SPECMAP stack. The 5 1 8Qocal curve represents variations in the surface water isotopic signal which are likely due to surface-layer temperature and/or salinity changes. Thus, past warming and freshening events in the Gulf of Alaska are evinced by negative excursions in 5 1 8 Q o c a i which appear throughout core 887B. These occur independently of global-scale glacial cycling (as suggested by the benthic 6 1 8 0 record) and are often synchronous with the deposition of intervals rich in biogenic material (Figures 5.2 and 5.11). 8 5^80]0cal is the subtraction of benthic composite from coeval planktonic (G. bulloides) values. 98 887 benthic comp. 5 1 8 0 (%o) - o - 887 local (pl.-ben.) S i a O (%>) Revised SPECMAP 6 1 8 0 (%o) Figure 5.11. 887 time series of benthic composite 5 1 8 0 ( % o ) and "local" 6 1 8 0 ( % o ) shown with isotopic boundaries and the SPECMAP stack (modified from Imbrie et al., 1984). Local 5 1 8 0 is the subtraction of benthic composite from coeval planktonic (G. bulloides) values; maxima and minima are constrained by replicate analyses where possible (Fig. 5.9); • represents estimated benthic S ^ O (%o) values (see text) . Numbered S^^iocai excursions are associated with biogenic-rich material and average 0.6.1 %o in magnitude (sd = 0.18%o). Figure 5.11 (completed) 5.33 Carbon Epibenthic Cibicides spp. 6 1 3 C is believed to reflect ambient bottom water 613C,2C02 and hence should record changes in deep water 6 1 3 C (e.g., Zahn et al., 1986). Conversely, Uvigerina spp. is an endobenthic species that lives in the top 2 cm of surface sediments, though living specimens are generally most abundant in the upper centimetre (Corliss, 1991). As a result, 6 1 3 C values of Uvigerina spp. likely reflect ambient pore water 6 1 3C2C02 (McCorkle et al., 1990). Gyroidinoides spp. are thought to inhabit a similar environment and as a result record comparable 6 1 3 C values. This supposition is supported by the data of Zahn et al. (1991), though Gyroidinoides spp. and Uvigerina spp. 6 1 3 C are not statistically similar in core 887B (n=9, A613G=0.17; a=0.18,T=1.76,P=0.04)9. Whereas 6 1 3 C values of DIC produced by carbonate dissolution are similar to those in bottom water DIC, the degradation of organic matter liberates 13C-depletedCC>2 and DOC to pore waters, producing lower 6 1 3 C in pore waters than in overlying bottom waters. Hence, the settling flux of organic matter to the seafloor which is fundamentally related to the DOC and DIC content of the pore waters (Martens et al., 1996) also influences the pore water 6 1 3 C gradient. As a result, variations in Uvigerina spp. and Gyroidinoides spp. 6 1 3 C relative to Cibicides spp. may reflect changes in the 6 i 3 C gradient (6 1 3Cg r) at the sediment-water interface which in turn are controlled by time-varying export production in the surface ocean (McCorkle et al., 1990). Unfortunately, interpretation of the 887B o 1 3 Cg r record is limited by a lack of paired data. Equivocal maxima in 5 1 3 Cg r are apparent in some biogenic ^Note that the Zahn et al. (1991) data set is much larger (n=33). 101 intervals (Figure 5.12, #1, 4, 8, 20); however the data are at present too sparse to draw significant conclusions. The 887B planktonic 6 i 3 C record is consistent with the global trend of more negative 6 1 3 C values in glacial than in interglacial times, with the clear exception of Stage 14 (Figure 5.12; Table 5.2), though positive values in Stage 14 are also apparent in other northeast Pacific cores (Al-Aasm & Bornhold, 1986). Benthic data also match the global trend through Stage 6, though stratigraphic resolution limits the description of glacial-interglacial shifts during earlier times. Nevertheless, observed shifts at stage boundaries are similar to those recorded at other sites (Zahn et al., 1991a; Zahn et al., 1991b). 102 887 benthic 6 1 3 C (%») 887 G. bulloides 5 1 3 C (%<>) 887 o 1 3 C G r a d i e n t • 887 A 6 1 3 C (pi. - ben.) - • -Figure 5.12. Time series of 887 benthic and planktonic 6^c (%o). A8^ 3 C is the difference between paired G.bulloides and Cibicides spp. data; & ^ C Q r a ( j j e n t represents the subtraction of Cibicides spp. from averaged coeval Uvigerina spp. and Gyroidinoides spp. values (note reversed scale). Numbered events correspond with 8^0] o c a ] excursions (Fig. 5.11). Legend on the following page.. (continued) 887 benthic 6 1 3 C (%») -1.5 -1 -0.5 o 887 G. bulloides 6 1 3 C (%o) -e-1.2 -0.8 -0.4 0 0.4 0.5 0 -0.5 887613C 'Gradient 0 0.5 1 887 AS 1 3C (pi. - ben.) 380 400 420 440 460 480 500 520 540 560 580 600 620 640 660 680 700 720 740 760 1.5 887 benthic613C (%») legend: -o— Cibicides spp. -A- Gyroidinoides spp. - • - Uvigerina spp. Figure 5.12 (completed) 104 Table 5.2. Glacial-interglacial 6 1 3 C shifts in core 887B. Transition Aci-int 0 1 3 C Uvigerina Gyroidinoides Cibicides G. bulloides 2/1 0.76 - 0.53 0.57 6/5 0.51 - - 0.38 8/7 - - 1.00 0.35 10/9 - - 0.31 12/1 1 - - 0.97 0.66 16/15 - - - -18/17 - - - . 0.30 A number of G. bulloides 6 1 3 C maxima (mean shift=0.32, cr=0.10) occur which are synchronous with & l 8 Q o c a i minima, intimating 1 3 C enrichments coincident with changing temperature and/or salinity. Increased 6 1 3 C during such episodes could be due to surface 1 2C-depletion produced by the preferential transfer of 1 2 C to deeper water via increased surface (export) production. 6 1 3 C enrichment in G. bulloides in upwelling zones has been observed in previous studies, and as suggested by Kroon & Ganssen (1988), "The fact that G. bulloides is enriched in the upwelling zones suggests that it thrives during the final stage of upwelling when the fixation of 12C in the phytoplankton blooms exceeds the amount of 12C brought into the surface waters". Alternatively, 6 1 3 C excursions may be influenced by thermodynamic fractionation, independent of gas exchange effects which predict change in the opposite direction. Although not as temperature-dependent as oxygen, the isotopic fractionation of carbon is influenced by changing temperatures: A 1 3 C (HCO3- »CaC03) = 0.035%o/°C(Emrich et al., 1970). Thus, equilibrium calcite 6 1 3 C should increase with temperature, though a change of 0.35%o would require an 105 improbably large temperature shift (10°C). Consequently, this effect is unlikely to account for the observed S l 3 C shifts at Site 887. Symbiotic algae, associated with some planktonic foraminiferal species, increase the internal concentration of 13C-depleted metabolic CO2 in foraminiferal cells, thereby altering the 6 1 3 C signal (refs. in Al-Aasm & Bornhold, 1986). G. bulloides are thought to be symbiont-free (refs. in Kroon & Ganssen, 1988), thus symbiont-related effects could not influence the observed 6 1 3 C maxima, except to attenuate the signal. The extent to which G. bulloides accurately records ambient 613Q>;co2 1 S unclear and differences between G. bulloides and equilibrium calcite 6 1 3 C of up to -2%o have been noted (Al-Aasm & Bornhold, 1986)1 0; as a result, the interpretation of foraminiferal 5 1 3 C in terms of seawater 613Q>;c02 is questionable. However, subtraction of epibenthic (Cibicides spp.) from coeval G. bulloides S 1 3 C data does provide a preliminary qualitative record of the benthic-planktonic 6 1 3 C gradient (A13C(pi_ben); Figure 5.12). Though constrained by low resolution, A1 3C(pi-Den) appears to be largely influenced by changes in the planktonic record which accompany S1 8Qocal and productivity events. With the exception of Stage 2, A 1 3 C( p i -b e n ) is generally independent of glacial-interglacial cycles, showing a gradual decrease in the lower Brunhes to a minimum during Stage 16, then increasing through Stage 18. This suggests that nutrient depletion of the deep subarctic Pacific via a thermohaline link to the surface layer has not occurred in response to glacial cycling throughout most of the Brunhes Chron. °This data remains unconfirmed. 106 5.4 dl 3Corg There is a large variation in organic carbon 6 1 3 C (813Corg) in 887B. The maximum amplitude of change exceeds 5%o (mean = -22.8%o; a = 1.2), with maxima and minima at -19.2 and -24.8%o, respectively (Figure 5.6; Table A2.1). These values remain primarily within the interval of -20 and -30%o established for higher-latitude marine phytoplankton by Sackett et al. (1965), implying that organic carbon at this site has originated from a predominately marine source. Using ancillary measurements such as C o r g / N ratios, it is possible to further assess the validity of interpreting bulk sediment 6 1 3Corg as a marine o 1 3Corg signal. The mean C o r g / N weight ratio is 7.56 (a = 1.18), well within the range of 6 to 9 typically seen in marine organic matter (Section 3.4.2.3). Such low and relatively invariant ratios throughout 887B and the virtual lack of correlation between C D rg/N and 5 1 3Co rg data argues against the mixing of terrigenous and marine matter as the cause of 6 1 3Corg variations (Figure 5.13). 107 -25 • 4 5 6 7 8 9 10 11 12 C / N org total Figure 5.13. Scatter plot of 887B 6 1 3 C o r g vs. C o r g / N . Mean 6 1 3 C o r g = -22.8%o, a = 1.2; mean C o r g / N weight ratio is 7.56, o= 1.18. R = 0.008. Peak 6 1 3 Co r g values occur at 69 ka (-19.6%o) and 82 ka (-19.2%o), coeval with the largest opal enrichments seen in the hole. Downhole 5 1 3 C o r g variations closely follow trends in the Si/Al profile and are broadly consistent with changes in C0rg content, though 6^ 3Coi- g infrequently varies with CQrg independently of Si/Al (Figures 5.2 and 5.6). This occurs notably at ~ 110 ka, where carbon concentrations and accumulation rates are maximal (613Corg = -20.4%o,%Corg = 1.03). Temperature-dependent fractionation (+0.28 to 0.35%o/°C) has been proposed as a mechanism by which marine planktonic 5 1 3 o r g values increase with SST (refs. in Fontugne & Calvert, 1992). However, it is unlikely that a 108 direct temperature effect is responsible for the 6 1 3 Co r g shifts in 887B for two reasons: 1) The trends are in the opposite sense; to illustrate, the diatomaceous band at ~ 79 to 92 ka is coeval with a maximum 6 1 3 Co r g excursion of about 4.5%o (from -23.7 to -19.2). Deposition of this layer took place from Stage 5.2 through 5.1 with an abrupt increase in 6 1 3 C o r g occurring on the inception of Stage 5.2, a time of relatively cool SST. The 513Cbrg values decrease during mid Stage 5.1 when SST should have been warmer. 2) As noted in the case of foraminiferal 6 1 3 C , an improbably large temperature shift of > 10°C would be required to account for the excursion of 4.5%o described in (1). Diagenesis can also influence the S 1 3 C o r g record since the rate of decomposition of different sedimentary organic matter components varies, and these are differentially enriched in 1 3 C . Variations in S13CorgOf up to 2.8%o are possible as a result of diagenesis (Fontugne & Calvert, 1992) and would be manifest as a secular trend downhole towards more negative values. Pre-Holocene sediments should thus be isotopically lighter than presumably less-altered core-top samples. Such is the case in 887B; glacial 6 1 3 C o r g values are on average 0.80%o lighter than Holocene values 1 1, though no further trend is evident downcore. However, diagenesis alone cannot explain the isotopic contrasts between lithologies in 887B. 6 1 3 Co r g shifts associated with biogenic strata are consistently positive and > 1.5%o, reaching as high as 4.5%o. For these changes to be the result of diagenesis, rather severe preferential degradation of isotopically light organic matter must have taken place. At present, there 1 1Note that this includes any inherent negative glacial-interglacial shift, predicted by Rau (1994) to be l-2%o for an 80 oatm increase in atmospheric CC>2. 109 are no data available that would support this possibility; hence, only minor deviations in 6 1 3Corg due to diagenetic effects are assumed here. Field studies have shown that algae such as diatoms and large filamentous cyanobacteria can be replete in 1 3 C ; in blooms where diatoms are important, 6 1 3 C P O M values are commonly -18.5 to -21.5%o (Fry & Wainright, 1991). The physiological basis of such algal enrichment is not well understood. In addition to decreased CC>2(aq) concentrations, diatom 6 1 3 C could be influenced by: 1) 1 3 C - r i c h bicarbonate fixation, which in diatoms may accompany rapid nitrate assimilation (Fry & Wainright, 1991); and 2) reduced fractionation in response to high growth rate (Section 3.4.2.4). Though opal-rich strata in 887B are characterized by large diatoms (e.g., Actinocyclus curvatulus), the role of aforementioned metabolic effects in determining bulk sedimentary 6 1 3 Co r g in such intervals is difficult to constrain. 110 5 . 5 Environmental Changes and Palaeoproductivity Throughout this section, the focus shall be on the interpretation of results over the past 200 ka. Trends observed in this interval are for the most part representative of variations in the lower core where both isotopic and temporal resolution are more limited. Unless otherwise indicated, Figures 5.14 and 5.15 illustrate the relationships between sediment properties which are discussed here. I l l Age (ka) Figure 5.14. 887B geochemical, isotopic records, and preliminary alkenone temperature records to 200 ka. Italicized numbers correspond with & 1 8 0 l o c a l excursions; numbered sections at the far right represent carbonate stages of Farrell and Prell (1989). Revised SPECMAP curve is shown for reference. 112 113 5.5.1 Carbon Dioxide 5.5.1.1 Calculation of [C02(aq)] from the 6 1 3 Co r g Record Previous work has shown that sedimentary S 1 3 Co r g values largely reflect a negative linear response to [C02(aq)] in overlying surface waters (e.g., Rau, 1994). Hence, with the caveats described in Sections 3.4 and 5.4 in mind, the S 1 3 Co r g record in 887B can be utilized as a direct measure of past variations in seawater ZCD2 concentrations. In order to make this interpretation, the 887B S 1 3 Co r g profile has been corrected for changes in surface water 6 1 3 C by subtracting variations in G. bulloides 6 1 3 C normalized to the modern 6 1 3 C of DIC in the Gulf of Alaska (0.801; C.S. Wong, personal communication). Since this correction requires the use of paired data, 6 1 3 Co r g values at G. bulloides sample depths are calculated based on a spline fit. Corrected 5 l 3 Cor g data are then converted to [C02(aq)] values using the relationship reported by Rau (1994): [cO2 (aq) ] ± Z 0 [ x M =-17.9-1.42 x 6 1 3 C o r g Estimated [CO2(aq)] ranges from 8.9 to 15.1 uM; results are shown in Figure 5.16. [C02(aq)] minima are consistently synchronous with diatomaceous intervals throughout the core, intimating an association between export productivity pulses and low dissolved surface-C02 concentrations. Due to the modest number of 5 1 3 Co r g sample points in the lower core, interpretation of the [C02(aq)] record is limited beyond ~ 200 ka. 114 Figure 5.16. 887B S 1 3 C o r g and estimated dissolved CO 2 concentrations ([C02(aq)j) using the relationship derived by Rau (1994). Corrected o-^c values represent the subtraction from measured 6 1 3 C o r g . of G. bulloides 6 1 3 C variations relative to modern surface ocean o 1 3 C D I C . Si/Al profile shown for comparison. (continued) Figure 5.16 (completed) 5.5.1.2 PC02 The conversion of reconstructed [CC»2(aq)] to provide a measure of past surface ocean PCO2 is easily done using Henry's law: PC0 2 * a =[C0 2(aq)] where a denotes the solubility constant as derived by Weiss (1974): In a = A, + A^lOO/r) + A, ln(r/100) + S [fl, + fl^T/lOO) + fi3(r/ioo)2] a is expressed in moles/l-atm, T is the absolute temperature, S is the salinity, and A's and B's are given constants. Variations in salinity of several %o produce relatively minor changes in CO2; of more significance to PCO2 calculations is temperature. In any case, past temperature and salinity estimates at Site 887 are problematic. However, dinoflagellate transfer function (TF) data from core PAR87A-10 have been used to reconstruct surface water properties through the past 27,000 calendar years (A. deVernal, unpublished data; Section 5.5.2.1). TF-derived August temperature and salinity values were used to calculate a, which was then applied to coeval 887B [C02(aq)] data to yield PCO 2. Downhole PCO2 estimates must be interpreted with care since uncertainties exist not only in the measurement and spline fit of 61 3Corg values but also in the calibration of [C02(aq)l to 613Corg, and in the transfer function-based data. Furthermore, the presence of minor amounts of terrestrial organic matter cannot be completely discounted. Excluding error in the TF data, the resultant PCO2 error is ~50 patm. PCO2 results for 887B are shown in Figure 5.17 and indicate that calculated Holocene values are comparable to those observed during between 1973 and 1978 in the Gulf of Alaska. Calculated (887B) core top PC0 2 is 311 uatm, whereas observations by (Wong, 1991) at Station P suggest that the modern (circa mid-1970's) annual range of surface ocean PCO2 is from 290 to 340 patm. 117 Weighted PCO2 monthly means range from 306 uatm (November) to 331 patm (June) with high values occurring primarily from May to July, and low values in late summer and fall; August mean is 313 uatm 3 2. Wong (1991) suggests that surface waters in the Station P region today are predominately a weak CO2 sink and occasionally a weak to neutral summer source. Direct comparisons between the 887B PCO2 (PC02(Sw)) record and the ice-core record of atmospheric pCC>2 (pC02(atm)) are limited by the extent of equilibrium between atmospheric and sea surface CO 2. At present, such CO 2 equilibria are uncommon, and regional ocean-atmosphere pC02 differences can reach >100 uatm (e.g., Tans et al., 1990). Nonetheless, the glacial-interglacial PC02(S W) shift apparent in 887B is on the same order as that observed in ice cores (~ 80 uatm; Figure 5.17). A notable discrepancy between the two records occurs from 14 to 10 ka, concurrent with the deposition of a biogenic-rich (diatom, carbonate) band. Across this interval, pC02(atm) rises steadily by ~ 73 uatm while PC02(Sw) increases by 29 uatm to 13.3 ka at which time it begins to decrease, reaching a minimum of 267 uatm at 10.3 ka for a net gain of only 5 uatm. It seems reasonable that the disparity between the two curves may be caused by increased phytoplankton production on the interglacial inception. G. bulloides 5 1 3 C values appear to support this scenario (Figure 5.14), though decisive interpretation is limited by the data resolution. Seasonal changes in sea surface CO2 levels at Station P have been attributed to a balance between productivity and mixed-layer depth (Wong, 1991). In April or May, the mixed layer shoals from the permanent halocline at 100-200 m to a summer depth of 25 m or less; the thickness of this layer determines the amount of CO2 available for photosynthesis and ocean-atmosphere exchange. Removal of seawater CO 2 occurs in response to spring 1 2PCO? values based on observations from 1973 to 1978. 119 and summer phytoplankton production and the concurrent sinking of organic matter below the mixed layer. In the winter months, PC02(Sw) rises as photosynthesis decreases, the mixed layer deepens and mixes with sub-surface waters replete in CO2, and cooler surface waters take up atmospheric CO2. During the period from 14 to 10 ka, it is possible that the summer PC02(S W) drawdown was enhanced due to CO 2 removal via increased new production. This mechanism could account for the "stalling" of the 887B PCO2 signal relative to atmospheric PCO2 on deglaciation. Such intensification of the summer CO2 sink has been observed in the subarctic Atlantic, Bering Sea, and western subarctic Pacific regions (refs. in Wong, 1991). 5.5.2 Evidence of Temperature and Salinity Variations 5.5.2.1 Relationships Between 6 l s O and Transfer Function Data Specific plankton species or assemblages are characteristic of distinct oceanographic settings; hence, varying environmental conditions are manifested by changes in those species. As a result, micropaleontological transfer functions can be used to obtain an independent record of changing sea-surface conditions. Comparison of measured 5 1 8 Q O C al with transfer function data is thus useful in estimating the relative effects of temperature, salinity, and sea-ice extent on down-core planktonic 6 1 8 0 values. In order to make this comparison, down-core foraminiferal 6 1 8 0 (6C) values were calculated based on dinoflagellate transfer function (TF) data from core PAR 87A-10 over the last 27,000 calendar years (A. de Vernal, unpublished data; Table A2.4). TF-derived salinity and temperature values were used to calculate seawater 6 1 8 0(6 W ; Zahn et al., 1991) and 6C (Shackleton, 1974). Once corrected 120 for temporal changes in mean seawater 6 1 8 0 due to continental ice-volume variations (Duplessy et al., 1988), 6C was subtracted from modern calculated values to yield the local (A6C) signal. The measured A6C record is based on pooled G. bulloides data from cores PAR87A-01, 02, and 10 (Zahn et al., 1991a), and from 887B. As with calculated A6C, G. bulloides 6 1 8 0 was subtracted from mean Sw (Duplessy et al., 1988) and modern (core-top) values to obtain measured A6C. There is some similarity between measured and calculated A6C profiles (Figure 5.18). R values for measured vs. predicted February and August A6C are 0.76 and 0.77, respectively; moreover, cross-spectral analysis suggests that measured and calculated records from both February and August are consistently coherent at the 95% confidence level (Figure 5.19). 121 C o E w > o o 0 o CC CD CO CO as to O o cb CO 3 < CD 4—' co (0 i CD CL E <D I— CU _3 > CU O P i CJ c O < •a cu >>.s BO C 03 P H P 0 OS a *o x: 3 C J P 03 01 P i CD CU _co p . >> a p * t3 P » fl " cu cu -a 3 u 15 o -a c as X ) cu p 3 to o3 cu CO cu °p cu co CU < 0 0 OH < a. cu p 3 60 o3 o3 (J CO o S P I o •a cu p i CU cu B b X ! p O •a cu CO 03 03 c j •d a o3 •a cu p 3 t o 03 cu DO £ ^ "o c c o O U fl u -fl cu o a cu T 3 g 6 C "C C CU oo — .2 co ' o CU to fl b <y o R <u CU x i & c C J CU P cu X tS x! s fl -03 . f l p b cu o <a CJ CO fl pa g P ! oo g xi cu co —' 03 ^ - ° TJ o cu a oi CU <-> M c O 03 < .y -a P CU P p CU 03 S 3 O u C CU X 5 •d c — o3 CU fl .b 03 p < CU pi 03 O cu . fl b 03 03 ^ -o ^ CJ P fl CO CU i >> cu £ u £ £ s £ o » *^  Cp . . .—i pi o O _o _y o o3 co _. 0) fl p» ^ 03 CU ^ 03 00 fl &0 =3 .5 x c j 03 CU P y 1 / 5 C 03 &0 co X co ' Q V C ^ o3 CO S 2 03 CU - I 122 1 0.9 8 0.8 c <L> v. 0.6 0.5 0 2 4 6 8 10 12 Frequency (cycles per metre) Figure 5.19. Coherency estimates for measured vs. calculated 5^0, core PAR87A-10. On the interglacial inception, calculated August 6 l s O values are more negative than measured 6 l s O . This phenomenon may be due to low salinities associated with seasonal melting of sea ice. To illustrate, if -0.50%o of the 0.73%o decrease in measured A6C from 14.7-13.8 ka is attributed to the synchronous (TF) 2.0°C warming, then the measured A6C signal accounts for less than 1/4 of the estimated salinity shift of 4 during that time. Moreover, the measured A6C excursion at ~18.3 ka reflects only 1.0 of a concurrent 2.8 salinity decrease. The glacial offset between measured and calculated 6^o values may be the result of changing foraminiferal calcification depths. Maximal 123 concentrations of G. bulloides are today found at 50-75m depth (refs. in Al-Aasm & Bornhold, 1986), though Zahn et al. (1991) suggest that the species may calcify at depths up to 150m. During glacial times, calcification may have occurred at shallower depths in response to cooler surface temperatures. Thus, the discrepancies between measured and calculated A5C may to some degree reflect temperature and possibly salinity gradients between G. bulloides calcification depths and dinoflagellate environments. 5.5.2.2 Alkenone-Derived Temperatures Alkenones 1 3 are a specific class of lipids produced by prymnesiophytes, of which the coccolithophorid Emiliania huxleyi is a characteristic and dominant species in the subarctic Pacific (Parsons & Lalli, 1988). The nanoplankton size fraction (2 to 20 um), composed largely of a variety of small flagellates including E. huxleyi, constitutes ~ 80% of phytoplankton biomass throughout the year in this area; hence the alkenone temperature signal may represent an integrated or mean annual temperature. Some evidence suggests that coccoid forms dominate the smaller flagellate size class (2 to 5 urn) in August, though at that time their distribution may extend through the surface water column over 120m (Booth et al., 1988). Thus alkenone-derived temperatures may represent both seasonal and depth integrations. A comparison of TF and preliminary 887B alkenone-derived temperatures (G. Cowie; unpublished data) is shown in Figure 5.20. Long chain unsaturated (C37-C39) ketones 124 Figure 5.20. Transfer function (TF) and alkenone-derived sea surface temperatures through the past 30 ky. "TF mean" represents the average of February and August transfer function estimates; alkenone temperatures are based on a field-derived Uk-37 and SST calibration. Holocene alkenone and mean TF-based temperature estimates of ~ 9°C are in good agreement and are similar to modern mean annual temperatures of 8-9°C at Station P (Tabata & Peart, 1985). This suggests that both alkenone and mean TF data may provide reliable downcore estimates of mean annual temperature changes at Site 887. Discrepancies between Holocene August TF and alkenone temperatures may represent vertical stratification between E. huxleyi and larger dinoflagellates, though this has not been noted in field 125 studies (e.g., Booth et al., 1988). In any case, the relationship between TF and alkenone data has altered through the last glacial-interglacial transition; the TF-based temperature change across the glacial-interglacial transition (AT; taking average Holocene and glacial values) is 7.4°C, while the alkenone AT is 3.8°C, more closely reflecting the CLIMAP (1981) estimate of 3°C. Moreover, alkenone temperatures increase by only 0.5°C from ~ 17 to 10 ka, in contrast to TF estimates which rise by up to 6°C through this interval. If we assume that the alkenone data reflect relatively constant SSTs during this period, then salinity shifts of almost 2 may be responsible for the observed 6 1 8 0 excursions at ~ 14 and 18 ka noted above (Figure 5.18). Additional alkenone-derived temperatures are shown in Figure 5.14. Temperatures at 6^0 events 4, 7, 8, and 10 are low relative to Holocene values, ranging from 4.7°C in Stage 6.5 to 7.5°C in Stage 5.4, whereas local 6^0 values are consistently lighter than core top by > 0.7%o. Although the alkenone data are at present sparse, the implication is that local 6^0 excursions are dominated by reduced salinity. 126 5.5.2.3 C. davisiana Records Morley & Hays (1983) have suggested that the predominance of the radiolarian C. davisiana in glacial sediments is due to conditions which were similar to those prevailing in the Sea of Okhotsk at present. Modern surface waters in the Sea of Okhotsk are characterized by: 1) low surface salinities resulting from seasonal sea ice melt, 2) a pronounced temperature minimum near the base of the low salinity layer, and 3) relatively invariant temperatures and salinities below the temperature minimum. Hence, if local excursions at Site 887 are the result of abrupt decreases in salinity one may expect a close correlation between C. davisiana abundance and light 6^0 events. However, such is not the case in 887B; in fact an inverse association is apparent between %C. davisiana (Morley et al., 1995) and local 6^0 (Figure 4.3). Modern conditions in the Sea of Okhotsk produce strong seasonal variability in the abundance of shallow-dwelling organisms (Kun, 1951), whereas a relatively stable environment exists for sub-surface fauna at depths greater than ~ 200 m. C. davisiana is restricted to depths of > 200 m, and highest abundances of this species occur in sediments below waters with the coldest sub-surface temperature minimum (Morley and Hays, 1983). In contrast, G. bulloides inhabits the near-surface layer (Sautter & Thunell, 1989), although this species may calcify at depths up to 150m (Zahn et al., 1991a). Thus, the negative relationship between %C. davisiana and local 5^0 values may be explained by environmental and/or ecological changes which were not conducive to high abundance levels of C. davisiana, yet were somehow related to decreasing surface water salinities. One such scenario in the Gulf of Alaska may have involved an increase in the depth of winter mixing which would destabilize temperatures and salinities at depths 127 approaching 200m and cause a drop in C. davisiana abundance. During the ensuing spring, decreasing salinities in the surface layer due to the influx of less saline water could produce the observed 5 l 8 q o c a l excursions observed in G. bulloides. 5.5.2.4 Mechanisms of Change Consideration of both transfer function and alkenone records implies that some fraction of observed negative S^O excursions are influenced by salinity changes. This suggests that in addition to seasonal fluctuation of sea ice margins, such events also reflect freshwater supply from coastal zones or iceberg meltwater. Hence, increased sea ice cover may have occurred at times in response to low surface layer salinities in the Gulf of Alaska. Evidence of low salinity on deglaciation has been reported previously in the western subarctic Pacific (Keigwin et al., 1992) and in the Bering Sea (Sancetta et al., 1985), though the origin of low-salinity water at these locations remains an enigma. The Alaska Gyre probably remained in existence throughout the Pleistocene, regardless of climatic change. This is indicated in the detrital (ice-rafted) fraction of core samples by the absence of rock types from places other than the adjacent coast. Most of these come from slate and greywacke outcrops along the central and western coasts of Alaska. The remaining rocks (granite, granodiorite, mica, metavolcanics) are found on smaller outcrops on Kodiak Island, in Prince William Sound, and in the Chugach - St. Elias Mountains. Most lithics fit the description of ablation debris on the Malaspina Glacier, the largest piedmont glacier flowing into the Gulf of Alaska (von Huene et al., 1976). Thus, IRD has been carried from the Chugach - St. Elias Mountains by 128 glaciers and then rafted by icebergs. The subarctic east Pacific is presently devoid of iceberg activity, and substantial sea-ice does not occur south of the Aleutian Islands (Kent et al., 1971). However, glaciers currently extend to tidewater in a number of bays and fjords, and at the location of La Perouse Glacier on the open coast (Plafker & Berg, 1994). During the last glacial maximum (LGM), a glacier complex composed of alpine glaciers, island ice caps, and piedmont lobes, which formed an extension of the Cordilleran Glacier Complex, terminated in ice cliffs which stood 50m above sea level along much of the outer continental shelf edge in the Gulf of Alaska (Mann & Peteet, 1994). Major ice retreat in the Gulf of Alaska probably began at 17-14 ka; glaciers had retreated from the continental shelf by 12-11 ka, and by 10.5 ka had reached positions comparable to those of present day (Hamilton, 1994). As ice retreat took place, sea level rose eustatically, though was still 50 to 75 m below present level. During this time, freshwater runoff to the Gulf of Alaska increased as rivers cut across the shelf, carrying glacially-derived fluvial sediments. Subsequent off-shelf transport of freshwater plumes over Site 887 could have occurred via circulation of the Gyre following routes similar to those taken by IRD, thereby providing a less saline source of water. TF data (Figure 5.18c) indicate that a substantial drop in salinity occurred from ~ 17 to 10 ka, corresponding to the interval of ice retreat from the continental shelf. Over approximately the same period, annual duration of sea ice cover decreased from 8 to 0 months intimating the withdrawal of a sea ice margin. Seasonal meltback along the margin of this zone would also contribute to lowered surface salinity, but would have little effect on 6 l s O since isotopic fractionation during freezing of seawater is of minor significance. 129 Alternatively, lowered salinity may also reflect changes in the evaporation-precipitation balance in the subarctic Pacific, though intensification of the Aleutian low which would act to enhance precipitation in the North Pacific was most pronounced during the LGM (Kutzbach et al., 1993). On deglaciation, pressure patterns were less strongly affected by the presence of ice sheets, with sea-level pressure and annual precipitation approaching modern values in the Gulf of Alaska. 5.5.2.5 High Frequency Climate Variability As mentioned previously, large glaciers and ice fields in the coastal Alaskan ranges are at present nearly within tidal range; hence small glacial advances could produce icebergs in the open ocean. Kotilainen & Shackleton (1995) present limited evidence for millennial-scale climate variability in the western subarctic Pacific, suggesting that periods of low atmospheric temperature were accompanied by discharge of icebergs into the North Pacific. Such rapid climate variability has also been noted in more southerly latitudes of the northeast Pacific Ocean (e.g., Thunnell & Mortyn, 1995), implying that these events may be driven by regional climate change. If, as suggested by Kotilainen & Shackleton (1995), major fluctuations in air temperature over Greenland and the North Pacific are synchronous, then the 5I80 record of sea surface conditions at Site 887 may reflect concurrent variations in Greenland ice cores. As shown in Figure 5.15, precise correlation between 887B records and GRIP 6 1 8 0 is equivocal, owing to the lower resolution of 887B data at this time. Nevertheless, surface water 6 l s O excursions appear to occur during periods of warmer air temperature identified in the GRIP 6 1 8 0 record. Moreover, such events also coincide with high D. seminae abundance in core PAR 87A-01 (C. 130 Sancetta, unpublished data) which intimates relatively warm interglacial conditions14. Thus, surface ocean conditions may be influenced by climate variability which is much more rapid in the Gulf of Alaska than that predicted by orbital forcing. Such changes may be apparent as light 6 1 8 Q O C a l events in the 887B record. 1 4 This species is today characteristic of the central portion of the Alaskan Gyre (refs. in Parsons & Lalli, 1988); however, its palaeoecologic significance is unclear. 131 5.5.3 Factors Affecting Surface Ocean Productivity In the modern Gulf of Alaska, episodic increases in new production are linked with a number of factors including upwelling, high SST, lowered surface salinity, a deeper spring mixed layer, and high solar radiation in spring and summer (e.g., Wong et al., 1995a). Iron limitation may also play a significant role (La Roche et al., 1996). This section examines the evidence for temporal changes in each of these parameters and their effects on the palaeo-productivity record. 5.5.3.1 Upwelling Episodic increases in surface primary production at Site 887 are indicated by down core variations in Si/Al, Ba/Al, %Co rg, and 6 1 3 C data. Productivity events are generally coincident with surface water 6 1 8 Q 0 cal minima, suggesting that enhanced primary production is frequently associated with temperature and/or salinity changes, though the relative dominance of each remains an open question. Both transfer function and alkenone data suggest that temperatures during these events were lower than at present; indeed temperature rise on the most recent interglacial inception appears to have "stalled" for a significant fraction of the interval from 14-10 ka, concurrent with a salinity minimum and increased biogenic deposition. The inference of colder surface water temperatures (presumably in summer) relative to today could signal increased upwelling of deep waters and as a result, higher surface concentrations of inorganic nutrients such as phosphate, nitrate, and silicate. As noted previously (Section 3.2), production rates high enough to allow the accumulation of biogenic silica in sediments 132 occur only in upwelling regions where a continuous supply of nutrients (notably silica) is available. Records from the high-latitude Atlantic show that increased discharges of ice through the last glacial cycle have been associated with low sea surface temperatures (Bond et al., 1992). Pulses of ice rafting into the North Atlantic (Heinrich Events) have occurred synchronously with Dansgaard-Oeschger cooling cycles which are evident in the GRIP ice-core record through the past 250 ky (Bond et al., 1993; Dansgaard et al., 1993). The large amplitudes (ASST up to 5°C) and sharp boundaries which characterize Dansgaard-Oeschger events are independent of global ice volume; for example, events occur on the last degradation, during the last glacial maximum, during Stage 3, and on the Stage 4 inception (Bond et al., 1993). These cycles are thought to be closely linked with North Atlantic Deep Water (NADW) production. Low SSTs associated with Dansgaard-Oeschger oscillations are generally synchronous with episodic reduction in NADW formation and a decrease in the strength of thermohaline circulation throughout the last 200 ka (e.g., Oppo & Lehman, 1995). During such episodes, a corresponding reduction in the rate of deep water upwelling in the subarctic Pacific probably occurs. Following the cessation of iceberg flux into the North Atlantic a warming phase begins, thermohaline circulation increases, and subarctic Pacific upwelling is strengthened. Such warming may have occurred in Stage 5.4, concurrent with deposition of the biogenic-rich interval in core 887B centered at ~ 107 ka. At that time, ice volume had reached ~ 50% of the full glacial maximum, yet significant atmospheric warming is apparent in the GRIP record, concomitant with the high production of NADW documented by Duplessy & Shackleton (1985). This event is also marked in 887B by a negative S18Qocal shift and 133 alkenone-derived SST of 7.5°C. The 6 1 8 Q o c a l excursion reaches l%o lighter than core top values; hence, even if we assume that "full glacial" alkenone temperatures of ~ 5°C preceded this event, then a salinity decrease (AS) of ~ 1 must have occurred, corresponding to a AS 1 8 0 of ca. - 0.4 %o. 5.5.3.2 Surface Water Properties Increased GRAPE density at Site 882 has been directly linked with high IRD abundance and by inference, to high IRD flux (Kotilainen & Shackle ton, 1995). Preliminary IRD abundance estimates in core 887B suggest that a similar association exists at Site 887 (Figure 5.1). The high frequency of IRD abundance variations in 887B suggests that IRD flux at Site 887 varies at "sub-orbital" timescales and could possibly reflect climate forcing which is in some way linked to Dansgaard-Oeschger oscillations. However, in the absence of a precise chronology that can be matched to the GRIP record, no attempt to map specific correlations is made here. In a broader sense, it does appear that episodic iceberg discharge into the Gulf of Alaska alternates with periods of high surface productivity associated with relatively warm surface water temperatures (albeit possibly lower than at present) and low salinities. If such productivity events are connected with atmospheric warming it is likely that during such intervals, which generally occur at during times of increased global ice volume relative to the present, glaciation was more extensive in Alaska than today. There is evidence of ice growth in both the Brooks and Alaska Ranges beginning in Stage 5.4, possibly due to warm SSTs in the North Pacific during Stage 5 which may have provided a source of water vapour, and thus snow. Hence, glaciation in Alaska may begin earlier in glacial-interglacial cycles than in other regions (J. Brigham-Grette, personal communication; Hamilton, 1994). Retreat 134 of glacial ice beyond tidewater produced by warmer atmospheric temperatures could result in the release of large volumes of meltwater into the Gulf of Alaska from both coastal glaciers and inland sources such as glacial lakes, thereby lowering surface salinity. To illustrate, one of the major glacial lakes in the Copper River Basin (Lake Atna; Chugach Range) started to form before 59 ka. Lake levels fluctuated through to the last deglaciation when complete drainage took place; evidence indicates that this lake was contained by ice dams and reached an areal extent of more than 5200 km 2 and a depth of ~ 270 m (Williams, 1989). Assuming a 6 1 8 0 value of -35%o for glacial ice (Dansgaard et al., 1993), meltwater from this lake would be sufficient to cause a salinity anomaly of 1 in a surface layer 100 m thick over an area of ~ 1.2 x 106 km 2 , thus covering much of the Gulf of Alaska. Planktonic foram abundance peaks correspond with 5 1 8Q 0cal excursions, though the distribution of species is variable between events. Counts of G. bulloides and N. pachyderma suggest that deposition of diatomaceous layers was associated with a relatively warm and stratified productive season, whereas colder conditions dominated during deposition of more carbonate-rich strata (Figure 5.8). IRD appears to be more common in the N. pachyderma-dominated intervals which occur in latter glacial stages (with the exception of Stage 2). Thus in glacial times the presence of icebergs may also influence surface water stratification, though this interpretation is contingent on a more comprehensive analysis of IRD at Site 887. During the interval of enhanced productivity from ~ 14 to 10 ka, model results indicate that both January and July SSTs were colder by about 2°C than at present (Kutzbach et al., 1993). Wind speeds were also higher than today as westerly flow increased with the reduced height and extent of the northern 135 ice sheets. High wind speeds would result in increased vertical mixing and a deeper mixed layer during winter which, along with increased upwelling, would combine to renew surface layer nutrients. Following deep winter mixing, decreasing salinities in spring would produce a shallower mixed layer, confining phytoplankton nearer the surface, thereby increasing available light and growth rate. Use of the aforementioned productivity event as an archetype for other apparently similar events downcore is somewhat tenuous. However, it does seem reasonable to assume that global ice volumes (as indicated by the SPECMAP record) and hence wind fields were similar, as in the case of S 1 8 Q 0 cal event 7 during Stage 5.4. Increased incident solar radiation has been shown to increase phytoplankton growth rate in at Station P, and is especially important in late spring through early summer since most net phytoplankton productivity occurs during this period (Wong et al., 1995a). Note that this conclusion is based on interannual studies in which atmospheric conditions were found to control both mixed layer depth and variations in spring incident solar radiation. Hence, a connection between productivity and fluctuations in solar radiation controlled by orbital parameters over much longer timescales may not be as straightforward. Nevertheless, intervals of enhanced productivity at Site 887 generally occur during periods where solar radiation is equal to or greater than at present (Figure 5.15). 136 5.5.3.3 Iron As described in Section 2.2.1, the persistence of major nutrients in surface waters of the subarctic Pacific may be due to a limited supply of Fe which in turn limits total new production. Consequently, in addition to physical parameters such as upwelling and surface water conditions, export productivity at Site 887 may be fundamentally controlled by the mechanism responsible for variations in Fe supply. Fe is primarily supplied to the ocean via long range transport of Fe-rich atmospheric dust derived from arid terrestrial regions (Duce, 1986). Episodic increases in aeolian influx to the North Pacific from Asia have been linked previously to increases in primary production of more than 60% (Young et al., 1991); moreover, the observed increases in production could be supported by the dissolution of only 10% of the Fe contained in aeolian particles. Young et al. (1991) observed that: "when the iron-rich dust first arrived, the primary producers were iron-limited, and subsequently, primary productivity was stimulated to the extent possible before it was probably limited by other nutrients". Thus, when a "dust event" occurs, adequate Fe appears to be available to the phytoplankton and productivity rates increase. In the past, episodic shifts to primary production at Site 887 dominated by large, rapidly-settling phytoplankton cells may have been stimulated by such events and preserved as biogenic-rich intervals in the sediment record. High-frequency variability in the transport of Asian dust to the Pacific has been hypothesized based on grain-size data from Chinese loess deposits of the last glacial period (Porter & Zhisheng, 1995). Porter & Zhisheng (1995) correlate high dust flux with times of atmospheric cooling (evinced by the GRIP record); however, within the temporal error limits of their study which approach ± 5 kyr, times of increased dust flux to the Pacific correspond to 137 61 8Qocal events 1 through 5 at Site 887. These events also correspond to intervals of high productivity, notably the productivity maximum centered on the 5/4 transition which matches a dust event at 71.5 ka in the Porter and Zhisheng (1995) record. Resolution of the temporal inconsistency between dust events during atmospheric cooling in China and productivity events at Site 887 which appear to occur during intervals of atmospheric warming may lie not only in discrepancies between age models, but also in the duration of events at both locations. For example, although grain size maxima culminate at times corresponding with atmospheric cooling events, the dust flux itself may remain high and extend temporally beyond the boundaries of such events due to continuing dryness on the Asian continent. During such times strong westerlies would provide a mechanism of transport across the Pacific. Such a description of enhanced Fe transport begs the question as to why primary productivity would not be stimulated to the same or even a greater degree during cold intervals at Site 887. It is possible that despite the availability of iron, other conditions requisite to intensified primary productivity were absent in the Gulf of Alaska during glacial times. Such conditions could include surface water properties discussed in the preceding section. It has previously been postulated that the release of nutrient-rich meltwater from numerous icebergs may have supported increased productivity in the North Pacific during the last glacial interval relative to today (Sancetta, 1992). However, in the absence of evidence for the presence of icebergs associated with productivity events seen in Hole 887, increased runoff may have provided an alternate source of iron enrichment in the Gulf of Alaska. This could have occurred as meltwater plumes extended into the Gulf of Alaska from coastal sources carrying glacial contact and fluvially-derived 138 sediments rich in iron. Fe concentrations in surface waters presently increase with proximity to the Alaskan Coast, coincident with nitrate depletion and increased productivity (Martin et al., 1989; Section 2.2.1). This attests to the availability of Fe released from particles as well as by diagenetic remobilization in shelf sediments. 139 6. Summary and Conclusions Downcore geochemical data from Site 887 indicate that episodic increases in primary production paralleled by rapid settling of biogenic matter throughout the Brunhes Chron are represented by diatom ooze and possibly carbonate-rich strata. The close relationship between Si/Al ratios and diatomaceous bands suggests that high Si /Al values reflect the increased deposition of biogenic silica; moreover, these intervals are characterized by an abundance of large (> 150 um) diatoms. An association exists between CaCC»3 and Si/Al, though CaCG*3 abundance variations reflect the balance between production and dissolution of calcite. Mass accumulation rates (MAR) suggest that the CaC03 content is not controlled by non-carbonate dilution. Peak Ba enrichments are closely linked with biogenic silica and carbonate deposition and are broadly associated with organic carbon enrichments. Significant CQrg enrichment does not consistently occur in association with Si/Al, %CaCC>3, or Ba/Al maxima; however, ancillary factors such as ecological changes combined with variations in bulk sedimentation rate and opal preservation may also affect the organic carbon content. If biogenic-rich intervals indeed reflect increased supply of biogenic matter to the sediments, then an increase in C o r g MAR would be expected during such events; i.e. although C o r g content is generally low in diatomaceous intervals, it is probable that the rate of carbon accumulation during the time of deposition was rapid. This supposition is supported by 2 3 0 T h data which indicate a doubling of sedimentation rate in the two large diatom bands occurring in Stage 5.2 and on the 5/4 transition (Figure 4.4). Episodic increases in the deposition of biogenic matter are also evinced by the occurrence of significant downcore Mn enrichments in biogenic-rich 140 intervals. This suggests that during deposition, a shoaling of the sediment redoxcline occurred in response to O2 depletion at the sediment-water interface due to degradation of an increased flux of organic matter to the seafloor. The occurrence of such redox changes is also indicated by trace metal enrichments concurrent with diatom ooze strata through Stage 5 in core PAR87-10 (McDonald, 1993). Episodes of high surface productivity are associated with a combination of relatively warm surface water temperatures (albeit possibly lower than at present) and low salinities. Such conditions appear to alternate with colder circumstances which are conducive to the discharge of icebergs into the Gulf of Alaska. Regional surface water temperature and salinity changes are evident in o 1 8 Q 0 caI variations, in transfer function data, and in alkenone temperature data. Such changes may reflect variable freshwater supply from coastal zones coupled with seasonally fluctuating ice margins. The seeming alacrity and amplitude of these signals suggests that significant regional climate variability in the Gulf of Alaska has been superposed on glacial and interglacial periods throughout the Brunhes Chron. This is consistent with an increasing body of evidence which indicates that large and abrupt climate change has occurred on sub-Milankovitch, millennial, and smaller timescales (e.g., Dansgaard et al., 1993; Keigwin and Jones, 1994; Thunnell and Mortyn, 1995). It is possible that S 1 8 0 depletion events are salinity-dominated due to seasonal meltback of ice which was likely present on the coast even during cooler interglacial substages (e.g., Stages 5.4, 5.2). This meltback may have been due to warmer air temperatures (indicated by the GRIP record) which would have been concurrent with increased NADW production. Hence, increased upwelling and slightly lower surface water temperatures could have 141 occurred at Site 887 (relative to today) during these events. If indeed a link exists between Greenland ice-core and surface-ocean 6 1 8 0 in the Gulf of Alaska, and hence between Atlantic and North Pacific climates, then ice-core and marine sediment records should provide similar records of climate change. Although tenuous correlations can be made between Greenland ice-core (GRIP) and Site 887 profiles, higher resolution records and associated development of detailed age models will be necessary to enable robust correlation between ice-core and North Pacific-sediment profiles. The accumulations of biogenic material that are interpreted here as reflecting episodic increases in productivity were sufficient to draw down surface ocean CO2, as illustrated by significant increases in 6 1 3 C o r g in diatomaceous intervals. Recognizing that terrestrial contamination of organic matter at this site is minimal, calculation of [C02(aq)l from the 6 1 3Co rg record shows that [CC>2(aq)] minima are consistently synchronous with diatomaceous intervals throughout the core; this intimates a strong association between productivity pulses and low dissolved surface layer CO2 concentrations. Moreover, conversion of [C02(aq)J toPC02 values in the upper core suggests thatC02 removal due to increased new production was responsible for PC02(Sw) drawdown during the last deglaciation. Thus, in contrast to the near neutral ocean-atmospheric PCO2 (APC02oc-atm) gradient which exists at present, an episodic sink for atmospheric CO 2 apparently existed in the Gulf of Alaska in the past. Although it has not been possible in this thesis to provide a definitive explanation for the root cause(s) of the productivity events that punctuate the Late Quaternary palaeoceanography of the Gulf of Alaska, it is clear that regional influences on the surface ocean played a strong role. One other potential influence that remains somewhat enigmatic is the role of iron. 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At the same time, the core was sampled at 10 cm intervals wherever possible. A total of 418 sample points were selected for geochemical analyses including samples from 887A and 887C (10cm3 ea.) which were provided by the ODP to fill in missing sections at 20 cm resolution. All samples were weighed (wet and dry), freeze-dried, and ground using a tungsten-carbide disc mill before further preparation. Following the geochemical survey, the core was again sampled at 10 cm spacing (staggered from the geochemical intervals by 2 cm) where foraminiferal presence was indicated by %CaC03. These samples, totaling 281, were freeze-dried, weighed, and rehydrated in distilled water, after which they were wet-sieved at 150 um and oven-dried at 50°C. Foraminifera were then picked from specific dry-sieved size intervals of each sample for measurement of 1 3 C and 1 s O isotopes (Table Al . l ) . Table A l . l . Criteria for selection of foraminiferal samples used in counts and in stable isotope analyses. Habitat # of Specimens picked/sample mean | ranqe Taxa Fraction (um) Planktic 28 17-33 Globiqerina bulloides 250-300 - Neoqloboquadrina pachyderma 1 80-21 2 Benthic 3 2-4 Gyroidinoides spp. > 425 8 2-10 Uviqerina (senticosta) 250-350 6 3-10 Cibicides (wuellerstorfi, kullenberqi) 250-350 158 A1.2 X-ray Fluorescence Spectrometry A13.1 Sample Preparation A1.2.1.1 Minor Elements Four grams of each finely-ground sample were pressed in a stainless steel die at 10 tons pressure into boric acid-backed discs. In order to prevent contamination, these discs were then stored face-down in a box lined with paper tissue until XRF analysis was conducted. A1.2.1.2 Maj or Elements In order to minimize surface heterogeneity (particle) effects which are critical for the longer wavelengths and hence for elements with lower atomic numbers, fused glass discs were used for major element analysis. 0.4 g of sample was combined with 3.6 g of Spectroflux 105 1 5; the melting temperature of the flux is lowered to 700°C by lithium tetraborate and lithium carbonate, whereas lanthanum oxide (heavy absorber) increases the sample mass absorption thereby lowering mass absorption coefficient differences between sample matrices. The sample-flux mixture was fused in a platinum crucible for 20 minutes at 1100°C and allowed to cool, then additional flux was added in order to offset weight loss during fusion from volatile components. The sample was then remelted using a Meker burner and cast into a disc at 400°C using an aluminum mold. 1 5 47.0% Li2B40y; 36.7% UCO3; 16.3% La2C>3; weights to 3 decimal places. 159 Al.2.2 Analysis Element concentrations were determined by X-ray Fluorescence Spectrometry (XRF) using a Philips PW 1400 spectrometer fitted with a Rh target X-ray tube. Although a full suite of elements was run, only those discussed in this study are referred to here; instrument settings for these are summarized in Table A 1.2. Separate sample preparation techniques were employed for major and minor elements and are described in the preceding sections. Element concentration results are listed in Table A2.1. Table A 1.2. XRF instrument parameters for elemental analysis. Detection by gas flow counter using 90% Ar and 10% CH4; kv = 60, ma = 40. Element Line Crystal Collimator Fe K a lithium fluoride (200) 480 um Ti K a lithium fluoride (200) 480 um Si K a thallium acid phthalate 480 um Al K a thallium acid phthalate 480 um Mn K a lithium fluoride (200) 480 um Ba Lis lithium fluoride (200) 160 um In order to correct for matrix effects, the Compton ratio method (Reynolds, 1967) was applied to elements with Ka wavelengths shorter than the Fe absorption edge (i.e. atomic numbers > 27). This is based on the premise that Compton-scattered radiation of the primary X-ray beam is inversely proportional to the mass absorption coefficient (\i) at a particular wavelength. Thus, the amplitude of the Compton component of incident Rh Ka radiation controls the mass absorption correction. However, the relationship between \i and the Compton peak does not apply to elements with Ka wavelengths longer than the Fe absorption edge (i.e. atomic numbers < 27), since a significant quantity of Ka radiation that fluoresces from these elements is absorbed by 160 matr ix -he ld Fe. This is also the case for Ba, w h i c h is measured using the LP l i ne . For such elements , w h i c h i nc lude those l i s t ed i n Table A 1 . 2 , ma t r ix correct ions were based o n the ra t io of peak ampl i tude to the in tens i ty of an adjacent b a c k g r o u n d wavelength . A d d i t i o n a l cor rec t ions were r e q u i r e d for spectral overlap (e.g., Ba and Cr o n T i peaks) w h i c h increases marked ly i n the l ong wavelength r eg ion . T h e sy s t em was c a l i b r a t e d a n d m o n i t o r e d fo r a c c u r a c y u s i n g i n t e rna t i ona l rock a n d sediment s tandards (Geostandards Newslet ter , J u l y 1989). Measured and recommended element concentrat ions for the standards used i n this study are l is ted i n Tables A1.3 and A 1 . 4 . After runn ing a fu l l suite of m i n o r elements, i t was apparent that salt mig ra t ion i n a n u m b e r of sample discs had affected the accuracy of Ba concentrations, since up to a m o n t h h a d elapsed between the date of sample disc manufacture and XRF analysis . This was n o t a b l y ev ident i n samples f r o m the uppe r core a n d i n low-dens i ty (diatomaceous) intervals . As a result , affected sample discs were remade a n d r u n w i t h i n 48 hours using new sample mater ia l and a specific Ba a l g o r i t h m 1 6 ; s tandard concentrations for this r u n are shown i n Table A 1.5. A n a l y t i c a l p rec i s ion was est imated for major elements by r u n n i n g the same disc seven times and by runn ing six subsamples f rom the sample in terva l 887B 2H7-25 (Table A1 .6 ) ; for m i n o r elements, two subsamples f rom the same sample i n t e r v a l were r u n six t imes each (Table A 1 . 7 ) . These values are i n relat ive agreement w i th those obta ined i n other studies w h i c h used the same i n s t r u m e n t s e t t i n g s a n d s t a n d a r d s ( e . g . , D r y s d a l e , 1 9 9 0 ) . 1 6 Instrument parameters as listed in Table A1.2; samples rerun for Ba are indicated in bold type in Table A2.1. 161 Table A1.3. Measured and recommended major element concentrations for standards used in XRF analysis. Standard FE203 MNO TI02 CAO K20 SI02 AL203 MGO P205 NA20 % % % % % % % % % % G2 measured 2.70 0.03 0.50 1.93 4.36 70.27 15.90 0.57 0.11 4.17 measured 2.73 0.04 0.49 1.96 4.35 70.17 15.81 0.54 0.11 3.94 measured 2.73 0.04 0.50 1.95 4.39 71.20 16.06 0.59 0.11 3.87 measured 2.69 0.03 0.49 1.92 4.40 69.99 15.56 0.65 0.10 4.09 measured 2.72 0.04 0.50 1.96 4.26 70.87 15.82 0.61 0.12 4.69 measured 2.72 0.04 0.48 1.95 4.27 70.00 15.55 0.69 0.11 4.28 measured 2.73 0.04 0.49 1.96 4.17 67.71 15.15 0.60 0.13 3.65 measured 2.73 0.04 0.48 1.99 4.20 67.10 15.06 0.68 0.13 3.93 Mean 2.72 0.04 0.49 1.95 4.30 69.66 15.61 0.62 0.12 4.08 Std Dev 0.02 0.00 0.01 0.02 0.09 1.47 0.36 0.05 0.01 0.31 %Std Dev 0.57 12.34 1.70 1.09 2.04 2.10 2.28 8.58 9.30 7.71 recommended 2.66 0.04 0.51 1.94 4.40 69.48 15.33 0.80 0.13 4.13 AGV1 measured 6.95 0.10 1.09 4.95 2.98 61.41 18.38 1.59 0.47 4.24 measured 7.02 0.10 1.11 5.00 3.01 62.21 18.58 1.64 0.47 4.21 measured 6.88 0.10 1.07 4.86 2.99 61.08 18.08 1.51 0.46 4.31 measured 6.96 0.10 1.09 4.95 2.98 61.97 18.36 1.46 0.47 4.22 measured 6.97 0.10 1.11 4.91 2.85 61.77 18.07 1.66 0.49 4.64 measured 6.92 0.11 1.07 4.90 2.83 61.75 18.08 1.63 0.49 4.26 measured 6.96 0.10 1.06 4.92 2.82 59.56 17.60 1.41 0.47 4.13 Mean 6.95 0.10 1.09 4.93 2.92 61.39 18.16 1.56 0.47 4.29 Std Dev 0.04 0.00 0.02 0.04 0.08 0.89 0.32 0.10 0.01 0.16 %Std Dev 0.62 3.73 1.83 0.91 2.90 1.45 1.74 6.27 2.39 3.85 recommended 6.70 0.10 1.09 4.92 2.93 60.09 17.23 1.56 0.49 4.41 JG3 measured 3.76 0.07 0.48 3.73 2.59 69.54 16.45 1.62 0.17 3.98 measured 3.80 0.07 0.50 3.76 2.64 70.88 16.80 1.67 0.15 4.21 measured 3.77 0.07 0.49 3.71 2.60 70.84 16.62 1.66 0.16 4.26 measured 3.73 0.07 0.48 3.67 2.61 69.99 16.25 1.70 0.14 3.86 measured 3.77 0.08 0.49 3.74 2.48 70.20 16.44 1.89 0.16 4.45 measured 3.75 0.08 0.47 3.68 2.47 70.41 16.45 1.87 0.16 4.37 measured 3.79 0.08 0.47 3.73 2.46 68.01 16.10 1.79 0.16 3.95 measured 3.79 0.08 0.48 3.75 2.46 68.30 16.19 1.84 0.17 3.80 Mean 3.77 0.08 0.48 3.72 2.54 69.77 16.41 1.76 0.16 4.11 Std Dev 0.02 0.01 0.01 0.03 0.08 1.09 0.23 0.11 0.01 0.24 %Std Dev 0.62 7.13 2.15 0.87 3.06 1.56 1.40 5.99 6.24 5.94 recommended 3.71 0.08 0.49 3.73 2.58 68.28 15.72 1.84 0.14 4.18 JB2 measured 14.54 0.22 1.21 10.06 0.41 55.74 15.85 4.73 0.07 2.19 measured 14.55 0.22 1.21 10.11 0.42 55.73 15.79 4.61 0.08 1.81 measured 14.43 0.21 1.20 9.93 0.42 55.80 15.60 4.67 0.09 1.73 measured 14.27 0.21 1.17 9.81 0.42 54.98 15.32 4.58 0.08 2.20 measured 14.49 0.23 1.20 10.03 0.39 55.46 15.40 4.90 0.10 2.48 measured 14.33 0.23 1.17 9.92 0.38 55.28 15.43 4.71 0.08 2.27 measured 14.45 0.23 1.17 10.00 0.41 53.90 15.24 4.77 0.09 2.19 measured 14.45 0.22 1.17 10.03 0.42 54.00 15.18 4.84 0.10 1.84 Mean 14.44 0.22 1.19 9.99 0.41 55.11 15.48 4.73 0.09 2.09 Std Dev 0.10 0.01 0.02 0.10 0.02 0.77 0.25 0.11 0.01 0.26 %Std Dev 0.67 3.77 1.61 0.95 3.80 1.39 1.60 2.31 12.30 12.63 recommended 14.40 0.22 1.21 9.89 0.42 53.98 14.82 4.71 0.09 2.04 162 Standard FE203 MNO TI02 CAO K20 SI02 AL203 MGO P205 NA20 % % % % % % % % % % JA2 measured 6.31 0.10 0.70 6.43 1.88 57.93 16.29 7.94 0.15 3.92 measured 6.23 0.10 0.68 6.39 1.87 57.07 16.05 7.81 0.13 3.84 measured 6.22 0.10 0.69 6.30 1.87 57.97 16.09 7.69 0.14 3.67 measured 6.19 0.10 0.68 6.22 1.88 57.21 15.85 7.69 0.14 3.63 measured 6.38 0.11 0.68 6.40 1.69 56.25 15.41 7.62 0.19 3.18 measured 6.21 0.11 0.67 6.27 1.74 57.16 15.81 7.79 0.14 3.86 measured 6.36 0.11 0.65 6.38 1.68 52.58 14.32 6.97 0.16 2.35 Mean 6.27 0.10 0.68 6.34 1.80 56.60 15.69 7.64 0.15 3.49 Std Dev 0.08 0.01 0.02 0.08 0.09 1.86 0.66 0.32 0.02 0.56 %Std Dev 1.23 5.13 2.32 1.23 5.20 3.29 4.23 4.12 13.33 16.08 recommended 6.14 0.11 0.67 6.48 1.80 56.18 15.32 7.68 0.15 3.08 NIM-S measured 1.42 0.01 0.03 0.66 15.45 66.97 18.57 0.26 0.10 0.75 measured 1.40 0.01 0.02 0.67 15.29 65.32 18.22 0.28 0.10 1.02 measured 1.39 0.01 0.03 0.65 15.17 65.21 17.96 0.34 0.09 0.95 measured 1.38 0.01 0.03 0.65 15.18 65.59 17.98 0.38 0.08 0.89 measured 1.42 0.01 0.04 0.66 14.48 67.12 18.43 0.44 0.09 0.43 measured 1.41 0.01 0.01 0.66 14.39 67.17 18.56 0.28 0.09 0.54 measured 1.42 0.01 0.02 0.67 14.29 65.03 18.10 0.33 0.10 0.65 measured 1.42 0.01 0.01 0.68 14.28 64.56 18.03 0.30 0.10 0.48 Mean 1.41 0.01 0.02 0.66 14.82 65.87 18.23 0.33 0.09 0.71 Std Dev 0.02 0.00 0.01 0.01 0.50 1.05 0.26 0.06 0.01 0.22 %Std Dev 1.12 0.00 44.66 1.56 3.37 1.59 1.40 18.46 7.94 31.39 recommended 1.40 0.02 0.04 0.69 15.07 64.81 17.67 0.35 0.11 0.29 163 X XS CD V) 3 -o LH •O a .2 k* -M C 0) cj C o (J a CD a S-H o CD XS C CD e o L J CD •o -a CD Tt i — i < Si .3 as as ro ~ i 1 :0 S: C_> > - I 3g . - E <3 i-i = P>| •o l l k_ E, 5 - Q.I Q. k_ E, Isl Q. 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Sample FE203 MNO TI02 CAO K20 SI02 AL203 MGO P205 NA20 % % % % % % % % % % 2H7-25A 5.86 0.11 0.71 2.92 1.95 64 .59 14.70 2.64 0.14 3.33 2H7-25A 5.84 0.12 0.69 2.89 1.92 65.38 14.52 2.85 0.14 4.02 2H7-25A 5.86 0.12 0.70 2.91 1.92 64.57 14.50 2.60 0.16 3.51 2H7-25A 5.88 0.11 0.69 2.91 1.88 63.52 14.39 2.61 0.14 3.55 2H7-25A 5.90 0.11 0.70 2.96 1.91 64.07 14.41 2.72 0.15 3.48 2H7-25A 5.87 0.12 0.69 2.94 1.91 63.59 14.48 2.66 0.14 3.53 2H7-25A 5.88 0.12 0.70 2.93 1.92 63.49 14.37 2.54 0.1 5 3.43 Mean 5.87 0.12 0.70 2.92 1.92 64.17 14.48 2.66 0.1 5 3.55 Std Dev 0.02 0.01 0.01 0.02 0.02 0.71 0.1 1 0.10 0.01 0.22 %Std Dev 0.33 4.62 1.08 0.78 1.08 1.11 0.77 3.78 5.40 6.20 2H7-25A 5.86 0.12 0.70 2.91 1.92 64.57 14.50 2.60 0.16 3.51 2H7-25B 5.53 0.11 0.69 2.83 1.82 62 .39 14.13 2.53 0.13 3.23 2H7-25C 5.59 0.11 0.69 2.87 1.86 63.02 14.26 2.74 0.14 3.35 2H7-25D 5.58 0.11 0.69 2.91 1.88 63.59 14.24 2.65 0.15 3.43 2H7-25E 5.55 0.11 0.69 2.93 2.00 63.77 14.43 2.75 0.16 4.68 2H7-25F 5.53 0.11 0.69 2.93 1.87 63.47 14.42 2.76 0.14 3.31 Mean 5.61 0.11 0.69 2.90 1.89 63.47 14.33 2.67 0.15 3.59 Std Dev 0.13 0.00 0.00 0.04 0.06 0.73 0.14 0.09 0.01 0.55 %Std Dev 2.26 3.66 0.59 1.36 3.28 1.15 0.99 3.52 8.26 15.20 Table A1.7. XRF analytical precision for minor elements determined by 6 runs of each of two separate discs of a single homogenized sample. Sample Nb Zr Y Sr Rb Pb Zn Cu Ni Co Mn V Cr Ba Na ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm % 2H7-25A 6.6 131 22.5 255 66.1 13 91.9 50.8 38.2 21.6 1003 164 95.9 975 3.58 2H7-25A 6.5 131 22.8 2 6 0 65.4 8.4 92.1 49.9 41.7 19.8 972 165 95.4 955 3.56 2H7-25A 5.8 134 18.8 263 67.3 17.1 96.5 51.9 36.8 22.4 959 1 57 92.2 944 3.47 2H7-25A 5.9 134 20.1 2 5 9 68.3 17.7 92.6 50.3 44.3 20 962 158 91.8 945 3.52 2H7-25A 7.7 133 21.3 261 67.4 24.4 96.6 50.9 40.2 21 950 158 91.1 948 3.5 2H7-25A 6.9 132 21.4 265 66 16.8 93.3 53.9 40.7 21.3 938 152 88.5 923 3.48 Mean 6.6 132.6 21.2 260.4 66.8 16.2 93.8 51.3 40.3 21.0 964.0 158.9 92.5 948.4 3.5 Std Dev 0.7 1.2 1.5 3.5 1.1 5.3 2.2 1.4 2.6 1.0 22.4 4.9 2.8 17.0 0.0 %Std Dev 10.6 0.9 7.1 1.3 1.6 32.8 2.3 2.8 6.5 4.7 2.3 3.1 3.0 1.8 1.2 2H7-25B 5.9 133 17.8 259 66.5 i i i i i i i& i i 94 50.7 37.2 23.2 999 168 97.7 973 3.52 2H7-25B 9.2 133 17.4 259 66.2 10.9 95.3 52.5 36.8 22.7 996 169 95.1 980 3.54 2H7-25B 6.6 135 19.1 257 66.4 16.3 95.7 48.6 39 22.4 981 162 94.1 968 3.57 2H7-25B 5.7 138 22.9 261 64.1 12.7 95.3 51.8 36.2 20.7 9 6 4 155 95.2 967 3.51 2H7-25B 8.2 134 20 2 6 4 64.7 11.4 98.1 50.5 42.3 21.7 935 151 89.8 916 3.53 2H7-25B 6.3 134 18.8 261 63.4 17.1 96.7 51.2 39.8 22 975 156 95.6 966 3.45 Mean 7.0 134.5 19.3 260.1 65.2 13.7 95.9 50.9 38.6 22.1 975.0 160.0 94.6 961.6 3.5 Std Dev 1.4 1.7 2.0 2.4 1.3 5.1 1.4 1.3 2.3 0.9 23.5 7.2 2.6 22.8 0.0 %Std Dev 20.1 1.3 10.2 0.9 2.0 37.2 1.5 2.6 5.9 3.9 2.4 4.5 2.8 2.4 1.1 The shaded value is thought t o be anomalous and has been omit ted from the mean and standard deviation calculations. 169 A1.3 Total Carbon and Nitrogen Total carbon and nitrogen were determined using a Carlo-Erba NA-1500 analyzer according to the following procedure. Approximately 20 mg of bulk sediment (weighed to 3 decimal places) was encapsulated in a tin container and introduced via autosampler into a quartz flash-combustion tube which was oxygen-enriched and maintained at 1020°C. The resultant combustion products were carried by a helium stream through an oxidation catalyst 1 7. Silver-coated cobalt oxide was placed in the bottom of the combustion tube to further ensure complete oxidation and to retain interfering compounds. The combustion product mixture of CO2, N2,NO x, and water was then passed through a second (reduction) tube containing copper heated to 650°C. This facilitated the removal of excess oxygen and the reduction of nitrous oxides to N2 which were passed with CO2 and water through a water-absorbing filter (molecular sieve) containing magnesium perchlorate. N2 and CO2 gases were separated in a chromatographic column maintained at 54°C and subsequently measured by a thermal conductivity detector. This signal was then sent to a microprocesser for peak integration. To provide a calibration curve, the standards listed in Table A1.8 were interspersed along with 2 blank cups in each run of 50 samples. Linear regressions (least-squares method) of C and N (u,g18) vs. area (counts) were calculated for consecutive groups of 3 sample runs using values for standards and for blanks. Concentrations were then calculated as: X T Area (counts) x Slope of calibration curve UAf M — .... ' Mass of sample 1 7 0703; chromium trioxide. This catalyst was selected since it inhibits the formation of nitrogen oxides, does not react with the quartz tube, and does not adsorb nitrogen. 1 8 ug calculated from element % and sample weight. 170 Table A1.8. Elemental analysis standards for C and N determination. Standard %C %N # in each run wt. per sample (mg) PACS 3.570 0.284 2 10-30 MESS 3.089 0.208 2 10-30 BCSS 2.220 0.202 2 10-30 Acetanilide 71.090 10.360 5 0.1-2.0 (CH3COiNHC6H5) Accuracy was estimated by running each of three sediment standards 8 times as unknowns; results are given in Table A1.9. No systematic offset is apparent between recommended and measured %C values, though measured %N is consistently high by 0.008 to 0.18%. No correction was made for this offset, since it is at or near the precision (la) of the standard replicates. However, it is possible that a analytic artifact of this type is responsible for the ordinate intercept of 0.02 (%N) in Figure A l . l which may also reflect bound inorganic N in Site 887 sediment samples. Analytical precision (relative standard deviation; la), determined by six runs of sample 887B 4H3-95, is 1.1% and 3.3% for C and N, respectively (Table ALIO). 171 Table A1.9. Measured and recommended values for standards used in C and N analysis. Standard %N %C PACS 0.295 3.617 0.294 3.661 0.313 3.720 0.322 3.708 0.286 3.549 0.301 3.671 0.309 3.729 0.293 3.659 mean 0.302 3.664 std dev 0.012 0.060 %st dev 4.002 1.625 recommended 0.284 3.570 mean - recom. 0.018 0.094 MESS 0.207 2.951 0.219 3.111 0.266 3.259 0.204 2.949 0.213 3.054 0.211 2.923 0.225 3.154 0.199 2.932 mean 0.218 3.042 std dev 0.021 0.124 %st dev 9.663 4.079 recommended 0.208 3.089 mean - recom. 0.010 -0.047 BCSS 0.196 2.153 0.206 2.190 0.207 2.132 0.223 2.253 0.203 2.189 0.209 2.229 0.214 2.207 0.222 2.285 mean 0.210 2.205 std dev 0.009 0.050 %st dev 4.409 2.287 recommended 0.202 2.220 mean - recom. 0.008 -0.015 172 Table A 1.10. Analytical precision for C and N analysis estimated by running six separate splits of a homogenized sample. Sample %C %N B4H3-95 0.56 0.06 B4H3-95 0.55 0.07 B4H3-95 0.55 0.06 B4H3-95 0.56 0.07 B4H3-95 0.57 0.07 B4H3-95 0.56 0.07 Mean 0.56 0.07 Std Dev 0.01 0.00 % Std Dev 1.11 3.30 173 Organic carbon concentration (%C o r g) was determined by subtraction of inorganic carbon (by coulometry) from total carbon values, thus relative standard deviation for % C o r g is about 3.0%. Downcore % C o r g , % total N, and Corg/N values are shown in Table A2.1. A1.4 Coulometry Inorganic (carbonate) carbon concentration was determined by coulometry using a Coulometrics Inc. model 5010 coulometer coupled with a model 5030 carbonate carbon (CO2) apparatus; the analytical procedure follows. For each prepared sample, approximately 25 mg (weighed to 3 decimal places) of sediment is discharged into a sample tube and attached to the instrument. After the system is purged with C02-free air (scrubbed with KOH solution) for 2 minutes, the sample is heated and 2 ml of 10% HCI is dispensed into the tube to evolve CO2 gas from the inorganic carbon fraction of the sample. Evolved CO2 is then swept through a scrubber to remove extraneous product gases and into a titration cell which is filled with a solution of ethanolamine and a colourimetric indicator. As the gas stream passes through the solution, CO2 reacts with the ethanolamine via C 0 2 + HOCH2CH2NH2 - HOCH2CH2NHCOOH to form a strong, titratable acid and is quantitatively absorbed, causing the blue indicator color to fade. The cell is situated between a light source and a photodetector. The increase in light transmittance as the colour changes causes a current to switch on, thereby generating OH" ions via the reduction of H2O at a silver electrode: Ag°-*Ag + + e _ H 2 0+ e" ~ ^ H 2 +OH-174 As the acid is neutralized via HOCH 2CH 2NHCOOH + OH" HOCH2CH2NOO " + H 2 0 the solution returns to its original color and the current stops. The current acts as a titrant, and the total amount required for the titration is integrated and displayed as u.g C. %Ci n o r ganic is derived as follows: %C inorganic 1 Cco 2(^g) -Cblank(u-g)^ sample weight (ug) J xlOO Blank values varied from 1.93 to 5.71 ug C with a mean of 2.82 ug C, and were run and used for correction with each group of analyses (~ 30 samples). Accuracy was estimated by mnning a standard of pure CaC03 (12.00 %C) 42 times throughout sample analysis, yielding a mean value of 11.97% with relative standard deviation (RSD; la) of 0.35%. Replicates were run for 64 samples of varying composition and the mean for each pair taken as the value for that interval; precision (mean RSD) for all replicate pairs is 2.04%. %CaC03 results for all samples are presented in Table A 1.2. 175 A 1.5 Mass Spectrometry Al.5.1 Calcite Analysis Isotope measurements were performed on a VG Isotech Prism mass spectrometer using a VG autocarb common-bath carbonate preparation system according to the following method. Following sample preparation, planktonic foraminifera G. bulloides and benthic foraminifera Gyroidinoides, Uvigerina, and Cibicides spp. were picked from specific size fractions (Section Al . l ) and stored in Nalgene thimbles. Immediately prior to analysis, specimens were loaded in groups of 20 into an aluminum carrying boat and pulverized in ethanol with an aluminum rod. Samples were then roasted in vacuo at 430°C (~ 10"4 torr) for a minimum of 30 minutes in order to remove organics. This temperature was selected as optimum since C o r g decomposes thermally at ~ 400°CandCaCO3 decomposition (i.e. CO2 evolved) begins at ~460°C. Roasted samples were transferred to the preparation/inlet system carousel where they were individually released into a 100% phosphoric acid bath at 90.0°C. Evolved CO2 was swept through a stainless steel water trap (sustained at -90°C by a cooling coil bathed in methanol) to a CO2 trap maintained at -196°C by liquid N. The CO2 trap (cold finger) was then warmed and the vapour passed through the mass spectrometer inlet where it is ionized in high vacuum by an electron beam for analysis. The system uses 3 plate collectors for ion detection. The instrument was calibrated to the PDB standard using NBS-19 and NBS-20 (National Bureau of Standards) as working standards. Measured isotope (6) values were taken as the mean of 5 X 20 aliquot measurements of a given sample; (internal) precision for these values was better than 0.008%o (la). Precision of the system, determined from a laboratory standard run with the 176 samples over a 2-month period, was 0.07 and 0.06%o for 6 l s O and 6 1 3 C, respectively (la; n=28). Moreover, analysis of 27 replicate pairs of 887B G. bulloides samples taken throughout the core (Figure 5.9) yielded a mean standard deviation of 0.10 and 0.07%o for S 1 8 0 and S 1 3 C, respectively. Isotopic results from all foraminifera samples are shown in Table A2.3. A2.5J? Organic Carbon Prior to analysis, accurately-weighed 500 mg samples19 were placed in glass vials and treated with 10% HCI solution in order to remove carbonate carbon. After oven-drying for 2-3 days at 50°C, samples were introduced into a Carlo-Erba NA1500 NC elemental analyzer coupled to the mass spectrometer. Evolved CO 2 was isolated by chromatographic separation and passed in a helium stream to the mass spectrometer preparation/inlet system where analysis proceeded as described in the foregoing section. samples used were splits of those prepared for geochemical analyses (Section Al.l). 177 A 1.6 Foraminifera Counts Foraminifera counts were done with a binocular microscope using samples selected for 6 l s O analysis in the sieve fraction > 150 urn. After samples were wet sieved and oven-dried at 50°C, they were distributed on a matte-black gridded tray and benthic species Gyroidinoides, Uvigerina spp,, and Cibicides spp. were counted using the entire sample. However, in the case of planktonic species G. bulloides and JV. pachyderma, it was frequently necessary to split the sample into fractions in order to count accurately due to the large numbers of foraminifera present. Samples were split up to six times (1/64); total individuals counted ranged from ~ 300 to 1000 in sample splits and from 0 to 600 in entire sample fractions. Relative standard deviation (la) of eight replicate counts of separate samples (200 to 400 specimens each) dispersed through the core is 5%. In addition, a qualitative survey was conducted to determine the presence of additional foraminiferal species and sediment parameters such as amount and type of ice-rafted debris and the presence of unique mineral phases like pyrite. 178 Appendix 2: Data 179 o TJ 0) OO _0j >x DO _o "o X (!) c OO CO 00 00 T3 a u < X < CD iJ o „ s BEE i ro " 18 E CD cu < U C T3 o d CO d Fe/AI (ratio) 0.6451 0.566| 0.4381 0.61 91 0.7251 0.655| 0.7471 0.6871 0.6261 0.689] 0.9931 0.661 1 0.736| 0.780| 0.546| 0.522| 0.5371 0.703| 0.634J 0.6241 0.528| 0.590| 0.551 | 0.5461 0.656| 0.6631 999'0 ] 0.736| 0.5621 0.6551 0.6051 0.593| 0.621 | 0.61 21 0.5721 0.690| 0.661 | 0.596| 0.5621 0.457| 0.5691 0.523| 0.3481 0.3451 0.356| 0.655| 0.7041 0.7881 0.635| 0.720| 0.6221 0.6751 Mn/AI (ratio) 0.0111 0.010 0.0091 0.011 0.009 0.013 0.015 0.018 0.022 0.034 0.028 0.033 0.022 0.021 0.010 0.009 0.010 0.01 o| 0.011 0.012 0.010 0.010 0.010 0.010 0.010 0.012 0.014] 0.014 0.010 0.010 0.009 0.010: 0.014 0.010] 0.009 0.010' 0.012 0.012 0.011 0.011 0.011 0.010 0.008 0.008 0.008 0.012 0.014 0.017 0.011 0.016 0.016 0.025 Ti/AI (ratio) 0.061 | 0.0591 0.0421 0.061 | 0.0641 0.061 | 0.0661 0.061 | 0.0571 0.062] 0.0651 0.0591 0.066| 0.0581 0.056| 0.054] 0.054| 0.0631 0.0631 0.058] 0.055] 0.0591 0.0541 0.056| 0.061 ! 0.059] 0.061 ] 0.058J 0.059] 0.062 0.061 0.064 0.065] 0.063] 0.056] 0.061 0.061 0.060] 0.056 0.046 0.055 0.050 0.030 0.031 0.031 0.075] 0.080 0.095] 0.060] 0.061 0.058 0.056] Ba/Al (ratio) 0.0131 0.017| 0.018| 0.0271 0.030| 0.030 0.031 0.030 0.032 0.038 0.046] 0.039] 0.035 0.026 0.010 0.011 0.015] 0.020 0.023] 0.025] 0.018 0.013 0.022 0.031 0.038 0.038 0.043 0.038 0.026 0.028 0.027 0.026 0.016 0.019 0.016 0.028 0.024 0.016 0.023 0.026 0.026 0.024 0.012 0.012 0.011 0.026 0.020, 0.006 0.034 0.033 0.021 0.026 Si/Al (ratio) 4.091 5.721 10.1 51 17.541 21.10| 20.08J 23.33] 20.33] 19.81 | 25.14 30.161 21.17 19.20] 13.36] 3.49 3.59 4.38 4.351 3.79] 3.76] 3.69 3.68 6.30 11.061 11.38 13.37 15.26 cn CD 5.09 5.24 CM LO 4.37 3.85 3.72 3.77 5.06] 4.57 3.70 4.47 4.48 3.80 4.06 4.20] 4.16: 4.17 3.48 3.37 2.95 3.15 3.23 3.26 3.82 AL203 (wt%) 1 3.481 10.36| 6.73| 3.93| 3.2ll 3.37| 2.90] 3.33] 3.40] 2.55 2.09 3.08 3.27 4.88 16.05] 15.96 13.74 13.01 14.38 13.83 15.09] 14.54 10.27] 6.07 5.78 4.98 4.09 5.26] 11.72] 11.16 11.47 12.89 15.14 14.86] 15.45 11.50 12.16] 14.39 13.04 13.33 14.62 14.21 14.88 14.96 14.95 14.90 14.78 17.02 15.77 15.06 14.14 9.96 SI02 (wt%) 62.43| 67.05| 77.331 78.03J 76.681 76.63| 76.61 | 76.65] 76.28] 72.59 71.37 73.81 | 71.09 73.80] 63.34] 64.89 68.11 | 64.10] 61.67 58.85 63.11 60.53] 73.28 76.04 74.50] 75.40 70.66 69.24 67.481 66.17 67.70 63.74 65.97] 62.55 65.94 65.94 62.92 60.36 66.00 67.59 62.90 65.26 70.70 70.41 70.64 58.69 56.34 56.85 56.20 55.13 52.24 43.11 CO E CO CL C L Ln Ol LO in C7> cn cn CD cn CD m o LTi o LO o CO •3-L O cn LO cn LO 60S m o LO cn cn CD o o CD o CO CD LO CO cn CM cn CD CO O o CD CO LO CO cn o o o CO cn LO cn cn CD LO o o m cn cn CO CD O CD CO LO LO CD LO LO CD LO r-LO O cn o o LO cn 00 CM LO CD CD LO LO cn CM o cn LO CD CO 2000 1794 LO cn LO oi cn ID CO 2088 1579 r*. LO 1/1 2821 2621 1556 1345 cz a CM CD CO CD CD O cn -3-CM CM cn o CM CD CM CM CM 01 CD CD CM CM •cr-•3-CO LO cn cn CD CO CD CO LO CO CD CM CD cn cn CD CO cn CD LO O CO CM CO 00 CO o cn <*-o CO LO CM CO CM LO cn CD CO CD CD cn LO cn CO cn CD CM cn cn CM cn h-CD CM cn 1123 00 LO CO CO CD CM 00 1037 o CD cn LO CO CO LO cn t^-CM 00 o cn CM CO cn cn CO cn 1074 1542 CM cn CO 1367 1284 1420 13Corg (%o) -23.12| -21.67| -19.341 -19.431 -19.27| -19.441 -19.21 | -1 9.721 -20.40| -20.571 -20.10| -22.01 | -23.74| -21.94| -24.1 21 -23.631 -23.10| -21.921 -21.74| -23.29| -23.781 -23.731 -22.18| -20.951 -21.43] -21.04] -21.61 | -21.961 -22.62] -21.44] -20.43] -20.17] -20.25] -20.59] -22.70] -21.26] -22.93] -23.59] -23.04] -23.36] -23.49] -23.25] -24.16 -23.49 -23.48 -23.74 -23.10 -23.21 -23.71 -22.99 Corg/N (ratio) 7.54| 7.71 8.71 | 8.56| 8.021 8.20| 7.90| 7.631 7.241 7.63J 7.00| 7.34] 7.08] 7.15 8.75] 7.90| 7.911 7.50] 7.62| 7.06] 7.58| 8.71 ] 8.1 51 8.45] 8.63 7.84 8.06 7.27 7.53 8.10 8.46 8.06 7.78 7.45 6.81 CO CO 7.82 7.76 6.87] 6.25 5.94 4.91 | 3.16] 2.83] 1.93] 5.59 5.78 2.62 5.41 5.57 10.73 7.99, 0.0631 0.0621 0.0851 0.081 | 0.0851 0.0841 0.0821 0.071 | 0.0651 0.058| 0.0641 0.0581 0.066| 0.0691 0.056| 0.0481 0.051 | 0.0791 0.0781 0.0661 0.0601 0.0551 0.0681 0.0861 0.0891 0.082] 0.071 | 0.071 | 0.0641 0.0871 0.122] 0.114] 0.0791 0.0831 0.058] 0.114 0.084] 0.057] 0.056] 0.046] 0.048] 0.037 L LO'O | 0.012] 0.014 0.042 0.032 0.011 | 0.062] 0.061J 0.080 0.072 Corg accum rate (g/cm2 ka) 0.041 0.031 0.031 0.021 0.021 0.021 0.021 o d o d o d o d o d o d o d 0.04 0.03 0.03] 0.03] 0.03] 0.03] 0.03 0.03 0.04 0.04 0.04 0.03 0.03] 0.03 0.05 0.07 0.10 0.10 0.07 0.08 0.05 0.08 0.06J 0.06 0.03 0.01 0.01 0.01 0.00 0.01 0.01 0.00 0.01 0.02 0.04 0.03 Corg (wt%) 0.471 0.471 0.741 0.70| 0.681 0.69| 0.65] 0.54] 0.471 0.45] 0.44] 0.42] 0.471 0.49 j 0.49 j 0.38] 0.40| 0.60] 0.601 0.471 0.46] 0.481 0.55] 0.73 0.77 0.64] 0.58] 0.51] 0.49 0.71 1.03 0.92 0.62 0.62 0.39 0.93 0.65] 0.44J 0.39 0.29] 0.29 0.18 0.03 0.03 0.03 0.24 0.19 0.03 0.33 0.34 0.86 0.58 CaC03 accum rate (g/cm2 ka) 0.01 1 0.01 1 0.001 0.001 o.ooj 0.01] 0.011 0.031 0.12| 0.1 71 0.091 0.1 51 0.13] 0.091 0.021 0.001 0.001 o.oo| 0.021 0.281 o.iol 0.01 | 0.001 0.00] 0.06] 0.13] 0.521 0.52] 0.071 0.01 j ______ 0.01] 0.01 | 0.01 ] 0.01 | 0.01 | 0.01 | 0.02 j 0.02] 0.01 j 0.07] 0.001 0.01 0.01 0.08 0.14 0.02 0.12 0.12] 0.49] 1.19 s * U * J TO > o & 0.1 61 0.1 Ol 0.081 0.061 0.061 0.251 0.43] 1.651 5.30| 7.751 4.571 6.441 5.61 ] 3.97] 0.24| 0.041 0.05] 0.091 0.341 4.90| 1.431 0.1 ol 0.061 0.091 __ _ 1.09 2.551 9.451 8.781 0.73] 0.10] 0.09] ___ 0.12] 0.06] 0.10| 0.1 Ol 0.10] 0.24] 0.13] 0.14] 1.67 0.14] 0.18 0.02 0.05 0.20 CD cn 3.72 0.30 3.18 2.44 9.96 26.25 Age, ka Calendar (ka) 77.091 77.731 78.641 79.551 80.45] 81.36| 82.271 83.281 84.69] 86.1 0| 87.51 | 88.921 cn cn o cn 91.74] 93.1 5 i 94.521 95.82] 97.12] 98.421 99.71 | 101.01 102.31 ] 103.61 | 104.46] 105.13] 105.79 106.45 107.11 107.78 108.44] 109.10 109.77 110.43] 111.36] 111.75] 112.42 113.08 113.74 114.40 115.20 116.80 118.80 119.80 119.80 119.80 121.00 123.00 124.80 126.20 128.16 129.72 130.19 131.29 Age 14C (ka) Composite depth (m) 5.521 5.591 5.691 5.791 5.891 5.991 6.091 6.19] 6.291 6.391 6.491 6.591 6.691 6.791 6.89] 6.99] 7.091 7.1 91 7.29] 7.391 7.491 7.59] 7.69] 7.79] 7.89] 7.99 8.09 8.19 8.29 8.39] 8.49 8.59] 8.69] 8.83] 8.89] 8.99] 9.09 9.19 9.29 9.39 9.49 9.59 9.69 9.79 9.89 9.99 10.09 10.19 10.29 10.39 10.49 10.52 10.59 LE Ident 2H2-38 1 2H2-45 1 2H2-55 1 2H2-65 1 2H2-75 1 2H2-85 | 2H2-95 | 2H2-105 | LO CM X CM 2H2-125 | 2H2-135 | 2H2-145 | 2H3-5 1 2H3-15 | 2H3-25 | 2H3-35 | 2H3-45 | 2H3-55 | 2H3-65 | 2H3-75 | 2H3-85 | 2H3-95 | 2H3-105 | 2H3-115 | 2H3-125 | 2H3-135 | 2H3-145 | 2H4-5 | 2H4-15 ] 2H4-25 j 2H4-35 | 2H4-45 ! 2H4-55 ! 2H4-69 | 2H4-75 | 2H4-85 | 2H4-95 2H4-105 2H4-11 5 2H4-125 2H4-135 2H4-145 2H5-5 2H5-15 2H5-25 2H5-35 2H5-45 2H5-55 2H5-65 2H5-75 2H5-85 2H5-88 2H5-95 SAMPi Hole 1 887B 887B 1 887B 1 887B 1 887B 1 887B 1 887B 1 887B | 887B | 887B | 887B | 887B | 887B | 887B | 887B | 887B | 887B | 887B | 887B ] 887B | 887B ] 887B | 887B ] 887B j 887B ] 887B | 887B 1 887B ] 887B 887B 887B | 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B Strat B i l l m a m • 1 Fe/AI (ratio) 0.575| 0.534| 0.629| 0.5321 0.532| 0.5601 0.5541 0.4951 0.51 51 0.256| 0.51 51 0.510| 0.5631 0.559| 0.5851 0.6221 0.5761 0.5421 0.6551 0.731 | 0.591 | 0.655| 0.527| 0.539| 0.6321 0.5801 0.6181 0.539| 0.522| 0.5331 0.632| 0.6561 0.619| 0.61 ll 0.6571 0.6651 0.582] 0.5941 0.561 ] 0.5681 0.5991 0.7381 0.6061 0.574] 0.6091 0.7001 0.5641 0.5381 0.5611 0.611 | 0.813] 0.700| Mn/Al (ratio) 0.0191 0.01 51 0.021 | 0.01 5| 0.01 31 0.011 | 0.011 | 0.01 21 0.011 | 0.0081 0.010| 0.0091 0.010| 0.011 I 0.017] 0.027] 0.013] 0.012| 0.0231 0.029] 0.012| 0.01 4| 0.011 | 0.01 4] 0.0201 0.0251 0.01 5.| 0.01 51 0.012| 0.01 21 0.01 71 0.01 51 0.016| 0.0151 0.01 31 0.020| 0.012| 0.011 1 0.010I 0.011 | 0.011 | 0.013] 0.011]   0.011 | 0.01 31 0.011 | 0.010| 0.011 | 0.01 3| 0.0221 0.026| Ti/AI (ratio) 0.056| 0.0551 0.0541 0.0541 | 0.0551 0.056] 0.056] I 0.048| 0.053| 0.020] 0.055] r0:053] [ 0.056| 0,055] 0.056| 0.055] |0.058| 0.056| f 0.059| 0.062] f 0.069] f 0.0571 | 0.055] | 0.054] 0.053] 0.052| 0.0521 | 0.055] 0.0551 0.056| | 0.056] 0.057] 0.0591 ro.o59i 0.0591 10.059] ro.057] 0.0601 0.0571 | 0.0581 0.059] ] 0.062] 0.0601 0.0571 0.060| 0.065] | 0.055] [ 0.0551 f 0.057| 0.061 ] 0.079] 0.061 | Ba/Al (ratio) 0.021 | 0.014 0.01 9 j 0.015] 0.014] 0.013] 0.014] 0.015] | 0.015 0.013 0.014] 0.012] 0.011] 0.012] 0.018 0.025 0.011 0.012] 0.029] 0.028 I o.ofTl | 0.0131 | 0.012] | 0.014l 0.022 0.027 | . 0.015] [ 0.013l 0.013] | 0.012] | 0.023 0.006 0.013] 0.024] | 0.035 0.029 [ 0.01 ll 0.011] | 0.010] | 0.014] 0.013 0.023 0.023] 0.016] 0.025] 0.027 r 0.019] | 0.013] 0.013] | 0.015] | 0.0251 0.045 Si/Al (ratio) 3.74| 3.41 j 3.79| 3.7T1 3.66] 3.57] 3.90] 3.89] 3.68 5.10] 3.68 3.48 3.45] 3.52 3.99 5.31 3.81 3.82 11.60 9.43 3.41 3.59] 3.88] 4.63] 8.40 11.96 5.98 f 3.78] 3.87 3.66 7.72 4.74 4.12 5.78 9.54 6.81 3.57 3.50 3.63] 3.78] 3.57 3.48 3.24 3.62 3.98 3.60 [ 3.671 | 3.85 3.52 4.02 8.33 12.50 AL203 (wt%) 11.531 15.17| 11.921 1 3.1 ll 14.60| 15.67] 14.27] 14.49] 15.31] 1 2.80] 15.39] 16.32] 15.75] 15.67] 13.16 8.78 14.37] 14.70] 5.00] 6.02 15.54 15.08] 14.70 12.18] 6.67 1- 4.76 . 9.54 14.97] 14.60 15.26 7.71 11.91 13.14 10.49] 6.70 8.98 15.24 15.39] 15.40 14.72 15.29 14.63 15.93 15.18] 13.89 14.19 14.94 14.69 15.22 13.48 7.17 5.00 SI02 (wt%) 48.76| 58.621 51.15 55.07| 60.471 63.331 62.94| 63.821 63.751 73.85] 64.21 | 64.31 | 61.451 62.441 59.40] 52.80] 61.921 63.591 65.69| 64.30] 59.92] 61.371 64.591 63.91] 63.40] 64.47] 64.55] 64.101 64.011 63.271 67.40] 63.89] 61.361 68.63| 72.35] 69.25] 61.68| 61.001 63.23| 62.971 61.86] 57.61 ] 58.49] 62.30| 62.55] 57.78 62.16] 63.97] 60.72] 61.39] 67.60] 70.79] « I CQ C l Q . CO o cn o CO LO o LO o O CO 1-CNJ cn LO r ~ CO CO IN- CNI O CNJ CT) CO CO CNJ o CD CO CNJ CNJ ~^ LO CO LO CNJ on CNJ C~ on CO LO CO CO [ 1044| CNI cn LO o on CD r~-CO CD CD CD 1042] 19 L0 L r~ co cn CNJ cn CO CD CO m CNI CD | 1322] | 1229] ! 1392] CO CO CD CO CO CO 11116] ! 1066] | 1804] | 1961| [ 1254] I 1824] ] 2010] ! 1503) | 1017] I 1063] ]990L | LO cn | 1195] c E S 3 1144| 1175| 1 367] 1035] CO on cn cn cn CO h~ o CO cn o cn CO CO CO CO CO CO cn fN. •3" CD CO CNJ CO cn 1409 ] 1564] 1102 I 1074 o LO CO 1219 1083 1037 cn CD cn 1056] CD cn r~ r~ CO r^. co O cn 1256 1040] Pl002 CD CD CD CO LO 1158 1046 CO LO 1195 CNI CNJ 1051 n-CD CO f--CM CT) CD O cn N . CD cn CO 0 cn IT) 00 CO CO 0 cn CNJ cn on 00 00 cn 00 1020 1080 CNJ O on CJ) o cn SI. -23.081 -23.85J -22.81 | -23.03] -23.90] -24.11] -24.05] -23.86] -24.19] -24.51 -24.35] -24.22] -24.35] -21.89 -22.13 -23.83 -23.12 -21.98 -21.47 -23.06 -23.94 -23.57 | -21.87] ^22.75 -22.56 -22.05 -23.60 -24.24 -24.28 -24.16 -22.75 -22.55 Corg/N (ratio) 8.291 8.091 7.31 9.491 7.861 7.57 7.62| 6.76| 7.33] 4.05] 7.541 7.27] 6.731 7.141 7.02] f 7.431 7.34| 7.53] 7.95] 7.70 8.26] "8.461 9.41 J 7.77 7.64] 7.92] 7.85] 5.90] 7.931 6.50] 6.26] 7.96] 8.12] 8.13] 7.57] 6.85 8.09] 8.94] 8.421 I 7.56! 1 7.43] | 7.59] 7.07 1 6,94] 6.37] 1 6.06] | 5.83] | 6.88] 8.37] 8.41 1 7.75 8.08] s | 0.0721 0.0631 0.070| 0.0641 0.0521 0.0531 0.0491 0.0471 | 0.0491 I 0.022] | 0.056] 0.056| 0.056] ( 0.052] j 0.055] i 0.0721 ( 0.055] ] 0.055] ! 0.070] | 0.076 | 0.058] | 0.061 ] ] 0.055] f 0.055] ! 0.068] | 0.058] ] 0.059] | 0.049 ! 0.046] | 0.044 | 0.053] ] 0.084] | 0.087] ] 0.066| I 0.068] ] 0.0631 [ 0.060] ] 0.045] ] 0.044] | 0.053] | 0.048] | 0.073] | 0.059] 1 0.049J | 0.054] 1 0.049] 1 0.043] | 0.042] | 0.053] | 0.052 | 0.047] | 0.055] Corg accum rate (g/cm2 ka) 0.031 [ 0.03 0.03] 0.04 0.03 0.02 0.03 0.02 ] 0.02 0.03 | 0.03 1 0.02 ] 0.02 ] 0.02 I 0.02 ! 0.02] ] 0.03 | 0.01 | 0.01 | 0.03 | 0.03 | 0.03 I 0.02 | _ ._ 0.02 | 0.01 | 0.02 | 0.02 | 0.03 | 0.03 1 0.02 | 0.04 | 0.03 | 0.02 | 0.01 | 0.01 | 0.02 | 0.02 | 0.02 | 0.02 | 0.02 | 0.02 | 0.02 | 0.02 | 0.02 | 0.01 | 0.01 I 0.02 | 0.02 | 0.02 | 0.01 | 0.01 Corg (wt%) 0.601 0.51 0.511 0.61] 0.41 ] 0.40 0.38 0.32] 0.36 0.09 1 0.42! 0.40 0.38 0.37 0.39 | 0.54 | 0.40 | 0.42 | 0.55 | 0.59 | 0.48 0.52 0.52 0.43 | 0.52 1 0.46 | 0.46: | 0.29 0.36 | 0.29 ] 0.33 | 0.67 0.71 | 0.53 ] 0.51 | 0.43 ] 0.48 0.41 1 0.37 1 0.40! | 0.36 [ 0.56 | 0.42 | 0.34 ] 0.35 | 0.29 ] 0.25 ] 0.29 | 0.45 ! 0.44; | 0.36 | 0.45 CaC03 accum rate (g/cm2 ka) 1.07| 0.40] 0.84 0.77] 0.44] 0.03 0.28 0.13 0.01 0.02] 0.01 | 0.06 0.06 0.32 0.60] 0.13] 0.02 0.21 0.19 0.09 0.01 | 0.04 0.25 | 0.39 L 0.34 |. 0.29 0.01 0.01 0.05 0.30 0.03 0.03 0.00 | 0.04 0.04] 0.01 ] 0.01 0.01 0.00 0.00 0.00 0.00 0.01 0.00 | 0.01 | 0.01 0.01 0.01 0.01 0.04 0.01 CaC03 (wt %) 20.1 91 6.331 16.18 12.82J 6.381 0.56) 4.041 2.491 L_ 0.141 0.02] 0.231 0.14| 1.081 0.85] 6.92 CO CO 2.17] 0.32] 9.06] 1. 8.21] 1 1-42] | 0.26] 1 0.71] 5.12 | 12.48] | 13.53] 1 • 7.65] 0.11] L 0.14] 1 0.521 cn CD 0.49] 0.78] 0.12] 1 1-70] 1.34 | 0.20] 1 0.23] 0.10] 0.05] | 0.06] 0.07] 0.09] 1 0.12] I 0.10] 0.15 1 0.16] ] 0.20] 0.17 1 0.14 1 1-74 | 0.38] Age, ka Calendar (ka) I32.85| 134.261 1 35.351 1 36.921 1 38.481 I40.05| I41.30J I42.86| I44.43| 145.52] 145.521 I46.93| 148.34] 149.90] 151.49 153.13] I 54.77] 156.09] 157.73] 158.88 161.01 | 162.65] 164.29] 165.93 166.59 167.57] 169.22] 171.35] 172.33] 1 74.471 175.941 179.72 183.00 186.66] 190.13 191.96 193.79] 195.62] 197,451 I 99.28] ] 201.11 | 202.94 CD r~ •<f' O CNJ | 206.59 | 208.42 | 210.25 | 211.71 | 213.54 | 215.37 | 217.20 | 218.85 | 220.86 | 222.14 Age, ka Calendar (ka) ] 201.11 | 202.94 | 206.59 | 208.42 | 210.25 | 211.71 | 213.54 | 215.37 | 217.20 | 218.85 | 220.86 | 222.14 Age 14C (ka) Composite depth (ml cn CD CD CT) CO CT) cn cn o cn cn CNJ on CO cn •3- cn LO o IN. cn CO CO cn cn cn o cn cn CNJ CD CO cn CD LO cn CD cn r~ cn CO CD cn CO o cn o cn CNJ CO on CO CO IO CNJ CD LO CO CO 1.05 LO CNI . J -<t •3- LO ^-CD Tj-•a- 1-00 cn 0 LO -^10 •3-CNI i n co i n Nt i n TJ-m i n •3-CD i n •3-m 00 LO -^cn i n CO 0 CD CD -3-CNI CD Composite depth (ml LL Ident LO o LO LO CNJ LO CO LO 2H6-5 | 2H6-15 | 2H6-25 | 2H6-35 | 2H6-45 | 2H6-56 | 2H6-65 | 2H6-74 | 2H6-85 | 2H6-95 | LO o LO LO CNJ LO CO CNJ 2H7-5 ] 2H7-15 | 2H7-25 ] 2H7-35 | 2H7-39 2H7-45 (2H7-55 ] |2H7-68 2H6-38 2H6-121 [2H6-61 I2H6-84 | |2H6-104 2H6-124 LO LO LO CD LO LO 00 LO on LO O LO m CNI m CO m |3H2-5 |3H2-15 |3H2-25 |3H2-35 |3H2-45 |3H2-55 | |3H2-64 |3H2-75 |3H2-85 LL Ident 2H5-1 2H5-1 2H5-1 2H5-1 2H5-1 2H6-5 | 2H6-15 | 2H6-25 | 2H6-35 | 2H6-45 | 2H6-56 | 2H6-65 | 2H6-74 | 2H6-85 | 2H6-95 | |2H6-" 2H6-1 2H6-1 2H6-' 2H6-' 2H7-5 ] 2H7-15 | 2H7-25 ] 2H7-35 | 2H7-39 2H7-45 (2H7-55 ] |2H7-68 2H6-38 2H6-121 [2H6-61 I2H6-84 | |2H6-104 2H6-124 « • « |3H2-5 |3H2-15 |3H2-25 |3H2-35 |3H2-45 |3H2-55 | |3H2-64 |3H2-75 |3H2-85 LL Ident 2H5-1 2H5-1 2H5-1 2H5-1 2H5-1 2H6-5 | 2H6-15 | 2H6-25 | 2H6-35 | 2H6-45 | 2H6-56 | 2H6-65 | 2H6-74 | 2H6-85 | 2H6-95 | |2H6-" 2H6-1 2H6-1 2H6-' 2H6-' 2H7-5 ] 2H7-15 | 2H7-25 ] 2H7-35 | 2H7-39 2H7-45 (2H7-55 ] |2H7-68 2H6-38 2H6-121 [2H6-61 I2H6-84 | |2H6-104 2H6-124 3H' .HE 3H' X CO 3H' 3H' X CO 3HP 3H-X CO |3H2-5 |3H2-15 |3H2-25 |3H2-35 |3H2-45 |3H2-55 | |3H2-64 |3H2-75 |3H2-85 SAMP Hole 1 887B I 887B I 887B | 887B | 887B | 887B | 887B | 887B | 887B ] 887B | 887B ] 887B ] 887B ] 887B j 887B 887B 887B ] 887B j 887B 887B 887B 887B ] 887B | 887B 887B 887B 887B 887B 887A 887C | 887A 887A 887A 887A 887B |887B |887B ]887B |887B |887B [887B ]887B |887B ]887B ]887B |887B |887B |887B |887B |887B |887B |887B |887B Strat I I m IJIillli 1 < 5 \ "ro Is-CO LO o Is-LO LO Is-LO CM CM to o to LO LO LO o CM to LO LO Is-cn Tf Tt Is-to CM Is-o LO to cn LO to CO Is-Is-Is-to LO Ti-to to LO ro CM N-LO CM LO cn CO LO CO CM LO LO LO cn LO LO LO CO LO cn Tj-LO Ti-to LO LO Tj-LO LO LO CM Tt-LO 0 0 Tf LO ro to to I S -Tj" LO Ti-Ti-LO o to LO Tj-LO LO to CM LO CO CM to o to CO to o to CM to cn CO to o LO LO to LO Ti-CO LO LO Tf to LO I S -0 0 LO LO CO to CO r-to o CO to LO OO to cn CO to LL. ^ o o o o o o O o o o o o O o o o o o O O o o o O o o o o o O o O o o o o o O o o o O o o to o o o o O o o O to ro CM LO CM t - ro «- to to to CO cn •- CM CM to •- cn CO CM CO O cn o o cn o CM cn CM o o .- o o •- 1- CM CM LO o «- LO CM CO ro ro Is- r- to Mn// (ratic o O O o o o O o o o o o O o o O o o o o o o o o o o o o o O o o o o o o o o o o o O o o o o O o o O o o o Mn// (ratic d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d /Al itio) CM to o to o Is-LO O CM to o CO LO o cn LO o Is-LO o LO o o LO o to to o cn to o o to o cn LO o to LO o cn LO o CO LO o cn LO o LO LO o CO LO o to LO o cn LO o CO LO o cn LO o CO LO o Is-IO o LO o Is-LO o CO LO o to LO o LO LO o CO LO o CO LO o CO LO o to o o to o r-LO o Is-LO o cn LO o o to o o to o CM to o cn to o ro to O ro LO o Ti-o to LO o to LO o I S -LO o o to o ro to O CO to o CO to o to o d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d /Al tio) CM CO o 0 0 o to o CO CM o to o CO CM o cn ro o cn CM o CO CM O CM CM O LO CM O o CM o CO o o CM o CM O CM o cn o o ro o to 0 0 o cn o o CM O CO CM o to o o o o cn o o cn o o o o ro o o CM O o to o cn o Tl-CM O CM O o CM O CM CM O CO CM o I S -o CO o cn o LO o to o cn o Is-o o CM O CO o LO CM o ro CM O o to CM O ro ro CO ^ d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d 3 "o o CM CO ro LO CM o CM ro to LO CO CO Tf LO CO CM ro to cn rvj cn to CO to cn Is-to 0 0 CO Ti-en CO to o to o to CO LO Is-T}-to Ti- to Ti- CO LO o to cn to LO Ti-to CM CO to cn cn to CO Ti-CO LO cn Is- Is-Tt-o ro o CM CM to to to I S -I S - 0 0 LO m LO CM Tf 0 0 cn CD ro Is-LO cn Is-Si/, (rati CO 0 0 Tf Is- Tf I S - ' to Tf m ro ro ro ro Tf LO cn Ti- to ro ro ro ro CO ro ro CO CO ro ro LO cn ro ro to to Is- Ti- CO CO ro TT Tf Tf LO I S - Is- I S - CO to o * O to Is-LO to O cn cn CM o ro rvi LO Is-CM LO Is-CM o ro ro cn ro to o to ro o Is- o to Is-to Tj- to ro 0 0 O CM LO ro CO cn CO Is-o Is-Tf CM cn CM Is-Tl-to Is- Is-LO o CM ro to o LO cn LO cn cn LO LO ro en LO Tj-en CM Ti-to to I S - ro Is- o to o to Is-LO CO o cn LO CNJ to LO ro Is- CM CO LO cn CO CM CO CM ,— CM o <— CO LO t— ,— CO LO LO LO lO LO LO LO LO LO Tj- Tl- o to LO LO 0 0 cn CO CO TJ- Tj" CM ro <— I S - Is- CO LO CO 5* £ * * " CM CM LO Is-0 0 cn ro Tf CM to Tf CM to I S -Is-Tf *f o LO CM N-cn cn 'tl-o LO LO o ro q CO CO Tf cn CM CM CO cn LO Is-ro CO LO cn Is-Ti-Is-CM cn LO Tt Ti- CO Ti-Is-CO tO o ro CO to Is-Tj-q Is-0 0 I S -o CM CM o cn Is- cn CM I S - LO Tl- cn CM CO cn to CO CO to I S - to to I S -to 0 0 to o g t cd to cri LO rd to Tf to to d Is-Is- CO to rd to d LO LO ^ ' LO CO LO LO LO to LO CO LO CO LO Is-! LO d to CM to cri LO LO to to O to CM to to CM to CM to to ro to to d to rd to CO to rvj Is- to CM to cri to cri to cri to CM to cri LO cri LO CM to LO to CM to to LO to 0 0 to CO to IS -d to to to ro "? CQ Q-Q. Is- I S -CO o CM CO O ro CO CO CO o CM o LO Tt CM Tf CO to ro Tf Is-o Is-CM to o CM CM CO CM CM to Is-CM I S -r-CO Ti-ro CO CO o CM LO CO CO LO to o o r-CM O O cn Tf LO CO cn to Is-CO I S -to 0 0 ro to O Is-cn CM CO cn CO cn CO m ro CM CM o ro Tj-CO CO I S -cn cn CO CO LO 0 0 en CM CO o CM ro o 0 0 I S -Tl-Tj-CM o Ti-ro CO to r- o r--CM CM to en to 0 0 CM CM CO 0 0 cn CM CM O Tf to CO CO CO CO to CM CM LO Tf cn CO Tf r--Tf I S -cn Tf LO LO o to Tf CM LO to to CO Is-CM LO CO o CO CM LO N-ro LO LO cn CO LO t^-co CM to to LO CM CO o Tj- cn CM LO LO o CM o CO CM cn CO Is-C 0 CM O cn to CO CO CO o cn to CO cn CM cn cn 0 0 o ro cn CO cn Ti-to Is-0 0 Tj-CO oi CO CO Ti-to Tj-Is-to to ro to LO ro CO Is-LO o to cn CM to cri cn CO cn LO CO ro Tf O CO CO CO to Tf CM CM CO Is-O CO CO LO I S -cn to ro CO LO CM LO CM ro cn N -ro cn to o to Is-Is- CO o Tt cn I S -O LO CO Is-LO o LO o LO Ti- to Ol to cn LO to Tj-T T Tf CO LO CO o ^ ro CM Tf (M ro CM ro CM CM CM CM CM co CM CM ro CM cd CM CO CM rd CM CO CM 4 CM CM CM CM CM rd CM CM CM CM CM rd CM 4 CM rd CM to ro CM o cn ro LO Is-ro to CM Tf cn to CM Is-LO to CO LO I S -LO ro CO CM Is- 0 0 ro o Ti- CO CM to q CO to CO cn CO LO to CM CO cn to CO Is-cn to Ti- CO CO Is-o CM cn to o Ti-lO cn CD Ti-ro Is-0 0 CM Ti-en CO Is-CM O Ti-CD CO LO cn ro ro Tj-Is-Is-co to Is-o ro CM LO o Is-cn cn o CM Corg (rati CO CO CO cd cd CO r-! CO to Tt Is-" to r-i cd cd ed CO* CO CO Is-" cri Is-! cd to IV Is-! IV CO Is-! CO N ! Is-' cd d ed N ! CO ed s! to N ! to N ! N ! cri CO CO CO d cri h-to o cn to o Tf to o CO to o o to o CO to o to LO o cn LO o Tt o ro Tt O 0 0 o Is-LO o CO to o to o 0 0 to O CO to o LO to o 0 0 to o CO to o to LO o Ti-to o CO to o O to o CO LO O LO o CO Ti-tO CO TI-O LO O CM LO O Tj-o LO LO o cn LO o to LO o CO Is-o LO to o ro LO O to LO o cn to o o to o to o to to o o LO o LO LO o LO Ti-ro CD ro O LO LO o to to o Is-LO o to o Tf to o Is-to o LO LO o CM to o s d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d O d d d d d d d d Corg accum rate (g/cm2 ka) o CM o CM O o CM o o o CM o o o o CM o CM O CM O CM o CM o CM O CM o CM o CM o Ti-to CM o CO o CO o CO o CM o ro o CO o CO o CM O CO o CO o CO o CM O o ro o LO o CM o CM o CM O CO o CM O CM O CM o CM o CM o o o o o o CM O o Corg accum rate (g/cm2 ka) d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d to LO cn *f Is-LO CO LO LO to LO Tt Tt cn Tf CM CM CO CO CO ro cn to LO I S -LO Is-LO LO LO LO LO LO LO Tt cn LO LO Tt Tf T]- CO Tt o Tj-CO CO CO ro CO ro I S -0 0 CO LO TT Is-Tf LO Tj- to LO CM LO Tj-Tf to o LO r-Tt- LO LO LO Is-0 0 CO CO ro 0 0 to CM Tf LO LO o LO CM LO o to LO LO Is-tO o 4^  U S, d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d £ "ro ro o o o Is-O Tt o o o CM O CM O Is-o CO to CM to to to o Is- cn CM cn Tj-to Tj" CM ro CO CO cn Tf CO o o CM o Ti-to o ro o o o CM O o o o o CM o o CO o o o o en o o o o o o o CM o Tf o o o o o o o LO o O CM d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d CaC< accum (g/cm, o * Tf r--to CO Is-0 1 to Is-cn to Tf LO CM Is- to LO CM Tj-CO CO o LO LO cn CO cn CO to 0 0 Is-Is-cn Tf LO LO Is- CO CM LO LO N - LO ro Tj" I S - CM CM cn Is- o I S - CM en CO Is-o CO Ti-LO N - CM CM Is-o Tf LO ro to o to I S -LO LO m LO O w ro 3 d d CM d d d d cri to CM Is-' r-i •4 LO to rd d d d d d d d d d d d d d d d d d CM d d d d d d ^ d to d d d CM CT) o CO ro to LO Tf CO CM Tt cn Is-Is-o to CO Tf to CM cn o Is-N-O cn C 0 Is-r-Is- LO to to cn cn ro LO cn cn LO CM Ti-to CO LO cn Tf CO Is-Ti-ro Is-Tj-LO CO CO LO Is-o o to ro to to o CM CO I S -to CM o CO 0 0 to CO en CO CO cn o cn cn CO CO CO r-. 0 0 to CO LO CO Tf CO Age, Calenc (ka] ro CM CM LO CM CM r-CM CM cn CM (M CO CM ro ro CM Tf 0 0 CM LD CO CM CO CO CM o Tf CM CM (M Tj" CM LO rj-CM Is-CM cn •vt CM LO CM co LO CM LO LO CM Is-LO CM CO LO CM O to CM to CM CO to CM LO to CM to to CM I S -to CM cn to CM o Is-CM CM I S -CM ro Is-CM LO I S -CM to I S -CM r-Is-CM cn Is-CM CO CM CM CO CM Tt 0 0 CM LO CO CM Is-0 0 CM CO CO CM o cn CM cn CM CO cn CM Ti-en CM to cn CM r-cn CM cn cn CM MO ro LO o ro CO o ro ro Ti-ro Is-co Age 14C (ka) CO Tf ro Tf Tf Tf LO Tf tD Tf Is-CO Tf cn *f O Tf Tf CM •ti-ro ro LO to Is-Tj-0 0 cn TJ-O Tj" CM Ti-ro ro Ti- Ti-LO Ti-to Ti-r- Tj-cq Ti-en ro O Tj- Tf CM Ti-ro CM Tf LO Tl-tO Ti-r- TJ-0 0 Tj-cn Ti-tO Ti- Tf CM Ti-ro TJ-Tt ro LO Ti-to TJ-Is- Tf CO Tf en Tf q Tf Tf CM Tf CO Tf Tf Tf to Compos (ml to to to to to to to r-! IS-! •si r-i r-i Is-" CO cd CO CO ed CO CO ed CO CO cri cri cri cri cri cri oi ed cri cri d CM d CM d CM d CM d CM d CM d CM d CM d CM d CM CM CM CM CM CM CM +-» CD LO cn LO o LO LO CM LO ro LO Tf LO LO LO CM LO CO LO LO LO •d-to LO Is-LO CO LO cn LO o LO LO CM LO CO LO Ti- Tf LO LO CM LO ro LO Ti-LO LO Ti-to LO Is- LO CO LO cn ro o CM LO CM LO ro LO Ti- LO LO LO CM LO 0 0 to Ti-LO LO Tj-tO LO Is-LO CO LO CD LO o LO LO CM LO CO LO Tf LO to "O LU CM X ro CM X ro CM X ro CM X ro CM X CO CM X 0 0 CO X CO CO X CO CO X CO CO X CO OO X CO CO X CO CO X ro oo X CO CO X CO ro X ro ro X ro ro X ro ro X ro CO X ro co X ro X ro X CO 4 X CO 4 X ro 4 X ro 4 X ro 4 X ro 4 X CO 4 X ro 4 X CO 4 X ro 4 X ro 4 X ro 4 X ro 4 X CO LO X CO LO X ro LO X CO LO X CO to X ro LO X ro LO X ro LO X ro LO X ro LO X ro LO X ro LO X ro LO X ro to X ro LO X ro to X ro to X ro Strat SAMP Hole 1 CO I S -CO CO I CO CO CO p CO Is-CO CO 1 CO N. CO CO 1 CO I S -CO CO I CQ Is-CO CO 1 CO Is-0 0 CO g CQ I S -0 0 CO CQ Is-CO CO CQ I S -0 0 CO CO r-0 0 co CQ Is-CO CO CO r-CO CO CQ I S -CO CO CQ Is-0 0 CO CO s-CO 0 0 CO Is-0 0 0 0 CQ Is-0 0 CO CQ r-0 0 CO CO Is-CO CO CQ Is-CO CO CO I S -co CO CO Is-0 0 CO CO Is-CO CO CO I S -C 0 CO CO Is-0 0 CO CO Is-0 0 CO CO Is-0 0 CO CO I S -co CO CO Is-0 0 0 0 CQ Is-0 0 0 0 CO r-0 0 CO CO Is-0 0 0 0 CO Is-0 0 CO 1 CO r-0 0 CO I CQ r-0 0 CO 1 CQ Is-0 0 CO CO I S -CO CO 1 CO Is-0 0 CO 1 CO I S -0 0 CO 1 CO Is-0 0 CO CO Is-CO CO CO I S -0 0 CO CQ N-0 0 CO CO Is-CO CO CQ Is-CO CO CO Is-0 0 0 0 CO Is-0 0 CO I CQ Is-0 0 CO 1 CO Is-0 0 CO 1 CO Is-0 0 CO I CO r-0 0 CO CQ Is-0 0 CO ro co Fe/AI (ratio) 0.6451 0.591 | 0.608] 0.61 9] 0.591 0.536| 0.582| 0.518| 0.5851 0.616| 0.61 81 f 0.6001 0.590| 0.654| 0.8381 L O CM CD d 0.6481 0.6031 0.5631 0.5391 0.5091 0.580| 0.51 71 0.541 | 0.508.| 0.5031 0.4991 0.5551 0.4991 0.6241 0.5151 0.641 | 0.71 31 0.6681 0.7321 0.574| ] 0.547] 0.496| 0.486| 0.490| 0.5281 0.574 0.561 | •<r o CD d 0.791 | 0.6891 0.652| 0.5051 0.580| 0.5991 0.593| 0.567| ! 0.531 | Mn/Al (ratio) 0.0291 0.0241 0.025] 0.022] [ 0.012] 0.014] | 0.013] | 0.011] 0.012] | 0.014] 0.01 o] r o.oi o] | 0.011 ] | 0.013] 0.022] 0.019] 0.01 6| fO.01 5] 0.014] 0.015] [ 0.012] | 0.01 4] 0.01 o| | 0.012| | 0.01 o] 0.011 | 0.010| [ 0.0091 f 0.0091 | 0.01 o| ] 0.009] 0.010] 0.011 | ILLO-OJ 0.01 31 ] 0.010] [ 0.010] | 0.0091 | 0.0091 j 0.01 o] 0.010] ] 0.009] ] 0.01 o| | 0.0101 ( 0.01 41 | 0.015] | 0.013] 0.01 o] | 0.013] I 0.014] | 0.01 31 | 0.013 | 0.010] Ti/AI (ratio) 0.060 0.060 0.061 0.063 0.059] 0.057] 0.057 0.057 0.058 0.058 0.056 0.057] 0.059] 0.061 | 0.061 0.058 0.059 0.060 0.060 0.060 0.055] 0.055] 0.054 0.055] 0.053 0.057 0.051 0.055 0.055] 0.055] 0.055 0.058 0.061 0.058 0.062 0.061 0.058 0.056 0.053 0.053 0.056 0.056 0.058 0.060 0.062 0.061 0.060 0.057 0.057 0.063 ! 0.062 0.062 0.060 Ba/Al (ratio) 0.0391 0.0381 0.0391 0.0371 0.0121 [ 0.0171 0.0231 0.023| [ 0.020| | 0.01 71 0.011 | 0.009| ] 0.0241 0.0291 0.030| 0.01 81 0.01 6| 0.01 91 0.0201 0.0231 0.01 51 0.01 3| 0.01 o| 0.01 3] 0.01 4| 0.01 21 0.011 | 0.012] 0.01 31 [ 0.01 71 0.021 I 0.031 | 0.0271 0.026| 0.0221 0.021 | 0.011 | 0.01 4| r 0.0121 | 0.017] 0.020| 0.01 8| | 0.0231 0.030] 0.034| 0.031] 0.026] 0.01 6| 0.024] 0.027] | 0.033] ( 0.028] 0.013] < % - 1 1 0.931 9.31 | 10.031 12.10] 3.771 5.01] 8.33] 6.91 3.44] 3.51 3.20] 3.08] \ 3.35] 3.18] 3.52] •3.40] 3.42] 3.72] 3.871 4.48] 3.66 f 3.80] f 3.62] f 3.64] | 4.12] 3.77] 4.10] 4.51] 3.94] f 5.551 | 6.42] 8.33] 9.14] 7.73] 6.72 5.53 3.71 3.91 4.05 5.28 5.80 5.44 | 6.11 6.58] | 9.82 7.66 6.57 4.20 CD CM LO* 7.25 [ 7.87] ! 7.80 [ 3.59 AL203 (wt%) 5.491 CM CO CD 6.48| 5.36] 1 4.53] 11.78] 7.74) 9.00] co L O 15.02 15.88] 16.40 16.30 15.06] 14.05] 14.96 15.00] 13.68] 13.34 11.76] 15.02 14.63 15.46] 15.04] 13.94] 14.99 14.13] 12.79 14.33 10.701 9.65] 7.19 6.86] 7.95 8.79 10.47 14.71 14.49 14.37 11.87 10.53 10.92 | 10.18 9.21] 6.40 7.96 9.12 13.60 11.15 8.51 f 7.87 ( 8.04 15.07 SI02 (wt %) 67.93| 69.781 73.57| 73.43| 62.061 r66.8l| 73.001 70.391 59.721 59.731 ] 57.45| 57.24| 61.781 54.291 | 55.95| 57.64] 58.01 | 57.551 58.431 59.66| 62.1 7 ] 62.91 | 63.451 61.96| f 65.06| 63.93] 65.54| 65.36| 63.991 67.28| 70.15] 67.81 ] 70.981 69.58| 66.92] 65.52] 61.82] 64.19] 65.93] 70.99 |69.19| 67.28 | 70.43] [ 68.66] | 71.19] [ 69.04] | 67.86 64.75 CD CD CD 69.82 f 70.10] I 70.98 61.31 ™ 1 CO ci 1140] 13221 1 3241 1052] CO L O CD 1056] ^ ~ CM CD 1098] 1600] 1316| o o CD CD CD 2090] 2351 ] 2246] 1386] 1 2781 1 3631 1418] 1409] 1175] 1026] CO L O CO 1059] 1034] CD CD L O CNJ CO CO L O L O CD CD CO CD CO •3-o CM r~ 968] CO o CD O r-L O r^-CM CO L O CO o CO o CD o o CD CM O CD •3-CM CO L O L O o CO c~ CD CM o CO ^ -CM o CM CM CD CO CD O CM 1011 c ? s a 11 591 [ 1209] 1 0831 CD CO 1090| 11401 CD CD CO CO CD CD 1048] 1160] CD O CD o CD CD CO CD o NT O CO L O r~ h~ CO CD COCD CD CO CM o o CM CD CNJ CM O o CM CD CO CO CD CD r~ N -CO CM O CD CO C ~ r~ CO -3 " 00 CD CO CD o CO L O L O •3-L O CO CO Tj-CD CD CD CD CM CO CD L O CD CM CO CD CO CO CD O CD CD CD L O CO CD CO CO f^ -CM CO o CO | 1018 CD CO FN-r^ . LO CD •3-LO N . C D CO O ^ -23.81 | -23.62] -23.84] -23.24 -24.37 -23.26 -22.55 -23.21 -23.82 -23.71 | -22.71 | | -22.70 -23.05] -22.34] 1 -22.41] | -22.39] 1 -22.63] -22.52 -22.98 | -24.30 | -24.37 Corg/N (ratio) 11.25| 7.021 7.871 8.1 41 8.65| 7.72 ] 8.671 8.45] 7.34] 6.67] CD O 11.46] 7.081 5.69] 5.08] -3 -CD CD 8.62J 10.18] 7.97] 7.061 7.24] 7.30] 8.771 7.28] 8.22] 8.59] 8.17] 8.60] 8.16] 8.41 CO 8.53] 8.86] 9.24] i 9-45] | 9.59] ] 8.21] 7.78 7.77 8.1 2 1 7.85] 7.56 8.10 7.49 7.14 7.55 7.19 L O CD CD f 6.13 [ 6.07 | 6.49 I 6.97 | 6.62 0.060| 0.056I 0.058] 0.059| 0.050| 0.0551 0.0601 0.0651 0.0551 0.0551 0.065| 0.0621 0.0571 0.0531 0.052] 0.057] 0.0621 0.065] 0.0621 1 0.056| 0.052] 0.042] 0.0471 0.051 | 0.054| 0.050] 0.042] ] 0.050J 0.049] 0.0601 0.057 0.057] 0.069] 0.077] | 0.082] | 0.083] ] 0.046] 0.049] 0.042] 0.057 | 0.054] 0.061 0.081 0.089 0.077 0.076 0.071 0.061 0.060 0.069 0.066 | 0.073 0.064 Corg accum rate (g/cm2 ka) o ci o d o d o d 0.031 0.021 0.021 0.021 0.031 0.021 0.041 0.051 0.031 0.021 0.01 ] 0.02] 0.03] | 0.041 0.03] | 0.02] 0.03] 0.04] 0.051 0.041 | 0.05] | 0.03] | 0.031 1 0.03J 0.03] 0.031 0.02] 0.02] 0.02] 0.03] | 0.03] | 0.04] | 0.04] 0.03] 0.03] 0.03 | 0.02] ] 0.03] | 0.03 0.03 | 0.02 | 0.02 | 0.02 [ 0.03 j 0.02 0.02] 1 0.02J 1 0.02] | 0.03 Corg (wt %) 0.681 0.39] 0.451 | 0.48] 1 0.44J 0.42J L 0,52] 0.55| 0.40] 0.37] | 0.66| | 0.711 0.40] 0.30] 0.26] 0.38] 1 0.54] 0.66] 0.49] ] 0.39] ] . 0,371 0.30] L 0,411 0.37] 0.45] 1 0.43] | 0.35] 0-43 0.40| 0.50] 1 0.47 | 0.49] 0.61] 0.71 | 0.77 I 0.79] | 0.38] 0.38 | 0.33 0.46 0.42 ] 0.46 [ 0.66 0.67 0.55 ] 0.57 0.51 1 0.41 I 1 0.37 0.42 L 0.43J | 0.51 1 0.42 CaC03 accum rate (g/cm2 ka) 0.09! 0.07] 0.041 0.04] 0.02] 0.00] o.pol 0.04] 0.01 0.01 0.01 | 0.0 ll 0.01 | 0.16] 0.02] 1 0-01 0.01 | 0.11] 0.28 0.32] 0.01 1 0.02] 0.02] 0.01 0.01 0.01 | | 0.00] 0.00] 0.00 0.00 0.00 0.05 0.00 0.00 0.01 1 0.01 | 0.02 | 0.00] | 0.00] 0.00 0.00 [ 0.00 ] 0.01 | 0.00] 0.02 |. 0.01] 1 .... 0.01 | 0.01 ] 0.01 0.02 [ 0.03] | 0.01 | 0.01 S £ C J *J cj 5. 6.281 4.481 2.81 | 2.01 | 0.291 0.081 0.08] 0.92| 0.10] 0.12] 0.1 0] 0.09] 0.1 9| 2.571 0.38] 0.23] 0.24] 1.79] 4.57] 5.611 0.16 0.15] 0.14] 0.12] 0.06 0.20] 0.03] 0.07] 0.05] 0.04] 0.06] 1.20] 0.05] 0.12 0.20] 1 0.18 0.1 81 | 0.03] 0.00] | 0.01 0.04 1 0.03 | 0.12] | 0.11] ] 0.49] ! 0.21 | [ 0.23] 1 0.14 0.13 I 0.59 | 0.91 | 0.39 0.20: Age, ka Calendar : (ka) 320.83| 323.82| 326.81 | 329.80| 331.67] 333.1 6| 334.50] 335.85| 337.20] 338.55| 339.89] 341.24] 342.45] 343.80| 345.15] 346.50] 347.84] 349.19] 350.49] 351.70] | 353.65] 355.72] 358.151 360.59] 363.02 365.45] 365.70| 366.91 | 368.13] 369.35] CD L O O CO 371.78 373.00] 374.21 375.43 | 376.65 377.87 | 378.96 | 380.18] 381.39 382.61 383.83 385.04! 386.26 387.48 388.69 389.91 391.13 392.35 | 393.56 [ 394.78 | 396.00 ] 396.97 Age 14C (ka) Composite depth (m) 21.64] 21.74| 21.84| 21.94| 22.03| | 22.141 ) 22.24] 22.34| | 22.44] 22.54] 22.64] 22.74] 22.84] ] 22.94| ] 23.04] 23.14 23.24| 23.34] 23.44] | 23.54] | 23.72] 23.92] | 24.12] | 24.32] 24.52] 24.72] | 24.74] 24.84] | 24.94] ] 25.04] | 25.14 | 25.24 | 25.34 ] 25.44 [ 25.54 | 25.64 ] 25.74 ! 25.83 | 25.94 | 26.04 | 26.14 | 26.24 | 26.34 | 26.44 | 26.54 | 26.64 | 26.74 I 26.84 26.94 | 27.04 I 27.14 | 27.24 | 27.32 LE Ident 3H6-25 ] 3H6-35 | 3H6-45 ] 3H6-55 ] 3H6-64 ] 3H6-75 ] 3H6-85 ] 3H6-95 I 3H6-105 | 3H6-115 | 3H6-125 ] 3H6-135 ] 3H6-145 | 3H7-5 j 3H7-15 ] 3H7-25 3H7-35 | 3H7-45 3H7-55 3H7-65 3H6-72 3H6-92 3H6-112 3H6-132 | 3H7-2 3H7-22 4H1-25 | 4H1-35 4H1-45 4H1-55 4H1-65 4H1-75 4H1-85 4H1-95 4H1-105 4H1-115 4H1-125 4H1-134 4H1-145 4H2-5 4H2-15 4H2-25 4H2-35 4H2-45 4H2-55 4H2-65 4H2-75 4H2-85 4H2-95 4H2-105 [4H2-115 |4H2-125 [4H2-1 33 SAMPI Hole 1 |887B ] |887B | |887B | |887B | |887B | |887B | |887B | |887B | |887B | |887B I |887B | |887B | |887B | |887B I ]887B | |887B | |887B | |887B | |887B | |887B | |887A | |887A | |887A | ]887A | |887A | ]887A | |887B | |887B ] |887B. | |887B | |887B j |887B ] |887B ] |887B ] |887B ] I887B j |887B ] ]887B ] I887B ] I887B I887B I887B | I887B I887B I887B | I887B I887B I887B ;::::::!887B I887B I887B I887B J887B Strat 1 ;::::::!887B CO Fe/AI (ratio) 0.5981 0.51 51 0.5371 0;476| 0.5331 0.666| 0.5971 0.586| 0.6921 0.596| 0.583| 0.585| 0.541 | 0.5341 0.636| 0.554| 0.525| 0.5071 0.520| 0.6281 0.5281 0.51 51 0.5421 0.5801 0.4931 0.4971 0.560| 0.553| 0.5471 0.522| 0.491 | 0.4791 0.4631 0.5531 0.6231 0.6581 0.572| 0.507| 0.565| 0.5851 0.549| 0.431] 0.595| 0.8491 0.608] 0.5691 0.575| 0.5291 0.716| 0.7431 0.735| 0.534| 0.496| Mn/Al (ratio) 0.010 0.009 0.009 0.011 0.01 31 0.030 0.038 -0.0281 0.060 0.027 0.01 51 0.015 0.014 0.013 0.015 0.011 0.011 0.010 0.013 0.018 0.013 0.011 | 0.013 0.01 3] 0.012 0.014 0.118 0.080 0.025) 0.013] 0.009 0.011 | 0.013 0.013 0.031 | 0.014 0.015 0,012) 0.011 0.014] 0.012] 0.019] 0.018] 0.028] 0.015] 0.011 j 0.011 | 0.010) 0.018) 0.031 ) 0.072] 0.014] 0.011 ] Ti/AI (ratio) 0.0601 0.0571 0.057! 0.048) 0.051 | 0.058) 0.057) 0.054] 0.052) 0.056] 0.059] 0.057] 0.0571 0.0551 0.058] 0.058 0.056| 0.0551 0.0531 0.058] 0.055] 0.0581 0.0571 0.0581 0.054) 0.054) 0.055| 0.0581 0.061 | 0.0571 0.054J 0.053| 0.0521 0.056| 0.0631 0.061 | 0.0591 0.056| 0.0591 0.0571 0.0531 0.0471 0.0601 0.0651 0.0601 0.059| 0.058| 0.0591 0.061 | 0.0631 0.060] 0.0571 0.0551 Ba/Al (ratio) 0.009'] 0.014 0.023 0.022 0.029 0.035 0.032 0.030 0.034 0.021 0.015 0.014 0.015 0.014] 0.010 0.010 0.012 0.013 0.012 0.016 0.013 0.017 0.015 0.01 ll 0.013 0.013 0.027 0.022 0.025] 0.012] 0.012 0.012] 0.013) 0.012 0.029] 0.018 0.011 ] 0.019) 0.011 ] 0.013] 0.017] 0.018] LO CO o d 0.036] 0.026] 0.011 ] 0.013] 0.019) 0.022] 0.023] 0.026] 0.01 5] 0.0091 Si/Al (ratio) 3.221 3.69 j 3.83 j 3.94) 3.67) 3.55) 3.50] 3.62] 3.86) 3.44] 3.32] 3.38] 3.40| 3.561 3.34] 3.34 3.54] 3.59] 3.651 4.15] 3.621 4.36| 3.691 3.40| 3.95] 4.30] 15.98] 17.64| 9.471 3.62] 3.881 3.971 3.891 3.64| 1 3.081 6.15] 3.751 6.421 3.651 3.851 5.381 5.331 1 3.831 1 3.951 6.821 3.651 4.501 6.321 8.64| 1 0.721 1 0.871 4.241 3.80] AL203 (wt%) 1 5.94| 1 4.881 14.02J 14.73) 13.82) 11.37) 11.27] 11.41 | 8.33) 12.40] 15.64] 15.38) 14.78| 1 5.581 15.50| 15.77] 15.41 | 15.27| 1 5.921 13.11) 14.92) 1 3.181 15.391 15.55J 1 4.40] 12.94] 3.091 3.30| 6.331 1 15.03J 1 4.571 14.591 1 5.201 1 5.251 4.67 9.281 14.381 9.881 14.56| 14.0S | 1 2.031 10.61 | 4.91 | 4.671 9.071 1 5.271 1 3.001 9.951 7.40| 6.1 21 5.681 1 2.921 14.90| SI02 (wt%) 58.061 62.21 60.84 65.64 57.48) 45.66 44.63 46.73] 36.43 48.28 58.80 58.80 56.96] 62.77] 58.54 59.72 61.851 62.01 ] 65.87] 61.58] 61.23 65.05) 64.22] 59.86! 64.37] 63.01 55.90] 65.90] 67.88] 61.54] 64.08| 65.57) 66.90] 62.80] 69.17] 64.65] 61.09] 71.82| 60.17) 61.25) 73.34) 64.001 76.89] 73.77] 70.051 63.13| 66.23] 71.21 72.40] 74.29] 69.93) 61.98 64.061 1 CO Cl Q. CM CO c~ 11 371 1670| 1710) 21 091 2080) 1 922 j 1840) 1488) 1391 ] 1279] 11 08] 1162| 1121 ) CD "1-00 o LO CO o CD CD 1020| CO CD CD 1110] 1066) 1 2091 1217| r~ CO CO CD LO CD CO CM CD CM •a-CD CO CO CD CM CO CM CD CD CM o CD CD 1011 | CO CD CD O CO CD CO CD CD CO CO CD CO r~ CO CM CD 1084| CD CD CO CD CO CD CO 1 2321 LO CO CO CD CD CD h-CD CO LO CO 5-LO CD 1 002 ] CO CO c ? s 1 i— cn CO CD CO c~ CO CO 1^ CO o CO 2219 2650 LD co CD 3109 CD CM LO LO o LO CM o CD o r~ CO o CO o LO CD CM C~ CO CM o r~ CD LO LO CD CM CO -3- LO CM 2821 | 2086] CD LO CD LO CD CO CO o h-CO o CM CO *t CD o N-CO CO CD CO O CM 1011 | LO CD CM CD CO CD NT LO c~ CD CO c~ CO 1035] Nt CM CO Nh CD CD CD CD CO 1311 ) 2799] 1108 o CD CO CD O -23.74) -23.80] -21.91 j -21.871 -22.11 ] -24.11 ] -22.86] -21.79] -21.31| -22.73] -23.171 -23.45) -22.78] -23.48] -23.73] -22.57 -22.31 ] -23.361 Corg/N (ratio) 9.34| 6.721 5.46| 3.971 5.20| 13.51 | 5.921 5.64| 4.891 6.441 7.61 | 7.35| 6.751 6.681 6.91 | 7.83] 5.56| 7.631 CD CO CD 5.84| 6.79] 7.441 5.031 6.491 7.31 ] 8.07) 9.30| 8.1 71 8.04| 7.14| 8.851 7.82] 7.901 6.801 7.271 7.681 6.851 6.801 7.841 8.001 8.24| 9.271 8.231 7.271 8.56| 7.531 7.031 7.881 7.96| 6.751 6.631 7.20| 7.15] s 0.0751 0.0671 0.0741 0.0541 0.0551 0.0581 0.0521 0.044] 0.0491 0.060J 0.066) 0.0621 0.0651 0.061 I 0.0581 0.059) 0.051 ] 0.0491 0.0481 0.0641 .0.0491 0.0591 0.100| 0.054| 0.045] 0.044] 0.041 | 0.0451 0.0731 0.053| 0.0411 0.0381 0.041 | 0.0631 0.0581 0.053| 0.0431 0.0551 0.0551 0.055| 0.0581 0.0641 0.0591 0.0541 0.0751 0.0531 0.056| 0.0681 0.0631 0.0661 0.0591 0.051 | 0.0381 Corg accum rate (g/cm2 ka) 0.061 0.041 0.031 0.021 0.021 0.021 0.00] 0.001 0.01 | 0.01 ] 0.02] 0.021 0.021 0.021 ZO'O 0.03] 0.01 ] 0.021 0.021 0.01 j 0.02] 0.021 0.021 0.021 0.02] 0.01 ] 0.01 | 0.01 | 0.01 | 0.02] 0.02] 0.02] 0.02] 0.01 ] 0.01 | 0.01 ] 0.01 | 0.02] 0.02] 0.02) 0.01 ] 0.02] 0.01 ] 0.01 | 0.01 | 0.021 o.oil 0.021 0.01 j 0.01 | 0.01 | 0.02] 0.02] Corg (wt%) 0.70J 0.45| 0.40J 0.21 ) 0.29) 0.79) 0.30) 0.25] 0.24] 0.38] 0.50] 0.46] 0.44] 0.41 ) 0.40] 0.46) 0.28) 0.38) 0.33| 0.37] 0.33) 0.44] 0.50| 0.35! 0.33] 0.36] 0.39] 0.37| 0.59] 0.37J 0.37] 0.29] 0.33] 0.43] 0.42] 0.411 0.30| 0.37| 0.43| 0.441 0.481 0.59| CO d 0.39| 0.641 0.401 0.39] 0.53] 0.50| 0.441 0.39J 0.371 0.27| CaC03 accum rate (g/cm2 ka) 0.01 | 0.01 | 0.01 ] 0.041 0.351 0.391 0.34] 0.32] 0.81 I CO CO d 0.021 0.04) 0.1 91 0.11 0.01 ] 0.01 ] 0.03] 0.01 | 0.01 | 0.01 ] 0.05] 0.01 | 0.031 0.041 0.01 ] 0.15] 0.491 0.21 | 0.1 31 0.071 ._ 0,01 ] o.oi 0.011 0.291 0.11 | 0.09) 0.01 | 0.01 | tO'O 0.001 0.001 0.001 0.01 | 0.06| 0.01 | 0.021 0.00| 0.00] 0.001 o-oi 1 0.071 0.15] 0.01 | CaC03 (wt%) 0.1 41 0.1 21 0.11 0.53 5.13 17.05 22.70 21.36 37.81 18.25 0.46 0.85 4.69] 2.63] 0.28 0.24 0.54 0.29] 0.16 0.33 0.87 0.17 0.67 0.69] 0.19] 3.67 25.33) 13.05) 6.12] . 1,151 0.1 51 0.11] 0.27] 10.11 | 5.72] 3.12 0.11 0.11 0.13 0.09 0.08 0.10 0.68 3.10 0.33 0.43 0.04] 0.08 0.08 0.61 3.38 3.49 0.13 Age, ka Calendar (ka) 398.43| 399.52| 400.74! 401.96) 403.1 71 404.39) 407.90] 413.71 | 419.52) 423.90] 426.13] 428.37) 430.61 | 433.071 435.09| 437.33] 439.57] 441.13| 443.37| 445.61 ] 447.85] 450.09] 452.33| 453.90] 456.361 458.37] 460.84| 463.07| 465.31 ] 467.55] 469.79] 472.03] 473.60| 475.84] 478.07] 480.31 | 482.55] 484.79] 487.03] 489.27] 491.51 j 493.75| 495.99] 498.22] 500.24] 502.48) 504.72) 506.96] 509.19] 511.43] 513.60] 515.62 517.64| co <_) Composite depth (m) 27.44| 27.54| 27.641 27.74| 27.84| 27.94! 28.041 28.14| 28.241 28.34| 28.441 28.54| 28.641 28.75| 28.84] 28.94| 29.041 29.14| 29.24| 29.34| 29.44] 29.54| 29.641 29.73] 29.84| 29.93| 30.041 30.1 4| 30.241 30.341 30.44| 30.54| 30.64| 30.74| 30.84] 30.94J 31.04] 31.14| 31.24] 31.34] 31.44] 31.54| 31.64] 31.74] 31.83 ] 31.94| 32.04] 32.14] 32.24] 32.34] 32.44) 32.54] 32.64] LE Ident 4H2-145 | 4H3-5 j 4H3-15 ! 4H3-25 | 4H3-35 | 4H3-45 | 4H3-55 4H3-65 | 4H3-75 | 4H3-85 | 4H3-95 4H3-105 4H3-115 4H3-126 | 4H3-135 4H3-145 4H4-5 J 4H4-15 4H4-25 | 4H4-35 4H4-45 4H4-55 4H4-65 4H4-74 | 4H4-85 4H4-94 4H4-105 4H4-115 4H4-125 4H4-1 35 4H4-145 4HS-5 | 4H5-15 4H5-25 4H5-35 4H5-45 4H5-55 4H5-65 4H5-75 4H5-85 4H5-95 4H5-105 4H5-115 4H5-125 4H5-134 4H5-145 4H6-5 4H6-1 5 4H6-25 4H6-35 4H6-45 4H6-55 4H6-65 SAMP Hole | I887B 1 887B j 887B i 887B J 887B ) 887B j 887B 887B | 887B ] 887B ] 9/881 887B | 887B 887B J 887B ] 887BJ 887B J 887B | 887B ] 887B J 887B 887B 887B 887B | 887B | 887B J 887B 887B | 887B 887B | 887B 887B 887B 887B 887B 887B 887B J 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B 887B Strat 1 H ID CO LO Tf o oo cn LO CM LO cn Tf __ CM CO Is- I S- Tf LO to CM ro o Is- cn to Tf CM Tf cn to CM CO CO o I S- CM CM to Tf Tf LO CO Is- LO cn Is- CO CO Is- O r- cn < .2 LO o to to O to Is-to 00 CM LO LO CO o to Tf to LO to to LO to LO LO ro LO LO LO LO LO LO cn CO CO LO o to LO LO LO LO IS-LO to to O to LO to Tf r— to Tf LO r-LO to LO CO LO C0 LO CO LO CO V ro LO CM LO ro LO o to cn LO LO LO ro LO LO LO to LO LO LO CM to CM to I S-to CO Tf to LO LO to LO Q3 ro u_ o o o o o o o o o o o O o o o o o o o o o o o o o o o o o o O o o o O o O o o o o o O o o o o o o o o o o Mn/AI (ratio) o d CM O d CM o d CM O d CO o d LO o d CM O d cn o O d CO o d CM CM O d o CM o d CM CM O d cn CM O d Tf o d o d o d ro o d ro o d o o d o d to o d LO CM o d o d LO o d CM o d CO o d Tf o d LO o d ro O d CM O d CM o d o d CM o d CM o d O CM O d cn o d ro o d o o d o o d o o d o d CM o d o d o d o d CM o d Tf o d Tf o d Is-o d CM CM O d LO CM O d V CM o d CO ro O d N - o o CO CO I S- CO LO to I S- CO Tf to I S- LO r- CO r- I S- cn CO CO cn CO CM CO o LO I S- to cn CO to CM I S- to to cn I S - to LO o o Tf o cn I S- LO ro Tf < 5 LO o to o o o o ro o LO co to LO o o LO co LO CO LO co LO co LO co LO LO o LO o LO CO LO LO CO LO o to o to CO to o to o LO o LO o LO o LO o LO o LO o to o LO o to o LO o to o LO o to o LO o LO o to o to o LO o to o LO o LO o LO o LO o LO o Ti/ (rat d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d ro I S- CM CM o ro CM CM CM CM to to CO CM to CM LO CM I S-CM CO CO LO cn o ro ro to LO CM to Is- CO CO Tf CM o CM CM CM CM o CM Is- I S- cn Tf CO CM to CO r-CM Tf CO o CM CM CO CM ro CM Is-(M CM ro to to CM CM IS-m LO ro Ol o CM CO cn Ba// (rati o d o d o d o d o d o d o d O d o d o d O d O d o d o d o d o d O d o d O d o d o d o d o d O d o d o d O d o d o d o d o d o d o d o d O d o d o d o d o d o d O d O d o d o d o d o d o d o d o d o d O d o d o d cn CO LO rf to Tf CM to to cn CM o to LO LO 00 Ti-ro cn o cn to o to CO CO LO CM Tf o o CO CO CO Tf CO to Tf o LO CO Tf LO Tf CO Tf cn to CM cn CO CO cn o LO to IS-. CM to CM CO IS-Is- CM cn CM o o cn LO CM I S -to IS-LO cn Tf o CO CO V Tf LO Ol Tf CM N . Tf LO o Tf I S-m Si/A (rati< rd rd rri rri cri cd CO Tf CO ed CO CO cri Tf" rd Tf Tf" to" Tf Tf N ! CO cri Tf cri rri rri rri CO Tf cri LO d r-! Tf" rd Tf LO Tf" to" ro" rd rd rri rri rri rri rri rri rd CO o CO OO ro to LO CO CO Tf Tf CM Tf r-to to Tf cn LO Tf cn 00 o oo CM CM Tf to LO CO ro Tf ro LO o ro ro cn o LO to Tf cn CM CO cn CO cn CO o CO cn LO to o to CM ro cn ro to Tf LO CO I S-ro cn to Tf cn cn o Is-o cn cn LO o LO cn Tf cn LO CO CO Tf LO LO CO to LO Tf cri LO CM d cri cri LO LO rd ro CO CO CO to to Tf CO Tf" Tf" Tf" Tf" Tf" cri Tf ri LO Is-! LO d LO LO Tf" Tf" ri CO M CM CM d ?! * CO ro o Tf LO cn cn Is- LO CO LO I S- Tf CM O ro O cn ro to to I S - Tf CM Tf to cn CO Tf CM cn r— CM O Tf Tf Is- I S- to cn Tf o CM r— Tf O Tf CM q «- cn I S- q to cn to <Tt LO LO q r-- Tf LO to to CO cn CM CM Tf r- q cn CO q Tf Tf •-. CM •~: q q CM to q to cq q N - CO CM to to CO I S- ro T ro q O ,j <" s co" to CO LO Is-! LO LO LO CO LO to LO to LO N ! to LO Tf to Tf cri Tf r-' Tf Tf CM to d to cd to Tf to rd to LO to CO to d to CO LO d to to to to" LO to to LO LO cri LO to CM to CM to r— to IS-to LO to o I S- CM to to to to Tf to Tf to Tf to CM to o to Ol LO o to Ol LO LO LO o LO CM LO O LO Tf Tf CM Tf Is- cn «- CM ro Tf CO I S- cn 00 cn O CM o Tf CM cn CM Tf ro o cn o cn o CM r- Tf CO I S - ro to o CO cn Tf I S- CM to CM o Tf CM CO CM CO O to ro E" CQ Cl Q. o CO IS- CM to to CO o to to LO LO to O I S- o en O i — * — r ~ CM ro o cn Tf to I S- LO CO ro ro o CM o o cn CO cn CM Tf LO cn o CM to I S- Tf CM ro o I S- CO LO LO cn o cn to CO CO o o LO to cn r_ LO o ro Tf CO o to LO CO CO cn I S- o cn CO I S- LO CO Tf ,— ,— ro LO I S- CM ro ro I S- I S- cn LO CM CO Tf ro LO LO CM o o CM CO CM CO Tf CM o cn to CM CM CO s CM CO LO r- ro to CO CM CO CO IS-s S cn o o Tf CM to CM LO CO en o o cn CM CM o CO cn to cn to 00 CM to CM ro CO to CO Tf Tf Ol q ^ £ CM Tf CM CM CM CM tO to to Tf CO CM Tf to Tf r- CO CM ^f CM to o CM o CO CO CO o CO Tf CO CO to Tf CM CO LO Tf Tf CM to CO CO ro Tf CO cn ro CM cn cn Tf I S-LO CO CM to o to CM O o LO ro CO CO CO ro CO o CO to cn CO cn Tf to cn Tf CO to CO LO CO LO to to LO CO to Corg/ (ratic tb cd Is-! LO to" to" rd rd Tf LO to rd CM r-! d IS-! LO LO CO IS-! to to IS-! N ! to to" to Tf" to" tb r-! v" tb r-! Is-! CO IS-! CO IS-' IS-! IS-! r-! tb Is-! r-! Tf Tf Tf CO LO tb tb I S- Tj- Tf Ti- LO cn Is- I S- CO CO to Tf r- cn __ LO Tf o LO CO LO o to Tf cn «— CO T— CO cn O «— CO CM Tf LO cn CM I S- I S- to CO O Tf Is- to LO CM LO CO CO to o z * LO tO o to o LO Tf CM o LO LO o Tf o Tf o Tf o LO o LO o Tf o LO co to co IS-o Tf o Tf o N -o Tf CO Tf o LO o LO o Tf o Tf o Tf o Tf o to o to o IS-o to o LO o to o LO o LO o to o IS-o to o LO o Tf o LO o LO o Tf o CO o LO o Tf o CO o Tf o m o LO o d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d O d Corg accum rate (g/cm2 ka) CM o d CM o d CM o d CM O d o d CM o d o d o d o d o d o d o d CM O d ro O d CM o d CM O d CM o d CM O d CM o d o d o d CM o d CM O d o d o d o d o d o d o d CM o d CM O d CM o d o d o d o d o d CM o d CM o d CM o d CM o d o d CM o d CM o d CM o d o d o d o d o d o d to d CM o d CM o d Co" CO LO to LO I S- CO cn LO o o to o I S-^f Tf CO CM to CO to CO o CO Tf LO CO o Tf CM CO CO CM o CM CO CM ro 00 Tf LO Tf LO LO Tf CO ro I S -Tf I S-Tf Ol ro cn LO o to LO LO Tf CM CO 00 CO Tf Tf ro CO LO Tf CM CM CM I S-CM I S-ro ro ro 5 t d d d d d d d d d d d d d d d d d o d d d d d d d d d d d d d d d d O d d d d d d d d d d d d d d d d d d a a- ro o CM o cn to o LO CO to CM LO o ro O CM o o o o to o CM to CM O o o cn CM CO CM co co co o co to o CO CM Is-o ro o o o CO o CO o o o o CM o I S-o Tf CO Tf Ol CO to LO CM to CM co ro t~) f\J d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d d CaC( accum (g/cmi LO to CO cn CM CO CO CM CM to o o cn to o Is-Tf to CO LO CO CO CM CO L O o to ro Tf cn LO o CO o CM Is-CM CM CO Tf Tf Tf to to o ro to ro CO 00 Tf LO ro to Tf to LO to Ol to oo LO LO CM CM o I S-I S-CO Tf o <S o o o TJ- O Tf ro r~ cn LO o LO .- o o o o o CM to o O o Tf o O o o o O CM CO 00 co co o O O o CM o o o o ro o CO Tf co to ro a ( J CM ro ro LO to r*-to cn to o CM ro LO N - 00 o CO LO ro LO Tf CO LO CM LO to Tf o LO CO cn CM CO CM CM to o o o o CM o Tf o to o CO o o o CM o Tf O to O CO o o o CM o Tf o to o CO O o o CM o Tf O to o CO o o o CM o Tf o to CM Ol CM LO IS-o cn cn cn ^ TD ^ OJ C 0 5 cri rd LO r-! cri .- rd Tf IS-! CO d Tf LO IS-! cri CO LO IN! cri rd LO r-! cri rd tb CO d CM Tf" I V cri cri to cd d CM Tf tb cri CO LO to" oi CO Tf tb j ? ^ ^ < ( J LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO LO to LO LO LO LO LO LO LO LO LO to LO LO LO to to to to to to to to to to to to to to Age 14C (ka) CO ci o Tf Tf Tf Tf Tf Tf to to Tf Tf Tf co o CM O CO o Tf o o to o o CO o cn o o o O CM o CO o Tf o LO o to o [—. o 00 O cn o o o o CM O CO o Tf o LO o to o I S -o 00 o cn o o o O CM o ro o Tf O LO o to o s. o CO o cn o o Composi depth (m) CM CO CM CO CM ro rd ro cd ro CO ro CO CO rri ro rd ro rd ro rd ro rd CO TfCO Tf CO Tf CO Tf CO CO Tf CO Tf ro Tf CO Tf"ro Tf" ro LO ro LO CO LO CO LCO LOro m LO ro LOCO LOCO LOCO tb CO tb CO tb CO to CO tb CO to" ro tb ro tb ro tb to to ro N ! ro N ! ro IS-! ro IS-! ro N ! ro Is-! CO r-! CO r-! CO r-! CO IS-! CO cd ro LO LO LO LO LO LO LO to LO LO LO LO LO to LO C LO LO LO o «- CM CO Tf LO N. LO CM LO CO Tf LO LO LO CM LO ro LO rf LO LO LO to LO I S- LO CO LO cn o r- CM ro Tf LO LO LO CM LO CO LO Tf LO LO LO to to r-- LO 00 LO cn o CM ro Tf to LO to CM LO m LO Tf to to LO to LO I S-LO CO LO cn 2 ti) to to to to to to to r— r-L I S- v to 1 1 1 1 CM CM M CM CM CM CM CM CM CM CM CM CM CM ro CO O CO CO CO ro ro COL U _ l Tt Tf Tf Ti- Tf Tf Tf Tf Tf Tf Tf Tf Tf Lo LO LO LO LO Lo LO LO LO Lo Lo Lo LO Lo Lo LO LO LO LO LO LO LO LO LO LO LO LO LO LO to LO LO LO LO LO to to LO in LO D_ | a; CQ co co en CO CO CO CO CO CO CO CO U CO CO CO CO CO CO CQ CO CO CO CQ CO CO CQ CO CO CO CO CO CO CO CO CO CQ CO CO CO CO CO CO CO CO CO CO CO CQ CQ CQ CO CQ 5) ° CO CO CO CO CO CO CO CO 00 CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO 00 00 CO CO 00 CO CO CO CO CO CO CO CO CO 00 CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO 00 00 00 00 CO CO CO CO CO 00 CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO CO 00 CO 00 CO CO CO CO CO 00 CO • a -i • 1 1 i m ra CO I 1 1 I 1 1 1 I 1 i Fe/AI (ratio) 0.617[ 0.603| 0.505| 0.506| 0.530| 0.582| CD ro L O d 0.539| 0.536| 0.4961 0.409] 0.5771 0.530| 0.548] 0.607| 0.539| 0.476| 0.506| 0.526| 0.542| 0.543| 0.51 4| 0.4971 0.501 | 0.479| 0.526| 0.8081 0.5071 0.564| 0.6751 0.5371 0.5981 0.5621 0.6231 0.5681 0.5101 0.6421 0.6671 0.8061 0.596] 0.532] 0.5401 0.555| 0.5681 0.573] 0.5791 0.5201 0.5081 0.535| 0.557| 0.608] 0.593] 0.6401 Mn/Al (ratio) 0.028 0.017 0.017 0.017 0.015J 0.017] 0.016 0.017 0.016 0.012 0.014 0.019 0.018 r 0.014 0.017 0.014 0.017 0.013 0.012 0.011 0.012 o.oiTl 0.010 0.011 0.01 ll 0.018 0.029 0.013 0.013 0.016 0.012 0.014 0.013 0.014] 0.012 o.oTTl 0.016 0.038 0.095 0.221 0.0341 0.015] 0.014] 0.013] 0.014] 0.015] 0.014] 0.01 31 0.017] 0.030 0.043] 0.070 0.058] Ti/AI (ratio) 0.0571 0.056| 0.051 i 0.052] 0.054J 0.055! 0.053] 0.055] 0.055 0.054 0.047 0.056 0.055 0.0581 0.056 0.056 0.055] 0.054] 0.059] 0.059] 0.054] 0.054] 0.054] 0.054] 0.054] 0.054] 0.059! 0.054| 0.0591 0.059] 0.0571 0.065] 0.0591 0.0661 0.0601 0.054] 0.0591 0.0591 0.056| 0.0541 0.050] 0.0551 0.0571 0.0581 0.056| 0.0581 0.056| 0.056| 0.0551 0.0591 0.0571 0.0551 0.056| Ba/Al (ratio) 0.019 0.016 0.016 0.017 0.015] 0.016 0.018] 0.020 0.018 0.014 0.012 0.023 0.015 0.013 0.013 0.015 0.013 0.015 0.012 0.012 0.011 0.016 0.013 0.012 0.015] 0.034 0.041 0.016] 0.023 0.030 0.021 0.012 0.015] 0.023| 0.017] 0.022 0.037 0.048 0.074 0.057] 0.031 0.016] 0.014] 0.011 ] 0.01 2] 0.015] 0.019] 0.020] 0.028] 0.053] 0.046] 0.051 0.0401 Si/Al (ratio) 3.651 3.54J 3.50J 3.58] 3.53] 3.57] 3.59] 3.73] 3.73 3.73 4.07 5.20 4.01 ] 3.58] 3.89] 4.01 3.81 ] 4.691 3.60| 3.46] 3.67] 4.03] 4.1 51 3.81 | 4.03] 7.11 9.1 61 3.831 5.17 9.64| 4.59! 3.57| 4.21 | 7.70| 5.56| 4.42] 3.481 4.01 | 7.34| 5.521 4.081 3.671 3.641 3.50| 3.41 | 3.62| 4.001 4.281 6.36| 1 2.451 8.991 6.97] 5.66| AL203 (wt%) 11.691 14.06| 1 3.271 1 2.641 1 4.57] 14.30] 13.51] 1 3.30] 14.00] 14.56] 13.72] 11.061 13.67] 14.961 13.80] 13.77] 14.81 | 1 2.491 15.26| 15.49] 15.12] 1 4.241 1 4.1 61 14.90| 14.281 8.14] 6.51 | 1 4.591 11.251 6.58] 11.871 14.92J 13.27| 8.12 10.84| 1 2.87] 13.48] 10.80| 6.31 | 6.501 10.921 15.001 1 5.021 1 5.1 31 15.46] 1 4.531 1 3.771 1 3.251 9.341 4.96| e. 131 6.441 8.32| SI02 (wt %) 48.26] 56.36] 52.57| 51.28] 58.22] I 57.78J 54.86] 56.13] 59.10] 61.56] 63.17] 65.06] 62.05] 60.68J 60.80] 62.53] 63.94] 66.27] 62.20| 60.69] 62.89] 65.031 66.531 64.331 65.15] 65.53] 67.54| 63.221 65.891 71.831 61.751 60.351 63.251 70.831 68.30| 64.38J 53.1 71 49.061 52.41 | 40.641 50.44| 62.261 61.83] 59.99] 59.65| 59.531 62.431 64.221 67.271 69.931 62.40| 50.801 53.29| ro £ CO d a. 1159! 1163| 1142| 1120] 1129] 1183| 13041 1418] 1314] 1063] CD cn CO 1351 ] 1087] 1038] CO cn 1068 1022] CO cn cn L O cn 10021 L O L O CO 1 21 31 L O cn cn CO cn 11 37 ] i 1444] 14021 1 2501 1 3581 1044| 1 293] cn CM CD 1 0771 ro CD CD cn L O C D 1 4921 CD CD CD CM 2751 | 2470| r~ CD CD 1764| CO o ro CM CO - 3 -cn CD O ro CO CO CO CO CD O • 3 " u-) co CM CO CO r~ • 3 -CM L O c E S 3 2050] * 3 -CD ro r^ . cn ro u-i c~ ro • 3 -L O L O CO ro ID C J ro OO lo-co ro ID O CO cn cn m CO c~ L O o o CO C J L O o o (NJ CO ID cn cn CO cn • 3 " O o CO CVJ o C\J L O cn CD CM CO CO cn CO cn cn cn c~ CO cn 1258] 1215] CO CO CM CO CD r~ ro cn 1220| 1036| CO L O CD CO o CO L O CO CO 1311| 2572| 401 6| 92271 241 21 CO • 3 -co CO CO CM CM CO CD CD CM CO o ro CD CO - 3 -LOO L O CO CD 1 8581 2996| 3175] O ^ -24.08] -23.08] -24.1 7] . -23.671 -23.01 | -23.35] -21.33| -23.17| -23.91 | -21.96| -22.78] -23.34] Corg/N (ratio) 5.391 7.91 | 7.681 7.04| 7.37] 7.001 5,961 6.431 6.75] 6.581 5.96] 7.83] 7.081 7.98] 7.06] 7.43] 6.76| 8.26] 7.70| 8.081 8.321 7.921 7.20| 8.3.1 | 8.42J 8.15] 8.14| 7.60| 8.001 7.851 7.20| 8.191 8.181 7.86| 7.001 7.14 5.60| 5.34| 6.031 5.871 5.1 31 6.26| 6.11 | 8.401 7.82| 7.68| __ 7,29] 7.81 | 7.18| 6.81 | 6.961 5.981 CO CD' s 0.058 0.057 0.050 0.046 0.044] 0.044] 0.055] 0.056 0.061 0.052 0.033 0.063 . 0.054 0.058] 0.057] 0.053 0.055 0.060 0.050 0.044 0.040 0.044 0.056 0.046 0.048] 0.053 0.052 0.051 ] 0.069] 0.053] 0.045] 0.049] 0.050] 0.044] 0.033 0.043] 0.048 0.048] 0.041 0.038 0.036] 0.044 0.046] 0.048] 0.054] 0.062] 0.060] 0.053] 0.047 0.048 0.049 0.048 0.044] Corg accum rate (g/cm2 ka) 0.021 0.021 0.021 0.021 0.021 0.021 0,021 0.02] 0.021 0.021 0.01 I 0.021 0.02] 0.021 0.021 0.02] 0.021 0.021 0.021 0.021 0.02] 0.021 0.021 0.021 0.02| _._ o.oi 0.01 | 0.021 0.01 | 0.01 | 0.01 | 0.02J 0.021 O.OlJ 0.01 | 0.011 _. _ 0.01] 0.01 | 0.01 | 0.01 | 0.01 | 0.021 0.021 0.021 0.02] 0.021 0.021 0.02] 0.01] 0.01 | 0.01 | 0.01 | 0.01 | Corg (wt%) 0.31 j 0.45J 0.38J 0.32] 0.32] 0.31] 0.33] 0.36] 0.41 | 0.34 0.19 0.50 0.38] 0.46] 0.411 0.39 0.37 cn - 3 -d 0.39] 0.35 0.33] 0.35] 0.40 0.38 0.40] 0.43 0.42] 0.39] 0.55] 0.42] 0.33] 0.40] 0.41 j 0.35] 0.23] 0.311 0.27] 0.26] 0.25] 0.22] 0.18] 0.28] 0.281 0.40] 0.42] 0.47| 0.44] 0.41 ] 0.34] 0.33] 0.34] 0.29] 0.30] CaC03 accum rate (g/cm2 ka) 0.83 0.24 0.77 0.95 0.21 0.23] 0.48 0.40] 0.18] 0.09 0.07 0.03 0.06 0.02 0.01 0.03 0.03 0.05 0.02 0.01 0.01 0.00 0.00 0.01 0.00 0.12 0.08 0.01 0.00 0.00 0.17 0.01 0.00 0.00] 0.00 0.01 | 0.37 0.56 0.52 1.44 0.90 0.01 0.01 | 0.01 | 0.01 ] 0.02 0.07] 0.03J 0.07] 0.16 0.34] 0.86 0.71 ] CaC03 (wt%) 16.81 | 4.91 | 13.51 | 16.47] 3.91 | 4.53] 9.34] 8.1 3j 4.00] 1.94J 1.77 0.89] 1.24] 0.37] 0.15] 0.70J 0.52! 1.19j 0.321 0.1 4] 0.12] 0.10] 0.10| 0.1 01 0.10] 5.72] 3.43] 0.24| 0.1 31 0.21 | 4.42] 0.21 | 0.1 21 0.031 0.061 0.17] 9.39| 17.35] 20.091 36.481 1 9.541 0.24| 0.1 71 0.20| 0.191 0.36| 1.591 0.671 2.46| 6.90] 12.34] 25.58] 1 8.74] Age, ka Calendar (ka) 628.83| 630.75| 632.67| 634.59| 636.511 638.43] 640.351 642.271 644.22] 646.45] 648.01] 650.47| 652.70] 654.93] ID LO ID 659.39] 661.391 663.18| 665.41 | 667.64] 669.86] 672.09| 674.32| 676.11 | 678.78] 681.01j 683.24] 685.47] 687.70| 689.93| 692.16| 693.721 .695.95! 698.18| 699.52| 701.75| 703.981 706.211 708.44| 711.10| 712.38| 715.13| 71 5.871 71 7.881 71 9.721 721.55| 723.381 725.22| 727.05| 728.88| 730.72| 732.55| 734.361 Age 14C (ka) Composite depth (m) 38.10 38.20 38.30 38.40 38.50 38.60] 38.70] 38.80] 38.90] 39.00] 39.07] 39.20 39.30] 39.40] 39.50] 39.60] 39.69 39.80] 39.90] 40.00] 40.10 40.20 40.30 40.38 40.50 40.60 40.70 40.80 40.90 41.00] 41.10 41.20] 41.30 41.40] 41.50 41.60 41.70 41.80 41.90] 42.03] 42.10] 42.25 42.29 42.40] 42.50] 42.60 42.70] 42.80] 42.90] 43.00 43.10 43.20 43.30] LE Ident 5H3-105 1 5H3-115 5H3-125 5H3-1 35 5H3-145 5H4-5 | 5H4-15 | 5H4-25 | 5H4-35 5H4-45 5H4-52 5H4-65 5H4-75 \ 5H4-85 5H4-95 5H4-105 ] 5H4-114 | 5H4-125 | 5H4-135 | 5H4-145 | 5H5-5 | 5H5-15 5H5-25 5H5-33 5H5-45 | 5H5-55 5H5-65 5H5-75 5H5-85 5H5-95 | 5H5-105 5H5-115 ] 5H5-125 ] 5H5-135 ] 5H5-145 | 5H6-5 | 5H6-15 | 5H6-25 | 5H6-35 ] 5H6-48 ] 5H6-55 | 5H6-70 ] 5H6-74 I 5H6-85 5H6-95 j 5H6-105 ] 5H6-115 ] 5H6-125 J 5H6-135 5H6-145 5H7-5 | 5H7-15 5H7-25 ] Strat SAMP Hole | :::;::]887B | :j:;:;j887B | :: ;:::J887B | :::.:.]887B | :: :.::]887B | ::.:::|887B | :::.:.]887B | :;.;.;J887B | ::.:::l887B | : :: :: :]887B | CO CO CO i J887B ] |887B ] [887B ] |887B ] BIIB8S7B 1 J887B ] J887B | |887B j ]887B ] ]887B ] |887B | ]887B ] ]887B | HB887B 1 i^H887B 1 B^H887B i .|.|!!.i.i!jiJ887B | 1M:]=!=HHM 88VB ] IJHH887B I 1 1887B ] •H887B ! : o.:i887B | HH887B 1 •H887B 1 ...i...i.!...!887B ] !:i!;i!;i;i;;|887B J i.ii;!;!{.i!ii887B | 1887B | :.:.:.!887B | :; :: :: J887B | ;:;:;:]887B ] :.:::.1887B ] ::::::]887B ] : i: i: i1887B j 00 < i -< o CM -J < ° I E m co ^ o r: cm y I E Si « • "•a ' < CJ OJ U S < o rsj ro to cn CO cn LO LO o cn CSJ d d d d r-- ,_ CO OJ o o o CSJ o d d d d to LO o LO o cn CSJ o o o o d d d d <— o Tf to rsj N- o o o o o d d d d (NJ lO to LO N- OJ Tf CD LO Tf d OJ CO cn CO Csj to oq to o O rd l< CSJ ,_ r- CO CO CM CO q Tf rsj to CO to ID h- ro O CO r— to cn CO CSJ CO CO CO ro CSJ CO ro f - (SJ to OJ cn r- OJ CD CM cn ,_ Tf CO i — CSJ CO rsj t-l CD Tf rsj CSJ tD cn r - ro Tf CO LO en n! CO [•^ Tf CSJ LO CSJ o O I— o d d d d rsj ,_ o o o o «— o d d d d CO Tf CO ro T— q O d d d to CSJ o i — CM LO o d d d Tf CO o o CD o Tf" to CO d CO z EA Q AX z 2 CO 2 2 — ro g o s S « ff cn E cj Table A2.2. and 887C. Composite depth model for Hole 887B with spliced sections from Holes 887A Hole Sample ID Actual core Composite depth, m depth, m 887B 1H1-5 0.05 0.05 887B 1 HI -1 5 0.15 0.15 887B 1 H I -25 0.25 0.25 887B 1H1-35 0.35 0.35 887B • 1H1-45 0.45 0.45 887B 1H1-49 0.49 0.49 887B 1 H I -55 0.55 0.55 887B 1 H I -65 0.65 0.65 887B 1 H I -68 0.68 0.68 887B 1H1-75 0.75 0.75 887B 1 HI -85 0.85 0.85 887B 1H1-95 0.95 0.95 887B 1H1-105 1.05 1.05 887B 1H1-115 1.15 1.15 887B 1 HI -1 25 1.25 1.25 887B 1H1-145 1.45 1.45 887B 1H2-5 1.55 1.55 887B 1H2-15 1.65 1.65 887B 1H2-25 1.75 1.75 887B 1H2-35 1.85 1.85 887C 1H2-44 1.94 1.92 887C TH2-64 2.14 2.12 887C 1H2-84 2.34 2.32 887C 1H2-104 . 2.54 2.52 887C 1H2-124 2.74 2.72 887C 1H2-144 2.94 2.92 887C 1H3-4 3.04 3.02 887C 1H3-24 3.24 3.22 887C 1H3-44 3.44 3.42 887C 1H3-64 3.64 3.62 887B 2H1-5 1.93 3.69 887B 2H1-15 2.03 3.79 887B 2H1-25 2.13 3.89 887B 2H1-28 3.92 887B 2H1-35 2.23 3.99 887B 2H1-45 2.33 4.09 887B 2H1-55 2.43 4.19 887B 2H1-65 2.53 4.29 887B 2H1-75 2.63 4.39 887B 2H1-85 2.73 4.49 887B 2H1-95 2.83 4.59 887B 2H1-105 2.93 4.69 887B 2H1-115 3.03 4.79 887B 2H1-125 3.13 4.89 887B 2H1-135 3.23 4.99 Hole Sample ID Actual core Composite depth, m depth, m 887B 2H1-145 3.33 5.09 887B 2H2-5 3.43 5.19 887B 2H2-15 3.53 5.29 887B 2H2-26 3.64 5.40 887B 2H2-38 3.76 5.52 887B 2H2-45 3.83 5.59 887B 2H2-55 3.93 5.69 887B 2H2-65 4.03 5.79 887B 2H2-75 4.13 5.89 887B 2H2-85 4.23 5.99 887B 2H2-95 4.33 6.09 887B 2H2-105 4.43 6.19 887B 2H2-115 4.53 6.29 887B 2H2-125 4.63 6.39 887B 2H2-135 4.73 6.49 887B 2H2-145 4.83 6.59 887B 2H3-5 4.93 6.69 887B 2H3-15 5.03 6.79 887B 2H3-25 5.13 6.89 887B 2H3-35 5.23 6.99 887B 2H3-45 5.33 7.09 887B 2H3-55 5.43 7.19 887B 2H3-65 5.53 7.29 887B 2H3-75 5.63 7.39 887B 2H3-85 5.73 7.49 887B 2H3-95 5.83 7.59 887B 2H3-105 5.93 7.69 887B 2H3-115 6.03 7.79 887B 2H3-125 6.13 7.89 887B 2H3-135 6.23 7.99 887B 2H3-145 6.33 8.09 887B 2H4-5 6.43 8.19 887B 2H4-1 5 6.53 8.29 887B 2H4-25 6.63 8.39 887B 2H4-35 6.73 8.49 887B 2H4-45 6.83 8.59 887B 2H4-55 6.93 8.69 887B 2H4-69 7.07 8.83 887B 2H4-75 7.13 8.89 887B 2H4-85 7.23 8.99 887B 2H4-95 7.33 9.09 887B 2H4-105 7.43 9.19 887B 2H4-11 5 7.53 9.29 887B 2H4-125 7.63 9.39 887B 2H4-135 7.73 9.49 189 Hole Sample ID Actual core Composite depth, m depth, m 887B 2H4-145 7.83 9.59 887B 2H5-5 7.93 9.69 887B 2H5-15 8.03 9.79 887B 2H5-25 8.13 9.89 887B 2H5-35 8.23 9.99 887B 2H5-45 8.33 10.09 887B 2H5-55 8.43 10.19 887B 2H5-65 8.53 10.29 887B 2H5-75 8.63 10.39 887B 2H5-85 8.73 10.49 887B 2H5-88 8.76 10.52 887B 2H5-95 8.83 10.59 887B 2H5-105 8.93 10.69 887B 2H5-115 9.03 10.79 887B 2H5-125 9.13 10.89 887B 2H5-135 9.23 10.99 887B 2H5-145 9.33 11.09 887B 2H6-5 9.43 11.19 887B 2H6-15 9.53 11.29 887B 2H6-25 9.63 11.39 887B 2H6-35 9.73 11.49 887B 2H6-45 9.83 11.59 887B 2H6-56 9.94 11.70 887B 2H6-65 10.03 11.79 887B 2H6-74 10.12 11.88 887B 2H6-85 10.23 11.99 887B 2H6-95 10.33 12.09 887B 2H6-105 10.43 12.19 887B 2H6-115 10.53 12.29 887B 2H6-125 10.63 12.39 887B 2H6-135 10.73 12.49 887B 2H6-142 10.80 12.56 887B 2H7-5 10.93 12.69 887B 2H7-15 11.03 12.79 887B 2H7-25 11.13 12.89 887B 2H7-35 11.23 12.99 887B 2H7-39 11.27 13.03 887B 2H7-45 11.33 13.09 887B 2H7-55 11.43 13.19 887B 2H7-68 11.56 13.32 887A 2H6-38 14.58 13.39 887C 2H6-121 14.72 13.53 887A 2H6-61 14.81 13.62 887A 2H6-84 15.04 13.85 887A 2H6-104 15.24 14.05 887A 2H6-124 15.44 14.25 887B 3H1-55 12.15 14.44 Hole Sample ID Actual core Com posit depth, m depth, m 887B 3H1-65 12.25 14.54 887B 3H1-75 12.35 14.64 887B 3H1-85 12.45 14.74 887B 3H1-95 12.55 14.84 887B 3H1-105 12.65 14.94 887B 3H1-115 12.75 15.04 887B 3H1-125 12.85 15.14 887B 3H1-135 12.95 15.24 887B 3H1-145 13.05 15.34 887B 3H2-5 13.15 15.44 887B 3H2-15 13.25 15.54 887B 3H2-25 13.35 15.64 887B 3H2-35 13.45 15.74 887B 3H2-45 13.55 15.84 887B 3H2-55 13.65 15.94 887B 3H2-64 13.74 16.03 887B 3H2-75 13.85 16.14 887B 3H2-85 13.95 16.24 887B 3H2-95 14.05 16.34 887B 3H2-105 14.15 16.44 887B 3H2-115 14.25 16.54 887B 3H2-125 14.35 16.64 887B 3H2-135 14.45 16.74 887B 3H2-145 14.55 16.84 887B 3H3-5 14.65 16.94 887B 3H3-15 14.75 17.04 887B 3H3-25 14.85 17.14 887B 3H3-35 14.95 17.24 887B 3H3-45 15.05 17.34 887B 3H3-55 15.15 17.44 887B 3H3-64 15.24 17.53 887B 3H3-75 15.35 17.64 887B 3H3-85 15.45 17.74 887B 3H3-95 15.55 17.84 887B 3H3-105 15.65 17.94 887B 3H3-115 15.75 18.04 887B 3H3-125 15.85 18.14 887B 3H3-135 15.95 18.24 887B 3H3-145 16.05 18.34 887B 3H4-4 16.14 18.43 887B 3H4-15 16.25 18.54 887B 3H4-25 16.35 18.64 887B 3H4-35 16.45 18.74 887B 3H4-45 16.55 18.84 887B 3H4-55 16.65 18.94 887B 3H4-64 16.74 19.03 887B 3H4-75 16.85 19.14 190 Hole Sample ID Actual core Composite depth, m depth, m 887B 3H4-85 16.95 19.24 887B 3H4-95 17.05 19.34 887B 3H4-103 17.13 19.42 887B 3H4-112 17.22 19.51 887B 3H4-125 17.35 19.64 887B 3H4-135 17.45 19.74 887B 3H4-145 17.55 19.84 887B 3H5-5 17.65 19.94 887B 3H5-15 17.75 20.04 887B 3H5-25 17.85 20.14 887B 3H5-35 17.95 20.24 887B 3H5-45 18.05 20.34 887B 3H5-55 18.15 20.44 887B 3H5-64 18.24 20.53 887B 3H5-75 18.35 20.64 887B 3H5-85 18.45 20.74 887B 3H5-95 18.55 20.84 887B 3H5-105 18.65 20.94 887B 3H5-115 18.75 21.04 887B 3H5-125 18.85 21.14 887B 3H5-135 18.95 21.24 887B 3H5-145 19.05 21.34 887B 3H6-5 19.15 21.44 887B 3H6-15 19.25 21.54 887B 3H6-25 19.35 21.64 887B 3H6-35 19.45 21.74 887B 3H6-45 19.55 21.84 887B 3H6-55 19.65 21.94 887B 3H6-64 19.74 22.03 887B 3H6-75 19.85 22.14 887B 3H6-85 19.95 22.24 887B 3H6-95 20.05 22.34 887B 3H6-105 20.15 22.44 887B 3H6-115 20.25 22.54 887B 3H6-125 20.35 22.64 887B 3H6-135 20.45 22.74 887B 3H6-145 20.55 22.84 887B 3H7-5 20.65 22.94 887B 3H7-15 20.75 23.04 887B 3H7-25 20.85 23.14 887B 3H7-35 20.95 23.24 887B 3H7-45 21.05 23.34 887B 3H7-55 21.15 23.44 887B 3H7-65 21.25 23.54 887A 3H6-72 24.42 23.72 887A 3H6-92 24.62 23.92 887A 3H6-112 24.82 24.12 Hole Sample ID Actual core Composite depth, m depth, m 887A 3H6-132 25.02 24.32 887A 3H7-2 25.22 24.52 887A 3H7-22 25.42 24.72 887B 4H1-25 21.53 24.74 887B 4H1-35 21.63 24.84 887B 4H1-45 21.73 24.94 887B 4H1-55 21.83 25.04 887B 4H1-65 21.93 25.14 887B 4H1-75 22.03 25.24 887B 4H1-85 22.13 25.34 887B 4H1-95 22.23 25.44 887B 4H1-105 22.33 25.54 887B 4H1-115 22.43 25.64 887B 4H1-125 22.53 25.74 887B 4H1-134 22.62 25.83 887B 4H1-145 22.73 25.94 887B 4H2-5 22.83 26.04 887B 4H2-15 22.93 26.14 887B 4H2-25 23.03 26.24 887B 4H2-35 23.13 26.34 887B 4H2-45 23.23 26.44 887B 4H2-55 23.33 26.54 887B 4H2-65 23.43 26.64 887B 4H2-75 23.53 26.74 887B 4H2-85 23.63 26.84 887B 4H2-95 23.73 26.94 887B 4H2-105 23.83 27.04 887B 4H2-115 23.93 27.14 887B 4H2-125 24.03 27.24 887B 4H2-133 24.11 27.32 887B 4H2-145 24.23 27.44 887B 4H3-5 24.33 27.54 887B 4H3-15 24.43 27.64 887B 4H3-25 24.53 27.74 887B 4H3-35 24.63 27.84 887B 4H3-45 24.73 27.94 887B 4H3-55 24.83 28.04 887B 4H3-65 24.93 28.14 887B 4H3-75 25.03 28.24 887B 4H3-85 25.13 28.34 887B 4H3-95 25.23 28.44 887B 4H3-105 25.33 28.54 887B 4H3-115 25.43 28.64 887B 4H3-126 25.54 28.75 887B 4H3-135 25.63 28.84 887B 4H3-145 25.73 28.94 887B 4H4-5 25.83 29.04 191 Hole Sample ID Actual core Composite depth, m depth, m 887B 4H4-1 5 25.93 29.14 887B 4H4-25 26.03 29.24 887B 4H4-35 26.13 29.34 887B 4H4-45 26.23 29.44 887B 4H4-55 26.33 29.54 887B 4H4-65 26.43 29.64 887B 4H4-74 26.52 29.73 887B 4H4-85 26.63 29.84 887B 4H4-94 26.72 29.93 887B 4H4-105 26.83 30.04 887B 4H4-1 15 26.93 30.14 887B 4H4-125 27.03 30.24 887B 4H4-135 27.13 30.34 887B 4H4-145 27.23 30.44 887B 4H5-5 27.33 30.54 887B 4H5-15 27.43 30.64 887B 4H5-25 27.53 30.74 887B 4H5-35 27.63 30.84 887B 4H5-45 27.73 30.94 887B 4H5-55 27.83 31.04 887B 4H5-65 27.93 31.14 887B 4H5-75 28.03 31.24 887B 4H5-85 28.13 31.34 887B 4H5-95 28.23 31.44 887B 4H5-105 28.33 31.54 887B 4H5-115 28.43 31.64 887B 4H5-125 28.53 31.74 887B 4H5-134 28.62 31.83 887B 4H5-145 28.73 31.94 887B 4H6-5 28.83 32.04 887B 4H6-1 5 28.93 32.14 887B 4H6-25 29.03 32.24 887B 4H6-35 29.13 32.34 887B 4H6-45 29.23 32.44 887B 4H6-55 29.33 32.54 887B 4H6-65 29.43 32.64 887B 4H6-75 29.53 32.74 887B 4H6-85 29.63 32.84 887B 4H6-95 29.73 32.94 887B 4H6-105 29.83 33.04 887B 4H6-115 29.93 33.14 887B 4H6-125 30.03 33.24 887B 4H6-135 30.13 33.34 887B 4H6-145 30.23 33.44 887B 4H7-5 30.33 33.54 887B 4H7-17 30.45 33.66 887B 4H7-25 30.53 33.74 Hole Sample ID Actual core Composite depth, m depth, m 887B 4H7-35 30.63 33.84 887C 4H6-41 31.21 34.04 887B 5H1-5 30.73 34.10 887B 5H1-15 30.83 34.20 887B 5H1-25 30.93 34.30 887B 5H1-35 31.03 34.40 887B 5H1-45 31.13 34.50 887B 5H1-55 31.23 34.60 887B 5H1-65 31.33 34.70 887B 5H1-75 31.43 34.80 887B 5H1-85 31.53 34.90 887B 5H1-95 31.63 35.00 887B 5H1-105 31.73 35.10 887B 5H1-115 31.83 35.20 887B 5H1-125 31.93 35.30 887B 5H1-135 32.03 35.40 887B 5H1-145 32.13 35.50 887B 5H2-5 32.23 35.60 887B 5H2-15 32.33 35.70 887B 5H2-25 32.43 35.80 887B 5H2-35 32.53 35.90 887B 5H2-45 32.63 36.00 887B 5H2-55 32.73 36.10 887B 5H2-65 32.83 36.20 887B 5H2-75 32.93 36.30 887B 5H2-85 33.03 36.40 887B 5H2-95 33.13 36.50 887B 5H2-105 33.23 36.60 887B 5H2-115 33.33 36.70 887B 5H2-125 33.43 36.80 887B 5H2-135 33.53 36.90 887B 5H2-145 33.63 37.00 887B 5H3-5 33.73 37.10 887B 5H3-15 33.83 37.20 887B 5H3-25 33.93 37.30 887B 5H3-35 34.03 37.40 887B 5H3-45 34.13 37.50 887B 5H3-55 34.23 37.60 887B 5H3-65 34.33 37.70 887B 5H3-75 34.43 37.80 887B 5H3-85 34.53 37.90 887B 5H3-95 34.63 38.00 887B 5H3-105 34.73 38.10 887B 5H3-115 34.83 38.20 887B 5H3-125 34.93 38.30 887B 5H3-135 35.03 38.40 887B 5H3-145 35.13 38.50 Hole Sample ID Actual core Composite depth, m depth, m 887B 5H4-5 35.23 38.60 887B 5H4-15 35.33 38.70 887B 5H4-25 35.43 38.80 887B 5H4-35 35.53 38.90 887B 5H4-45 35.63 39.00 887B 5H4-52 35.70 39.07 887B 5H4-65 35.83 39.20 887B 5H4-75 35.93 39.30 887B 5H4-85 36.03 39.40 887B 5H4-95 36.13 39.50 887B 5H4-105 36.23 39.60 887B 5H4-114 36.32 39.69 887B 5H4-125 36.43 39.80 887B 5H4-135 36.53 39.90 887B 5H4-145 36.63 40.00 887B 5H5-5 36.73 40.10 887B 5H5-15 36.83 40.20 887B 5H5-25 36.93 40.30 887B 5H5-33 37.01 40.38 887B 5H5-45 37.13 40.50 887B 5H5-55 37.23 40.60 887B 5H5-65 37.33 40.70 887B 5H5-75 37.43 40.80 887B 5H5-85 37.53 40.90 887B 5H5-95 37.63 41.00 887B 5H5-105 37.73 41.10 887B 5H5-115 37.83 41.20 887B 5H5-125 37.93 41.30 887B 5H5-135 38.03 41.40 887B 5H5-145 38.13 41.50 887B 5H6-5 38.23 41.60 887B 5H6-15 38.33 41.70 887B 5H6-25 38.43 41.80 887B 5H6-35 38.53 41.90 887B 5H6-48 38.66 42.03 887B 5H6-55 38.73 42.10 887B 5H6-70 38.88 42.25 887B 5H6-74 38.92 42.29 887B 5H6-85 39.03 42.40 887B 5H6-95 39.13 42.50 887B 5H6-105 39.23 42.60 887B 5H6-115 39.33 42.70 887B 5H6-125 39.43 42.80 887B 5H6-135 39.53 42.90 887B 5H6-145 39.63 43.00 Hole Sample ID Actual core Composite depth, m depth, m 887B 5H7-5 39.73 43.10 887B 5H7-15 39.83 43.20 887B 5H7-25 39.93 43.30 887B 5H7-35 40.03 43.40 887B 5H7-45 40.13 43.50 887B 5H7-55 40.23 43.60 193 CD • — * LO * o o 8 S <-< CO CD O 5 •cr U D 5 O CJ + o r^ 00 •cf o — CD O <3-CO CO o § 00 •=r d d d o o d o d o d o d o CO d d •3-o d o o o d CO d o d d cn d o d CO d o d o d 5 ® o d o d o d o d oq d o d d d CO cb cn d CO o § CJ) CO < U 1; o (O CO CO co d co u cn n-CO E o -4J co a. -a co I hn CD OJ e g co S CJ) C\J i X CD cn < Local (pl.-ben.) -1.36 -1.73 -1.21 -1.28 -1.12 -1.04 -1.38 -1.31 -1.37 -1.44 -0.94 -0.54 -0.84 -0.39 -1.90 -1.45 o • -0.51 -0.85 -0.94 -1.85 -1.68 -1.84 -1.40 -1.59 I Composite Benthic** 4.20 4.35 4.09 3.90 3.98 4.25 4.00 3.74 3.73 3.61 4.27 4.27 3.73 3.07 3.20 3.06 4.51 4.72 4.65 4.68 4.62 | 4.69 Cibicid +0.64 4.20 4.35 4.09 3.90 3.98 4.25 3.68 4.27 4.27 4.56 4.65 4.62 4.55 4.62 CD id 3 3.63 3.73 3.07 3.20 3.06 4.60 4.91 4.67 4.81 4.69 d18C Gyroid 1 4.00 3.81 3.73 3.60 4.36 4.60 4.66 G. bull 1 2.84 2.62 2.88 2.74 2.80 3.21 | 2.87 2.89 1 2.74 | 2.61 3.06 3.20 | 2.77 3.88 2.37 2.80 3.16 3.22 2.21 3.57 2.87 2.97 2.85 | 3.22 | 3.10 Cibicid -0.50| -0.49| -0.82| -0.53| -0.69] -0.86| -0.38I -0.75 -0.81] -0.90 -0.84 -1.30 -0.73 -0.70 t_ D) -1.13 -1.16 -0.70 -0.93 -0.85 -1.27 -1.40 -1.45 -1.37 -1.36 o , CO *5 2 >v 00 -1.21] -0.97] -1.06] -1.35 -1.54 -1.63 G. bull | 0.051 0.231 0.011 0.111 -0.041 0.18| 0.241 0.131 0.211 0.311 0.16| 0.10| 0.021 0.39] 0.32 0.42 0.36 0.11 | -0.31 | -0.25 | -0.24 | -0.33 | 0.03 I -0.26 | -0.42 G. bull 16.01 30.8|| CO NT CO CO d 55.8|| d o d o d o d o d o d o d o d CNI d 58.0|! 48.7| 477.5| CD d 187*. 11 CT> I N ! o od o d CO o d 0-00 CD 31.1| oq 157.9| 74.8| d CTl CO q d o d o d CO CO d 43.3 1163.6 1139.9 309.4 153.9 722.6 (#/g) Planktic N. pach 1 en d CO LO CO ID I N ! q d CO CO o d o d o d o d o d o d o d o d o d 27.81 CD od 45.51 CNJ d 20.4| 38.6| to CO o d d q d 11.7 tt 13.1 21.61 26.8 38.2 o d CO d o d CNJ d co o d 90.5 1507.8 2480.5 358.2 894.6 2328.9 ndance Cibicid d (NJ d d O d o d o d q d q d q d o d O'O o d q d cn d CO CNJ rN. o d o d o d q d o d d o d 00 CNJ d CNJ d CNJ d tn CNJ o d d d CNI d q d CNI d Nf d CNI d q d oq d CO 00 d d CO -Abu Benthic Uviger 1 o d o d o d O d o d o d q d o d o d o d o d o d o d o d O'O o d o d o d o d o d o d CO d o d o d d CD d o d O d o d o d o d o d CO d o d oq q CNJ LO o d q CO CO CO Nt; °! CNJ o Gyroid 1 q d o d o d O'O q d q d o d o d q d o d q d o d o d o d o d q d o d o d o d q d q d to d o d LO d q cn d o d o d q d CNJ d d o d •cr d q d CO d d "3" d O'O q CNJ d •3-d O d d I 0.2 Q DC LO LO LO to <D LO LO - ( N J CO NT LO LO If ••3- m to to LO LO CO - CNJ LO CO CO LO CO tn CO in CNJ in CNJ N* CD LO to •cr CO in oJ CO •Ll- LO LO CNJ CNJ Age* Cal. (Ka) 69.41 71.09 72.00 73.09 74.00 74.82 75.36 76.27 77.36 78.36 79.27 80.18 81.09 82.00 82.91 84.27 85.68 87.09 88.36 89.91 91.32 92.73 95.43 98.03 99.32 100.62 104.93 105.59 to CNJ ID o 106.92 107.58 112.88 114.87 1 18.20 120.40 122.40 125.60 127.60 129.25 130.82 132.38 133.79 135.35 136.29 Age* 14C (Ka) Compos. Depth (mbsf) 4.74 4.84 4.94 5.06 5.16 5.25 5.31 5.43 5.55 5.66 5.76 5.86 5.96 6.06 6.16 6.26 6.36 6.46 6.55 6.66 6.76 6.86 7.06 7.26 7.36 7.46 7.86 7.96 8.06' 8.16 8.26 9.06 9.36 9.56 9.96 10.06 10.26 10.36 10.46 10.56 10.66 10.76 10.89 10.95 Sample Identifier B 2H1-110 B 2H1-120 B 2H1-130 B 2H1-142 B 2H2-2 B2H2-11 B2H2-17 B 2H2-29 B 2H2-41 B 2H2-52 B 2H2-62 B 2H2-72 B 2H2-82 B 2H2-92 B 2H2-102 B 2H2-112 B 2H2-122 B 2H2-132 B 2H2-141 B 2H3-2 B2H3-12 B 2H3-22 B 2H3-42 B 2H3-62 B 2H3-72 B 2H3-82 I B 2H3-122 B 2H3-132 B 2H3-142 B 2H4-2 B 2H4-12 B 2H4-92 B 2H4-122 B 2H4-142 B 2H5-32 B 2H5-42 B 2H5-62 B 2H5-72 B 2H5-82 B 2H5-92 B 2H5-102 B 2H5-112 B 2H5-125 B 2H5-131 CO CD CO 00 CO CM 00 •cf CD CD to o CO LO CD LO CO •cr 00 CO O r^ CD 00 CD CO O q LO CM r^ CD CO CD to r^ CO CO •cr CM •cr CD (NJ CM LO Lo (pl.-t d q" ' • ' • • d q" d q q q CM • o o o CD o CD CM LO LO CD LO CD CO O CNJ CO CM M" •cr CO CM LO CO 0 CD CM LO LO CD CD LO •cr LO CM c o Benth Nt N+ •cr Nf NT •Cf Nt r^ •cr CO CO •* Nt N+ Cibicid 1 +0.64 4.90 4.60 4.32 4.07 3.60 3.52 2 Uviger 4.41 4.59 | o CO Gyroid 4.48 4.52 4.51 4.67 4.31 4.40 4.42 4.44 4.31 4.35 4.57 4.69 4.45 | 4.51 4.42 I 4.34 bull 1 CD CM CD CD CD CD •cr N+ CO 00 o IN-CD CM CD 00 1-00 CO o q q CO CD •cr 00 CD CO 00 q r - 00 CO CD to CO CO 0 h~ CO CM LO LO CO 0 r o CM 00 r o 00 e> CO CO CO CO CNJ CM CO CM CM CM CO CO r o CO r o c o r d r o CO r o CM CM CO CM CO r o CO CM !2 CM CD CD O cn 0 00 LO 00 q CJ l a q d d d 1 d "CJ g Uviger | -1.12| -1.24| I3C CO •cr CM o CO CM LO CD LO CD O |N-CM tN. LO LO •d- to I D CM h~ LO LO O IN. •*> Gyre " 7 ' 7 T 1 1 1 1 1 1 = CD LO 5- CO o CO CO CM 1^  CO LO CM 00 O CM 00 CM |N- •cr CO 0 CM CO CD 0 CM Nt CNI CD CM Nt CD CO r o 00 LO CO LO X ! q q d i d d 1 q d q" d t d d 1 d 1 d • d d 1 d d d d d d d d d 1 d d q CO o o o CM j— ,— 00 tn 00 r-. IN. o •cr I D •cr 0 0 Nf r o O 0 t~ 00 T— O 0 0 LO (NJ 0 0 0 q CD CM CO CNJ M- CM 00 r o G. bull LO LO CD CD d d d CM 1033. d CD d CD CM d CM d CM o CM LO LO, CO IN. d CO CO d CM d d d CD CO d d d d d d CM CNJ d d d CO d CD CM CM O r^ r^ LO CO 00 •cr CM NT O CM Nf NT r~ C) LO CO LO o CD CM 00 CO CO LO o CO LO q •cr IN. CO 0 LO LO CM CO r- 0 0 q CO 0 0 0 00 LO CM LO q CD In. CO CO 2- Planktii N. pach IN! CM CD 1^  d d d CM 1235. CD O 00 CNJ 00 LO d d CM CD CNJ CNJ LO LO CO r o r o IN. d to to d d d CD LO d d CO d d d d | N ! CM d d d LO r o LO CO •cr •cr LO CD CD IN! •cr 00 CM IN! 00 cu CD ,— CNJ o CO q CNJ ,— O •cr o q t^ LO 00 0 0 0 0 0 CO CM CO 0 0 0 0 O 0 0 • - 0 q — CNJ CM CO 00 I D CD CO indanc icid O d d d • - d d d d d •-" d d d d 0 d 0 d d d 0 d 0 d d d 0 d d CM d d d d O O d indanc Cib CJ CO o o o o LO o o o o o o q r- O O 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 q CO O 0 0 0 — O 0 < Benthi Uviger d d d d d d d d d d d d d d d d d d d 0 d d d d d 0 0 d d d 0 0 0 d d d d d d d d d d d Benthi Uviger 00 CO 00 o CNJ q CM o o LO 10 CM q 0 ,— 0 N+ 0 0 q CM 0 0 0 0 0 0 0 0 0 LO LO CD CD r o LO Gyroid d d d d d d d d d d d d d d d d 0 d d d d d d d d d d d d d d d d d d d d O d IRD M- CM CM LO CO LO - LO CM LO CM to LO LO CM LO CM •cr CO CO LO LO CO CM LO CM LO CO LO r d LO r d CO CO LO CM LO CM LO CM 10 CO LO r o CO CO CD q 00 LO o CO CD CO r -00 r^ o q <£> CM •cr CM N -CO o LO o 00 r^ 00 q CM r~ N+ O CO CO CD (NJ IN. O q to CD CM CO •cr CNJ 0 CD CD q CO (Nl •cr CD CO In. CD CO (NJ CM LO q CO 00 r i LO CD CNI CO to CM CD r o CO q 0 q CO CO CD CO ? CM N+ r-! •o- CD •cr LO CM LO LO IN! LO CD LO CD CO CO LO CD CO 00 CD d CO N . LO In. CD IN. CO CO to" 00 d CD CO CD CD CD •crCM 0 (NJ CM r d r o r o CO CM CM 0 r o CM CO CM co" CO CM 00 CO CM CD CO CM M-CM r d 1-( M LO ( M IN! •cr r o D t CM LO CM CO LO CM LO LO CM IN! LO CM CD CO < CJ Age* 14C(Ka) If) CO q CO CD CO CO CO LO 00 CD CD CD o CO LO CM CO •cr CD LO CM In. CO 00 CO CD CD q to •cr CM CD CO CM q LO 00 LO 0 LO ( M LO r^ LO CO CO IN. CO CD IN. CD 0 r— CM CO *t LO to IN. 00 CD 0 — O . x : 4J to CM CN] CM CM CM CM CO CM CO CO CO CO r o CO •cr •cr •^ LO CD co CO CD CD IN! IN! IN! IN! IN! r-! In! IN! IN! IN! 06 00 Com Q . - a CD c Q £ Sample Identifier CM •CT CM LO CNJ CM IN. CM 00 CNJ CD CM O CM CO LO •cr 00 CM CM CM CO CM •cr CM LO O CO 00 CO to •cr 00 r^ 0 r^ 00 Nt •cr (Nl CD CM IN. CM CD 00 CO CM r~ CM CD CM r o CM r o CNJ (NJ CM CM CM r o CM CM LO to r o h- CM 00 CM CD -102 -112 -122 Sample Identifier 2H5-2H6-2H6-2H6-2H6-2H6-2H6-2H6-2H6-2H6-2H6-2H7-2H7-2H7-2H7-2H7-2H7-2H6-2H6-2H6-2H6-2H6-2H6-3H1-3H1-SHI-3H2-3H2-3H2-3H2-3H2-3H3-3H3-3H3-3H3-3H3' 3H3' 3H3 3H3 3H3 3H3 3H3 3H3 3H3 Sample Identifier CO 00 c o o o 00 00 o a c o 00 00 OQ 0Q CQ 00 < < < < < < CO CD 00 CO 00 CO CO CO CO CQ 00 CO 00 00 00 00 CO 00 0 3 CO 0 CJN Local (pl.-ben.) -1.55 -1.29 -1.52 -1.41 -1.43 -1.16 -1.55 -1.41 00 -1.69 -1.63 -1.55 -1.14 -1.28 -1.54 -0.60 Composite Benthic** 4.53 4.50 4.35 4.40 4.19 4.28 3.88 3.77 4.70 4.69 Cibicid 1 +0.64 4.69 4.35 4.40 4.19 3.77 cu CO d18C Gyroid 1 4.53 ] 4.30 | 4.28 3.88 4.70 4.69 G. bull 1 2.99 | 3.22 | 2.98 | 3.07 | 2.97 | 3.03 | 2.72 | 2.62 | 2.81 2.26 2.24 | 2.22 | 3.56 3.41 3.15 3.02 Cibicid -1.751 -0.57| -1.07| -0.56| -0.731 (%») Uviger | co "o *5 2 >v CJ3 -1.64| -1.76| -1.07| -1.64| -1.55 -1.45 G. bull | -0.56| -0.491 -0.601 -0.59| -0.26| -0.11 0.001 -0.08| -0.27| -0.491 -0.16| -0.061 -0.37| -0.35 -0.23 | -0.20 G. bull 415.3|| . 430.1|| 171.0| 72.5 o d |00 CO d d d d 423.7] 163.91 d CO d 00 CO 11.0| 28.0| 84.7| 142.2| m d o d o d o d q d CO d 18.3| 34.9| . 85.3| o d o d o d o d o d CO CNJ d q d d o d CNJ d o d CO CNJ O'O | o d Planktic N. pach | 575.61 522.71 333.0| 42.9) o d o d CO d d o d o d 493.2] 141.5| q d d CO d CO CNJ 00 CO in T f 21.7] o d o d o d o d o d JN-d 11.91 51.2: 28.4 O'O | d o d o d o d CO d o d o d o d o d o d o d LO d o d o d ndance Cibicid 00 d T f d CNJ CNJ d O'O o d q q d o d o d LO CO d CNJ d d o d CNJ d O d d o d CO d o d o d o d o d o d o d d d o d o d o d o d o d o d d o d o d ro d o d o d d OJ d o d o d Abu Benthic Uviger 1 o d o d o d o d o d o d o d q d o d o d o d o d o d q d o d o d o d o d o d q d o d O'O o d o d q d o d o d o d o d q d o d o d O'O | o d q d q d q d o d o d o d q d o d o d o d Gyroid 1 CO d T f d ro d o d o d p d o d o d o d o d d o d o d o d o d o d o d T f d CO d d d q d q d o d o d o d CNJ d q d CO d o d o d o d o d q d o d o d o d O'O q d o d o d q d o d o d IRD in CNJ in CNI - - ro CNJ LO (NJ in m T f in (NJ CO m CO T f m rj LO LO CNJ in CNJ OJ LO LO m m CNJ OJ LO OJ CO CO in CO ro in T f LO T f m CO Age* Cal. (Ka) 258.531 260.07 261.60 263.13 267.12 270.03 280.61 283.67 288.27 297.47 299.00 304.98 307.97 310.96 316.94 319.93 322.93 325.92 328.91 331.00 335.45 343.40 344.74 346.09 347.441 348.79 350.12 351.34 352.56 356.08 357.67 360.10 362.53 364.97 371.42 375.07 387.11 388.33 389.55 393.20 394.41 395.63 396.60 401.59 Age* 14C (Ka) Compos. Depth (mbsf) CNJ 00 CO CO od LO 00 CO CO 19.00 |N-cri CT> 20.21 20.81 20.91 CNJ CNJ CNJ CO CNJ in CNI CD CNJ OJ oq OJ CXI CNJ 21.98 22.31 22.91 23.01 23.11 23.21 23.31 23.41 23.51 23.63 23.95 24.08 24.28 24.48 24.68 25.21 25.51 26.51 26.61 26.71 27.01 27.11 27.21 27.29 27.71 Sample Identifier B 3H3-132 B 3H3-142 B 3H4-2 B3H4-12 B 3H4-42 B 3H4-61 B 3H4-132 B 3H5-2 B 3H5-32 B 3H5-92 B 3H5-102 B 3H5-122 B 3H5-132 B 3H5-142 B 3H6-12 B 3H6-22 B 3H6-32 B 3H6-42 B 3H6-52 B 3H6-59 B 3H6-92 B 3H7-2 B3H7-12 B 3H7-22 B 3H7-32 B 3H7-42 B 3H7-52 B 3H7-62 A 3H6-63 A 3H6-95 A 3H6-108 A 3H6-128 A 3H6-148 A 3H7-18 B 4H1-72 B 4H1-102 B 4H2-52 B 4H2-62 B 4H2-72 B 4H2-102 B 4H2-112 B 4H2-122 B 4H2-130 B 4H3-22 CN Local (pl.-ben.) -1.13 1 -1.15 | -0.98 -1.19 -1.36 -0.52 -0.88 | -1.79 -1.52 -1.52 -1.69 -1.13 -1.13 -1.45 | -1.44 -1.32 -1.53 -0.92 -0.71 -0.90 -1.12 -1.24 -1.53 -1.08 CO -1.10 | Tf 1 Composite Benthic** 3.41 1 3.08 | 3.28 1 3.82 | 5.01 4.95 | 4.56 4.18 | 4.35 4.12 3.86 4.03 3.85 4.10 4.08 4.16 4.01 Cibicid +0.64 3.46 | 3.32 | 4.95 | 4.18 | 4.35 | 4.12 3.86 ] 4.03 3.91 4.10 4.02 3.98 >(%.) Uviger 3.36 | 3.08 | 3.25 1 3.78 4.15 4.16 4.04 d18C Gyroid | 3.82 | 5.01 | 4.56 | G. bull | 2.28 | 1.94 i 2.31 1 2.42 1 2.46 3.68 | 3.59 | 3.23 1 3.43 1 3.35 | 2.87 3.16 3.05 | 2.73 | 2.78 | 2.95 j 2.83 2.90 ] 2.95 | 3.22 2.75 2.80 2.32 3.02 2.78 3.06 2.60 Cibicid -0.42| -0.44| -1.401 -0.89| -0.91| -0.66| -0.42] -0.39] -0.30] -0.48] -0.04] 0.00, aj -0.76) -0.52| -0.651 -0.70 -0.70 -0.66 -0.68 o ^ CO TJ "5 2 >> o -1.40| -1.361 -1.40| G. bull | -0.35| -0.36J -0.211 o.ool 0.011 -0.59| -0.711 -0.621 -0.76| . -0.80| -0.441 -1.09| -0.81| -0.30| -0.681 -0.661 -0.531 -0.67] -0.46] -0.57] -0.15 0.00 -0.16 -0.02 0.19 | 0.32 | 0.04 G. bull o d 46.11 58.0| 00 cri 15.5| LO d CM Tf CM d 504.0| Tf d 123.6| CO CM q d o o 538.6] q d 229.9| 1097.4| 41.7| 11.9| q d CM d 646.5| 66.2| 11.6| CO d CM Tf d CM d CTl d 44.7| o d o d q d oq CO LO CD LO LO 86.7] 10.5] 19.3] (#/g) Planktic N. pach | o d TT q od OJ CM CM Tf o d d 00 LO o d 276.8| d 42.7| cr> CM CM d d 196.8| o d 94.8| 548.7| 26.1 LO 00 o d d 185.4] 22.9] oq CM o d OJ CM LO d d o d CM 11.9 o d d o d o od CO 00 oj 50.6 CO CD CD 00 ndance Cibicid o d CM d OJ d CD d OJ d o d o d q d q d o d d OJ d O d O'O o d o d o d O d CO d d o d o d o d o d CD d o d O d O'O o d o d o d LO d OJ CO d O'O | d o d q CO OJ Tf f-d 00 d LO d Abu Benthic Uviger | d 00 d q CM q d d o d o d o d d o d o d o d q d q d q d q d o d o d o d o d o d q d q d o d o d o d o d o d o d o d q d q d o d o d o d o d d Tf d Tf oq d Gyroid 1 q d Tf d CO d CD d CM d d q d q d q d 00 d o d o d o d o d o d CM d O'O | o d o d d q d o d o d o d o d o d O'O | o d CM d o d q d o d o d d o d d q d q d o d d CO d CO d Tf d IRD Tf LO CM LO CM CO CO ro CO LO LO LO to LO CM Tf LO Tf LO CM LO oj - Tf LO Tf LO Tf LO Tf - - CO Tf CO LO CM OJ CO Tf Tf Age* Cal. (Ka) 402.811 404.031 406.161 411.971 417.771 421.261 423.221 425.46] 427.701 429.941 432.401 434.421 436.661 438.901 442.701 444.941 447.18] 451.661 453.901 457.701 460.161 462.401 464.64] 466.88] 474.27] 476.51] 478.75] 480.99] 495.31] 497.55 499.57] 501.81 510.761 513.00] 515.02 519.05 525.10 527.11 529.13 531.15 533.16 535.58 538.00 539.94 Age* 14C (Ka) Compos. Depth (mbsf) 27.811 27.91 28.01 28.11 28.211 28.271 28.311 28.41 1 28.511 28.61 28.721 28.811 28.911 29.011 29.211 29.31 1 29.41 1 29.611 29.731 29.901 30.011 30.111 30.21 j 30.311 30.67 30.771 30.871 30.971 31.611 31.71| 31.80] 31.91] 32.31 32.41 32.51 32.71 33.01 33.11 33.21 33.31 33.41 33.57 33.71 33.81 Sample Identifier B 4H3-32 1 B 4H3-42 B 4H3-52 B 4H3-62 B 4H3-72 B 4H3-78 B 4H3-82 B 4H3-92 B 4H3-102 I B 4H3-112 B 4H3-123 B 4H3-132 B 4H3-142 B 4H4-2 1 B 4H4-22 B 4H4-32 B 4H4-42 B 4H4-62 1 B 4H4-74 B 4H4-91 B 4H4-102 B 4H4-112 B 4H4-122 B 4H4-132 B4H5-18 B 4H5-28 B 4H5-38 B 4H5-48 B 4H5-112 B 4H5-122 B4H5-131 B 4H5-142 IB4H6-32 1 B 4H6-42 B 4H6-52 B 4H6-72 B 4H6-102 |B 4H6-112 B 4H6-122 B 4H6-132 B 4H6-142 1B 4H7-8 B 4H7-22 |B 4H7-32 00 CJs Local (pl.-ben.) -1.74 -1.49 -0.51 -1.23 -0.87 -1.73 -1.75 -1.68 CVJ t r t -1.10 -1.38 -1.57 -1.72 -1.72 -1.31 -1.37 -1.44 -1.66 -1.40 -1.41 -1.32 -1.15 -1.38 -2.05 -0.98 -1.28 -1.17 | Composite Benthic** 3.81 3.98 3.51 3.37 3.69 3.73 3.82 4.03 4.84 . 4.83 4.94 4.91 4.82 4.67 4.63 4.89 4.90 4.63 4.78 4.55 4.38 Cibicid +0.64 3.98 | 3.51 | 3.82 | 4.88 | 4.67 4.63 4.90 | 4.63 4.55 4.38 cn 3.37 | 3.69 d18C Gyroid | 3.81 1 3.73 | 4.03 1 4.81 4.83 4.94 4.91 4.82 4.89 4.78 G. bull | 2.18 | 2.41 | 3.37 1 2.75 | 2.89 I 2.02 1 1.98 | 2.05 | 3.09 | 3.74 | 3.45 1 3.37 | 3.19 3.10 ] 3.35 3.26 3.45 3.24 3.23 3.30 3.23 3.34 3.09 2.39 3.40 3.08 CO Cibicid -0.77] -2.06| -0.73| -0.91| -0.69| -0.70| -0.64| -0.47] -0.87] -0.71 <%o) Uviger | -0.86] -1.601 o CO ^ 5 2 -1.62| -1.56| -1.411 -1.47 -1.40 -1.30| -1.52 -1.62 -1.65 -1.47 G. bull | -0.231 -0.67| -0.851 -0.361 -0.33| -0.581 -0.441 -0.37| -0.38| -0.551 -0.491 -0.481 -0.39| -0.48| -0.541 -0.69 -0.43 -0.73 -0.57 -0.63 -0.74 -0.83 00 -0.83 -0.72 -0.89 -0.82 G. bull 26.81 cn c\i CO (NJ tr d 31.1 CVJ d o d o d o d o d 55.3|| 667.4| 90.2|| CD cvj o d o d o d o d o d d o d o d o d q LO 15.7| 911.4| 2995.3| 1480.0| 449.2| 163.5| 599.0| 193.0| 12.8| 107.5| 61.4| 69.8| CO CVJ rv! 29.8| CD cvj q 10.2] CVJ d 21.5] (#/g) Planktic N. pach | 25.5] - CO d q d 12.5| o d o d o d o d o d 42.0| 250.3| 28.11 CO o d o d o d q d q d q d o d d o d CD d 20.3| 1548.4] 3964.9| 3191.2| 604.9| 1029.11 4069.9 694.11 94,1J 869.8| 165.3 259.3 126.2 CD CD 122.8 lv. t r 15.2 23.4 q 23.2 ndance Cibicid d d o d rvi d q 00 d d o d o d o d o d o d o d O d o d o d o d d CO o d o d o d o d CVJ d t f d LO d d LO d CO d CD d q CVJ d LO d CD d d d LO d d d d CVJ d d o d d Abu Benthic Uviger | o d o d o d O d q d CO d CD d o d CVJ d d o d q d q d o d o d o d o d o d o d o d q d o d o d d d o d o d d q d q d o d d o d CVJ d o d o d o d o d o d o d o d o d q d q d Gyroid 1 o d o d q d ID d tT d d o d o d o d 0.0 LO d o d CVJ d o d o d o d o d o d o d o d o d d CVJ d d LO d LO d lv. d |v-d d CO d oo d CVJ d CVJ d d CVJ d CO d o d d O'O | o d o d o d o d IRD •- CO LO LO CO LO cvj LO rvj ro LO CVJ CVJ LO LO LO LO LO LO cvj ro LO CVJ LO tr CO ro LO r d CO LO Cvj CO LO r d CO LO CVJ LO CVJ LO CVJ LO LO LO Age* Cal. (Ka) 544.991 546.941 550.821 556.651 560.531 568.741 570.941 573.141 579.741 584.141 586.341 588.541 590.741 592.941 595.141 597.341 599.541 601.741 603.94 612.741 614.941 616.261 618.461 620.581 622.50! 624.421 626.341 628.261 630.181 632.101 634.02 635.941 637.86 639.78 641.70 643.62 645.78 649.80 651.58 654.26 658.72 660.50 662.51 664.74 Age* 14C (Ka) Compos. Depth (mbsf) 34.07 34.17 34.37 34.67 34.87 35.27 35.37 35.47 35.77 35.97 36.07 36.17 36.27 36.37 36.47 36.57 36.67 36.77 36.87 37.27 37.37 37.47 37.57 37.67 37.77 37.87 37.97 38.07 38.17 38.27 38.37; 38.47 38.57 38.67 38.77 38.87 38.97 39.17 39.25 39.37 39.57 39.65 39.77 39.87 Sample Identifier B 5H1-2 B5H1-12 B 5H1-32 B 5H1-62 B 5H1-82 B 5H1-122 B 5H1-132 B 5H1-142 B 5H2-22 B 5H2-42 B 5H2-52 B 5H2-62 B 5H2-72 B 5H2-82 B 5H2-92 B 5H2-102 B 5H2-112 B 5H2-122 B 5H2-132 B 5H3-22 B 5H3-32 B 5H3-42 B 5H3-52 B 5H3-62 B 5H3-72 B 5H3-82 B 5H3-92 B 5H3-102 B 5H3-112 B 5H3-122 B 5H3-132 B 5H3-142 B 5H4-2 B 5H4-12 B 5H4-22 B 5H4-32 B 5H4-42 B 5H4-62 B 5H4-70 B 5H4-82 B 5H4-102 B 5H4-110 B 5H4-122 |B 5H4-132 CjN I ffi 'O to •Q d cj + o *3 ^  5 n .S>l O) TO < O CD CJ < — o o c v c p E CD O 00 d o q t r co 8 CO d CO d o tf d o i d o d t r d o d o d o d q i d ' t r d o d o d o d q d CD d o d q d o d o d o d o d CD CO I N - -o CM CD 0) .— * J II CD C CD CO r w ra to CD ro T3 c a) ro II 13 -St1 O -s; ° ro o CJ rt CD J> >- ri t r r j o o Tj C o Tj CJ CO OS o SO t—I 'O •o 1) co rt X) .2 <J c .3 -t-H u .cu CO -2 § a) yj 00 ^ ii <• ° s s P " ° 3 ro °° < 8 "3 „ 11 o < CO c co tu _ E , o CD O CJ-CD § -a *-<u CO ^ i3 o <« E D O CO _ o 3 ro oo CD O JO « ~ < CD <0 i l a-l < a C "o 00 Si § • 0 0 E S-2 -a T3 N O) cn _J CjO 00 T „; cn CO _ CO II CM ? co oo ro " " ro 10 co S cr> " ro m- z» 9 co 1 -11 P £ ro -8 <5 8 o- -S -2 _ ro £ a. > CO c CO i E ™ E S o ^ E LO E cu - ° "8 © © s 5 CO o 

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