The Modern Organic Carbon Cycle in Hudson Bay, an Arctic Coastal Sea Undergoing Change

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The Modern Organic Carbon Cycle in Hudson Bay, An Arctic Coastal Sea Undergoing Change

Zou Zou Anna Kuzyk

A Thesis submitted to the Faculty of Graduate Studies of

The University of Manitoba

in partial fulfilment of the requirements of the degree of

Doctor of Philosophy

Deparlment of Environment and Geography

University of Manitoba

Winnipeg, Manitoba, Canada

FACULTY OF GRADUATE STUDIES

The Modern Organic Carbon Cycle in Hudson Bay, An Arctic Coastal Sea Undergoing Change

Zou Zou Anna Kuzyk

A Thesis/PI'acticum submitted to the Faculty of Graduate Stutlies of The University of

Manitoba in partial fulfillment of the requirement of the degree of

Doctor of Philosophy

ZouZou Anna KuzykO2009

Permission has been granted to theUniversityof Manitoba Libraries to lencl a copyof this thesis/practicum, to Library antl Archives Canada (LAC) to lend a copy of this thesis/practicum, and to LAC's agent (UMlÆroQuest) to microfilm, sell copies and to publish an abstract of this thesis/practicum.

This reprocluction or copy of this thesis has been macle available by authority of the copyright ovvner solely for the purpose of private study and researclr, ancl may only be reproducetl and õpied as permitted by copyright lalvs or rvith express rvritfen authorization from the copyright owner. Abstract

Arctic coastal seas are important sites of organic carbon cycling. Projected changes in Arctic temperatures, dver inflows and sea ice conditions will affect this cycling, with consequences for local ecosystems and global carbon cycles. To predict and measure the effects of change requires in-depth understanding of organic matter (OM) sources and processes controlling production, transport and burial. In this thesis, various geochemical tools were applied to study the modern organic carbon cycle in Hudson Bay, alarge but poorly-known Arctic inland sea, which is undergoing change more rapidly than other Arctic areas. Organic compositional data for sediments and suspended particulate samples from marine and river waters, together with a biogeochemical box model for nitrate, revealed that new marine primary production is concentrated in inshore surface watets, where there is increased upwelling of deep, nutrient-rich waters. This is probably supported in part by nutrient entrainment related to the large volume of river inflow to the Bay, which circulates through the inshore region. River inflow also provides the largest source of allochthonous (terrigenous) OM. Bulk (C/Ìrtr, ðr3C, ðr5N¡ and specific organic biomarkers (lignin) showed that heterogeneous terrestrial materials undergo hydrodynamic sorting in the coastal zone, resulting in the coarse fraction being retained near river mouths while a fine fraction undergoes transport by marine currents. Seasonal sea ice cover interacts with winter/spring river inflow to influence OM production and transpott indirectly. Sediment and particulate organic carbon budgets showed that resuspension and lateral transport of fine-grained coastal sediments is also an important process, supplying most of the sediment for contemporaly burial of OM and as much terrigenous OM as river inflow and subaerial coastal erosion combined. Resuspension also supplies offshore Hudson Bay with old (glacigenic) marine carbon, supporting slightly enhanced burial of marine OM in the Bay's sediments, compared to other Arctic shelves. The importance of resuspension likely reflects the exceptional postglacial isostatic rebound (relative sea-level fall) ongoing in Hudson Bay. Hypothesized transitional sedimentary and OC regimes in Hudson Bay pose challenges for interpreting responses to climate change and using Hudson Bay as a sentinel for change in Arctic coastal seas. Acknowledgements

First, I would like to thank my supervisors, Dr. Robie Macdonald and Dr. Gary

Stern, and ArcticNet, for providing me with the incredible opportunity to work in Hudson

Bay, and allowing me the freedom to take the research in my chosen direction. I also thank my supervisors for their mentorship and laying before me the challenge of pursuing good science and writing good papers; and for their patience and encouragement. I am grateful to Dr. Macdonald for his generosity with his knowledge of organic carbon cycling and the Arctic Ocean, for his efforts to train me to be a better scientist, and for debating with me about the game of science. I also benefited greatly from the mentorship of Dr. Miguel Goñi and Dr. Christine Michel, and from good science discussions with my committee members (Dr. Michel, Dr. Tim Papakyriakou, Dr. David Lobb) and Mats

Granskog, Greg McCullough and Alex Hare. I thank Monica Pazerniuk, teamZAM,

Mary o'Brien, cJ Mundy, Johannie Martin, the officers and crew of the ccGS

Amundsen, Churchiil community members, and fellow ArcticNet scientists for assistance and support with field work. E. slavicek, M. soon, Y. Alleau, and B. LeBlanc are thanked for laboratory support and P. Kimber for assistance with figures. Financial support from the Natural Sciences and Engineering Research Council of Canada

O{SERC), the Northern Scientific Training Program (Indian and Northern Affairs), the

Province of Manitoba and the University of Manitoba is gratefully acknowledged.

Finally, I would like to thank my grad room buddies, other good friends in Winnipeg and beyond, my family in Ontario, and especially Jason Stow.

111 Table of Contents

Abstract...... i Acknowledgements...... iii Table of Contents ...... iv List of Figures...... vi List of Tab1es...... x List of Copyrighted Material for which Permission was Obtained...... xi Thesis Format and Manuscript Claims...... xi Chapter 1 : General Introduction...... 1

Motivation and Thesis Objectives ...... 1 8ackground...... 6 Outline of Thesis Chapters .... 16 References ...... 18 Chapter 2 : Sea lce, Hydrological, and Biological Processes in the Churchill River Estuary Region, Hudson 8ay...... 31 Abstract...... 31 Introduction...... 32 Materials and methods...... 35 Results...... 43 Discussion ...... 62 Conclusions...... 71 References ...... 73 Chapter 3 : Sources, Pathways and Sinks of Particulate Organic Matter in Hudson Bay: Evidence from Lignin Distributions...... 80 Abstract...... 80 Introduction...... 81 Methods ...... 85 Results...... 91 Discussion ...... 104 Summary and implications...... 118 References ...... 121 Chapter 4 : Toward a Sediment and Organic Carbon Budget for Hudson Bay...... 131 Abstract...... 131 Introduction...... 132 Overview of the Hudson Bay System ....134 Methods ...... 138 Results and Discussion ...... 150 References ...... 182

1V Chapter 5 : Elemental and Stable Isotopic Constraints on River Influence and Patterns of Nitrogen Cycling and Biological Productivity in Hudson Bay ...... 196 Abstract...... 196 Introduction...... I97 Study area ...... 201 Methods ...... 203 Results and Discussion...... 209 References ...... 234 Chapter 6: General Discussion and Conclusions...... 246 Marine primary production and its controls ...... 251 Terrigenous OM: sources, composition and distribution ...... 257 Role of resuspension...... 260 Sediment accumulation and OM burial ...... 262 Implications for Hudson Bay's responses to climate change ....265 References ...... 269 List of Figures

Figure 1-1. Location of Hudson Bay and the Arctic Ocean shelf seas...... 4 Figure 2-1. Layout of sampling sites (circles) in the Churchill study area and general ice conditions (March-May 2005) .....36 Figure 2-2.Tidal elevations in Churchill Harbour, air temperatures in Churchill, and Churchill River discharge (solid line) and temperature (dashed line) measured at the weir over the five main sampling periods (winter, pre-melt, early melt, peak flow and break up). Black symbols show specific sampling dates...... 44 Figure 2-3. Evolution of the (A) landfast ice edge off the Churchill River, showing where rubble ice was lost during the peak flow (white arrows), and the fast ice then ablated in the break up (black arrow) and (B) surface reflectance, largely a function of open water, in the Churchill River estuary. Images are modified from A) Quickbird RGB composites, May 16 and27, and B) MODIS "MOD09GQK" BAND 2, Aprll?7,May 16, May 29...... 47 Figure 2-4. Salinity profiles for Button Bay and El under the ice in the pre-melt, early melt and peak flow periods (spring 2005) and in open water in the fall of 2005..... 50 Figure 2-5. ôl80 - salinity relationships for water samples from the Churchill estuary region (a) and ôr8O profiles for ice cores from the estuary area (April 8 at E1, solid square, May 19 at T3, open squffe) and Button Bay (April 6, solid triangle and May 5, open triangle) (b) ...... 52 Figure 2-6. Seasonal trends in surface water salinity in Button Bay and the estuary region (El and T3) derived from ice core records of ðr80...... 52 Figure 2-7. Nitrate - salinity and silicic acid - salinity in Churchill Estuary surface water samples. Note the linear relationships in the winter & pre-melt periods (open circles) and departures from linearity in the early melt & peak flow periods (closed squares)...... 60 Figure 2-8. Seasonal trends in surface water chlorophyll a concentrations and ratios of chlorophyll a to phaeopigment, and the ðl5N and õr3C values of suspended particulate material in Button Bay, the estuary region (81, T1-T5), and the Churchill River...... 61 Figure 2-9. Conceptual model of the winter-spring transition in the frozen Churchill estuary area. Transect extends from the weir, at the upper end of the estuary, across the landfast ice zone, and out into the flaw lead. Contours show isohalines (dashed portionsspeculative)...... 64 Figure 3-1. Maps showing the Hudson Bay watershed, major rivers and the position of tlre tree line (after Rouse e|a1.,1997; Stewart and Lockhart, 2006) and the location of the 11 sediment cores examined in this study (filled circles labelled with core number)...... 85 2rOPb Figure 3-2. Profiles of the natural log of excess activity in the sediment cores. Points represent measured values while lines represent the model results. Supported

v1 2l0Pb levels, surface mixed layer (SML) depths and other parameters are listed in Table 3-1...... 90 Figure 3-3. Total lignin (ÀS) (yields (mg/i00 mg OC) in sediment core sections (A) and relationships with distance from shore (B) and latitude (C) (mean +SD)...... 94 Figure 3-4. Profiles of lignin compositional ratios in the sediment cores, including ratios of syringyl, ciruramyl, and 3,5-Bd to vanillyl phenols (s/v, crv, and 3,5- Bd/V, respectively) (A) and acid to aldehyde ratios for vanillyl ([Ad/Al]v) and syringyl ([Ad/Al]s) phenols (B). Dates of sediment deposition estimated from 21oPb dataare shown on the y-axis, however the intrinsic time resolutions of the cores varied from 6-57 years (shown as vertical lines in each plot)...... 100 Figure 3-5. Ratios (mean +SD) of syringyl and cinnamyl phenols to vanillyl phenols (S/V and C/V, respectively) vs. latitude (A) and distance from shore (B) and ratios of acidic to aldehydic syringyl phenols ([Ad/Al]s) ploned against the same x- variables (C and D)...... 102 Figure 3-6. Plot of C/V vs. S/V ratios. Core numbers identify the surface sample of each core. Filled boxes show ranges for major vascular plant types (Goni et al., 1998; Hedges and Mann, 1979; Hu et a1.,1999). Open box shows range for suspended particulate organic matter (POM) collected from 13 northern rivers, including the Mackenzie River (Goni et a1.,2000) and 12 Russian rivers (Lobbes et a1.,2000). .... 103 Figure 3-7. Plot of C/V vs, 3,5-Bd/V ratios. Core numbers identify the surface sample of each core. Dashed boxes show approximate ranges for plant debris and peat (Goni et al., 2000; Louchouarn, 7997), surface soils (Louchouarn, 1997; Prahl et al., 7994), and subsurface soils (Houel et al., 2006; Ugolini et al., 1981)...... 103 Figure 3-8. Plot of [Ad/Ad]v vs. syringyl [Ad/Al]s ratios. Filled box shows typical range for fresh plant tissues (Hedges and Mann, 1979; Hedges et al., 1982; Goni et al., 1993). Sd yields were below detection in samples shown on the x-axis...... 104 2l0Pb-derived Figure 3-9. sediment accumulation rates (black bars: g cm-2 a-l) and modem lignin accumulation rates (white bars: mg u-'), shown as an overlay on a preliminary sediment classification map for the Bay,^-' based on interpretation of Huntec D.T.S. and 3.5kHz seismic data (Josenhans et al., 1988)...... :...... 105

Figure 4-1. Map of Hudson Bay showing major rivers sampled in 2005 and 13 sediment box-core sites (filled circles labelled with core number)...... 135

Figure 4-2.Hypsometry of Hudson Bay in 50-m depth intervals (bars) and as a curve showing the area (kt"') and proportion(%) of the seafloor above the depth indicated 21oPb-derived on the y axis. Horizontal dotted lines and circles show sedimentation rates by depth for the cores from this study (filled circles) and four previously published values (open circles) from d'Anglejan and Biksham (1988) and Lockhart et al. (1998)...... 136 tl0Pb Figure 4-3. A) Profiles of the natural log of excess activity in the sediment cores. Points represent measured values while lines represent the model results. Arrow indicates the bottom of the surface mixed layer (SML). Supported 2l0Pb levels, SML

vlr depths, and other parameters are listed in Table 4-2. B) Porosity profiles in the sediment cores. The shaded zones indicate textural changes...... 146 Figure 4-4. Downcore profiles of organic carbon content (%OC) and ôr3C (%o) in cores a) 5, b) 6, and c) 8. The %OC profiles in these cores are interpreted as reflecting OC oxidation in the surface sediments...... 149 2l0Pb-derived Figure 4-5. sediment accumulation rates (black bars: g cm-2 a-l¡ and average proportions of marine and terrestrial carbon in the sediments derived from õl3C data, shown as an overlay on a preliminary sediment classification map for the Bay, based on interpretation of Huntec D.T.S. and 3.5kHz seismic data (Josenhans et al., 1988; Geological Survey of Canada Open File Report #2215). Open bars show sedimentation rates for cores (HUD-4, FOGO-4) collected in 1992193 (Lockhart et al., 1998) and previously published rates for two cores collected in 1985 (5583, S40; d'Anglejan and Biksham, 1988)...... 155 Figure 4-6. Relationship between ô13C values (%o) and sedimentary lignin yields (z\8), defined as the sum of vanillyl, syringyl and cinnamyl phenols (Hedges and Mann, 1979), in units of mg/100 mg OC. ..159 1) Figure 4-7. Budgets (106 t a for sediment (A), marine organic carbon (OC) (B) and terrigenous OC (C). Straight arrows represent particulate organic carbon (POC) and wavy arrows dissolved organic carbon (DOC). Arrows into the water column represent inputs from autochthonous primary production (PP), river input, coastal erosion, resuspension and lateral transport of coastal deposits, and atmospheric inputs. Arrows out of the water into the sediments represent sediment burial and arrows out of the water column to the right represent export by ice and advection or oxidation of DOC. Circular arrows represent internal losses (oxidation or leaching). The resuspension term in the sediment budget was derived by difference to balance the sediment budget and the unbalanced terms in the OC budgets assigned to water column leaching/oxidation...... I79 Figure 5-1. Map of study area and location of samples, including dissolved nutrients (X), particulate organic matter (POM) from surface waters (open squares), POM from subsurface waters (open squares labelled I-I), and sediment boxcores (closed circles)...... 204 Figule 5-2.Yertical profiles of subsurface POM OC/TN ratios (a), ôr3C values (b), and ô"N values (c). The location of the five profiles (labelled I-'t) is shown in Figure 5- 1. Depth of the water column is indicated by gtey shaded area. Properties of nearby sedimentary OM are indicated with filled squares...... 212 Figure 5-3. Surface sediment CAtr ratios (a¡, 813C values (b) and ôlsN value s (c)...... 217 Figure 5-4. Relationships between sediment ôl3C values and N/C ratios (a) and ôlsN values (b). Lines show best fit line in a) and example two end-member mixing lines in b), the latter non-linear because they incorporate the difference in C/ÌlI ratios of marine and terrigenous OM (here 6.6 vs. 21). The two end-member model still provides a poor fit to the ôl3C - ðlsN relationship. Surface samples of the cores show core label. Dashed outline boxes show ranges for river SPM and marine POM. ...218

vllr Figure 5-5. Relationship between estimated ðlsN values in marine-derived sedimentary OM (ôr5Nn'u,) and distance from shore ...... 225 Figure 5-6. Relationship between estimated ôr5N values in marine-derived surface water POM (õrsNn.,u) and surface water salinity...... 225 Figure 5-7. Nitrate (+nitrite) concentrations (¡rM) plotted against salinity for samples collected in September-October 2005. Sample locations are shown in Figure 5-1. (Plot produced using Ocean Data View, R. Schlitzer, http://odv.awi.de.) ...... 227 Figure 5-8. Biogeochemical model for Hudson Bay for freshwater, salt and nitrate (+nitrite). The Bay is represented as a two-layer stratified water column (0-50 m surface layer) and separate inshore and offshore regions to produce four boxes (i.e., four sub-basins). Net volume flows are represented by one-way anows and mixing exchanges by two-way anows...... 230 Figure 6-1. Schematic illustrating major sources and processes in the Hudson Bay OC cyc1e...... 248 uî;o ''*:1 1 l:::l: i:i ::::::.:ïi:t:: ::i:::::::::: :::ï:: ï:T::t: ":i:::

IX List of Tables

Table 2-1. Summary of sampling ...... 37 Table 2-2. Oxygen isotope (õ'tO) values (%o) for western Hudson Bay (offshore), Button Bay, the estuary surface waters, and the Churchill River, in fall 2005 and over the winter-spring sampling periods ...... 51 Table 2-3. Equivalent heights (m) of river water (RW) and sea ice melt (SIM) in a 10 m water column in the Estuary (E1) and Button Bay and the top 4 m (the plume) at El, and estimated estuary flushing times based on Churchill River discharge...... 57

Tabie 2-4. Estimated heat budget elements contributing to sea ice melt in the estuary region ...... 57 Table 2-5. Concentrations (mean (range) in uM) of dissolved nutrients in Button Bay, the estuary surface waters, and the Churchill River, over the winter-spring sampling periods...... 59 Table 3-1. Core properties and sedimentation parameters ...... 93 Table 3-2. Yields (mgl100 mg OC) of lignin phenols, total lignin (z\8) and 3,5- dihydroxybenzoic acid (3,5-Bd)...... 95 Table 3-3. Lignin accumulation rates (-g *-' a-'; for three time periods (1875-1925, 1925-197 5, 197 5-2005) calculated from lignin yields (mg/l 0 g sediment) in sediment intervals corresponding to these periods multiplied by the sediment accumulation rate within each core (Table 3-l) ...... 97 Table 4-1. Suspended particulate matter properties and DOC concentrations for Hudson Bay rivers ....140 Table 4-2. Sediment core properties and sedimentation parameters ...... 144 Table 4-3. Organic carbon burial rates, surface fluxes and oxidation rates ...... 149 Table 4-4. Sediment and organic carbon budget (x106 t a-r) ...... 158 Table 4-5. Suspended sediment concentrations measured in Hudson Bay drainage areas and estimated regional sediment inputs to the Bay...... 168 Table 5-i. Characteristics of suspended particulate matter (SPM) and dissolved nutrient concentrations in Hudson Bay rivers ...... 205 Table 5-2. Properties of particulate organic matter (POM) from Hudson Bay surface waters ...... 207 Table 5-3. Properties of cores and surface sediment sections (0-l crn) ....216 Table 5-4. Choices for 'best estimates' (and ranges) for representative salinity and nitrate in budget ...... 229 List of Copyrighted Material for which Permission was Obtained

Reproduced in this thesis is a hgure adapted from "Digital Maps of Surficial Geology of Hudson Bay", Zevenhuizen, J; Josenhans, H. Geological Survey of Canada, Open File 2215, 1990, with the permission of Natural Resources Canada 2009, couftesy of the Geological Survey of Canada (Open Flle 2215).

Thesis Format and Manuscript Claims

This thesis includes four manuscripts (Chapters 2-6), each containing an Abstract, Introduction, Methods, Results and Discussion section. Chapter I provides a general introduction and Chapter 6 a general discussion and conclusion. Chapter 2: Kuzyk,Z.A.,Macdonald, R.W., Granskog, M.4., Scharien, R.K., Galley, R.J., Michel, C., Barber, D. and Stern, G., 2008. Sea ice, hydrological, and biological processes in the Churchill River estuary region, Hudson Bay. Estuarine, Coastal and Shelf Science 77, 369-384. Z.Kuzyk collected and analyzed physical, chemical and biological data from the Churchill River estuary region and prepared the manuscript with the participation of co- authors. R. Scharien analyzed the MODIS image and R. Galley provided output from a one-dimensional thermodynamic model for sea ice growth. Chapter 3: Kuzyk, 2.A., Goñi, M.4., Stern, G.A. and Macdonald, R.W., 2008. Sources, pathways and sinks of particulate organic nxatter in Hudson Bay: evidencefront lignin

distr ibut ions. Marine Chemistry 1 12, 27 5-229, doi : 1 0. 1 0 1 6/j.marchem.200 8. 08. 00 I . Z.Kuzyk collected the sediments, did the lignin analysis under the supervision of M. Goñi, and wrote the manuscript with the input of co-authors. Chapter 4: Kuzyk,Z.A.,Macdonald, R.W., Johannessen, S.C., Gobeil, C. and Stern, G.4., 2009 . Towards a sediment and organic carbon budget.for Hudson Bay. Marine Geology 264,190-208. Z. Kuzyk collected the sediments, co-ordinated lab analyses, and prepared the manuscript with the input of co-authors. S. Johannessen and R. Macdonald performed the numerical l37Cs modelling of and contaminant Pb distribution. Chapter 5: Kuzyk, Z.A.,Macd.onald, R.W., Tremblay, J.-É. and Stern, G., Elemental and stable isotopic constraints on river inJluence and patterns of nitrogen cycling and biological productivity in Hudson Bay. Continental Shelf Research In review; submitted April2009. Z.Kuzyk collected and co-ordinated analysis of the sediment and particulate samples and J.-E. Tremblay's team collected and analyzed the nutrient samples. Z.Ktnykprepared the manuscript with the input of co-authors.

xi Chapter 1: General Introduction

MOTIVATION AND THESIS OBJECTIVES

The ocean provides one of the major natural sinks for actively cycling carbon on a global scale and consequently is a key component of the Earth's climate system (Denman et al., 2007). Although dissolved inorganic carbon provides most of the oceanic sink, the organic carbon stored in living substances and the physical and biogeochemical processes that control its production and preservation play important roles, providing a 'motor' that keeps cycling in motion ((Dietrich, 1963) as cited in (McCarthy, 2000)). Collectively, the transformation of carbon between organic and inorganic and dissolved and particulate forms, the physical redistribution of organic carbon within the water column through transpotl and mixing, and the interactions of these processes serve to modulate atmospheric carbon dioxide concentrations over geological timescales (Eglinton and

Repeta, 2004; Siegenthaler and Sarmiento, 1993).

The coastal ocean, including continental shelves, estuaries, and inland seas (about

10% of the global ocean area), is a particularly active site for carbon cycling because of larger autochthonous production of organic matter (primary production) relative to the open ocean and additional inputs of allochthonous organic matter (terrigenous material and freshwater aquatic production) from the surrounding watershed (Liu et al., 2000;

Smith and Hollibaugh, 1993). An estimated 80%-95% of the organic matter preserved in the oceans presently undergoes burial in coastal ocean sediments and is thereby removed from active cycling (Chen et al., 2003; Hedges and Keil, 1995). Organic matter burial also plays an important role in the cycling of other associated elements, including contaminants (e.g., Jonsson et al., 2003). Nonetheless, the intensity of physical and biogeochemical processes in coastal seas provides potential for rapid exporl, modification or turnover of organic matter (Liu et al., 2000), and consequently changes in the rates of these processes, their interactions, or where they occur (cf. Rippeth et al., 2008), can cause OC burial to change. Thus, the coastal ocean has been more profoundly impacted by past human activities (e.g., damming, eutrophication, over-fishing) and is now expected to be the most sensitive part of the marine environment to global climate change

(Liu et a1.,2000; Smith and Hollibaugh, 1993).

Arctic shelf seas share many features with other coastal seas including large river inputs and enhanced primary production compared to offshore areas, but provide a uniquely seasonal and ice-influenced environment, with significant consequences for many aspects of their organic carbon cycles (Carmack et al., 2006; Chen et a1.,2003;

Stein and Macdonald,2004c). Climate models consistently show the Arctic to be one of the most sensitive regions to climate change, with larger and more rapid temperature changes and greater responses to these changes relative to more southerly areas (ACIA,

2005; Johannessen et al., 2004: Lawrence and Slater, 2005). Among the responses to climate change that have already taken place, many have the potential to substantially impact the OC cycles of Arctic shelves, including decreases in extent and thickness of sea ice (Johannessen et al., 2004), permafrost thawing (Lawrence and Slater, 2005; Payette et al,2004), coastal erosion (IPCC,2001 St. Hilaire etal.,2007), and altered distribution and abundance of marine species (Grebmeier et al., 2006).

Hudson Bay, which is the world's largest inland sea (ca. 840 x 103 km2;, lies at the southern margin of the Arctic (Figure 1-1), in the centre of subarctic and Arctic North America (52'N to 82'N; Rouse et al., 1997). With Arctic Ocean waters comprising much of its inflow and a complete sea-ice cover developing every winter, Hudson Bay has much in common with the other Arctic seas. However, owing to its southern latitude and close contact with terrestrial systems, Hudson Bay may be subject to greater and perhaps faster climate-related changes than other Arctic seas (ACIA, 2005). Changes in river discharge (Dery et al., 2005; Dery and Wood, 2005; Lammers et al., 2001), permafrost in the surrounding watershed (Payette et al., 2004), and sea ice conditions (Johannessen et a1.,2004; Parkinson and Cavalieri,2002; V/u et a1.,2005) have already been documented, with further dramatic changes projected for the end of the century (Gagnon and Gough,

2005; Gough and Wolfe, 2001; Westmacott and Burn, 1997).

Recent syntheses of data describing the sources, composition, distribution, and burial of organic matter in marginal seas of the Arctic Ocean (Carmack et al., 2006;

Gebhardt et al., 2005; Macdonald et a1.,2004a; 1998; Stein and Fahl, 2004) have significantly advanced our understanding ofthe general processes controlling organic carbon cycling in Arctic coastal marine environments. From this basis, broad predictions have been advanced about how Arctic coastal systems will respond to climate change

(e.g., Carmack et al.,2006; Chen et a1.,2003). However, because of the variability of the

Arctic shelf seas, accurately predicting actual responses to change will require an understanding of the important processes at the regional scale (Liu et al., 2000). The scarcity of oceanographic data from Hudson Bay (e.g., Stewart and Lockhart, 2005), and limited understanding of the physical, chemical and biological processes that govern terrigenous inputs and autochthonous production of organic matter in this system mean that it represents a key regional gap. The limited knowledge about the Bay has hampered adequate assessment of the impacts of natural climatic variability and past anthropogenic effects related to dams and diversions (Dery et al., 2005; LeBlond et a1.,1996;

Prinsenberg, 1983; 1980) and now make global climate change impacts in Hudson Bay difficult to predict.

Figure 1-1. Location of Hudson Bay and the Arctic Ocean shelf seas.

The overall goal of my research is to address some of the critical knowledge gaps surrounding the rnodern cycling of organic carbon in Hudson Bay. Advancing our understanding of the processes that control organic carbon (OC) cycling in Hudson Bay, and similarities and differences between this system and better-studied Arctic coastal seas, represents an important first step towards ultimately predicting how Hudson Bay will respond to change. If Hudson Bay responds to climate change more quickly than other Arctic areas, there may be valuable insights to be gained about how climate change will affect the Arctic coastal marine environments, provided important differences between these systems are clearly understood.

The studies in the thesis address the sources, pathways and sinks of terrigenous and marine organic matter in Hudson Bay and the major controlling processes within the marine system including circulation, ice rafting, primary production, and sedimentation, and at the interface between rivers and the Bay.

The approach taken to study the OC cycle in Hudson Bay was to collect sediment cores and suspended particulate matter samples from the marine environment and rivers discharging to Hudson Bay, apply various geochemical tools (t'oPb dating, bulk organic proxies, specific organic carbon biomarkers) to characterize the organic content of samples and determine rates of key processes (e.g., sedimentation, oxidation), and then synthesize the new data with literature data to construct conceptual models (e.g., of estuarine structure and function) and box-model budgets (sediment, organic carbon, nitrate). Each Chapter in the thesis has been prepared as a separate manuscript, which includes its own introductory material and explanation of the specific geochemical tracers and models being applied. In order to keep repetition to a minimum, below I provide some background information concerning the nature and fate of organic matter in the oceans (sotuces, pathways and sinks), an overview of the geochemical approaches that were applied, and an introduction to the processes that impact carbon cycling in Arctic coastal seas in general and may specif,rcally impact OC cycling in Hudson Bay. BACKGROUND

Overview of the Organic Carbon Cycle - Sources, Pathways, Sinks

The major sources of organic matter in marine systems may be either marine

(primary production) or terrestrial (eroded soil and plant material), with freshwater

aquatic production generally contributing an insignificant component compared to the

other two sources (lttekkot, 1988; Lobbes et al., 2000; Meybeck, 1982; Onstad et al.,

2000). Marine phytoplankton produce organic matter in the euphotic zone of the ocean in

response to the controlling factors that modi$r nutrient supply and light. Terrigenous OM

is supplied to marine systems by both river input and coastal erosion, with supply by

aeolian processes relatively minor in coastal (shelf) seas (Stein, 1991b).

In global carbon budgets (Eglinton and Repeta,2004) and most local or regional

budgets (cf. Durrieu de Madron et al., 2000; Johannessen et al., 2003; Macdonald et al.,

i998; Thomas et al., 2005), marine primary production supplies a much larger quantity of

OM than the combined terrigenous sources. Nevertheless, the terrestrial and marine

components of the organic carbon cycle both need to be studied because their behaviour

and fate are affected by intrinsic (compositional) differences and by the influence of

different impofiant pÍocesses (Hedges and Keil, 1995). Terrigenous and marine sources

and the major active processes affecting each component may also respond differently to

global climate change (Chen et al., 2003).

Terrigenous organic matter generally includes a large fraction of biologically resistant (recalcitrant) compounds such as lignin and kerogen and OM already strongly degraded by soil microbes (Bergamaschi et al.,1997; Hedges and Oades, 1997). Marine

OM is generally much more labile and, for example, grazing by zooplankton frequently limits phytoplankton biomass (Wassman et al., 2004). Only about 20o/o of netprimary production settles out of ocean surface waters and fluxes decrease exponentially down the water column, with less than 70Yo of surface production generally remaining at depths of hundreds of metres and only about 1o/o in deep ocean areas (Betzer et al., 1984; Hedges and Keil, 1995; Suess, 1980; Wassman et a1.,2004). Despite its intrinsic recalcitrance, terrigenous OM can also be recycled in the marine environment. Modern terrigenous OM is recycled quite efficiently in some places, €.9., >650/0 on the Beaufort Sea shelf (Goni et al., 2005), and ancient terrigenous OM fractions, comprising kerogen or old soil organic matter also, albeit to a lesser extent (<25%o; Goni et al., 2005). Teruigenous OM supplied by rivers undergoes intensive reworking in estuaries and deltas (Eglinton and Repeta,

2004), with some estimates suggesting losses of as much as two-thirds of the global fluvial input of POM in these environments (Keil et al., 1997). The fractions of terrigenous OM lost to leaching (conversion to dissolved organic matter) vs. completely mineralized (to COz) vs. exported or buried are probably highly variable and the controlling mechanisms in most systems are not yet well known.

The small portion of marine OM and variable portion of terrigenous OM evading recycling in the water column and surface sediments is ultimately preserved and buried in seafloor sediments (Hedges and Keil, 1995). Although the detailed mechanisms of OM preservation in sediments are not yet clear, several principal factors contribute (cf.

Eglinton and Repeta,2004).In addition to the intrinsic properties (recalcitrance) of the compounds comprising the OM, preservation and burial is enhanced by high sedimentation rates, stabilization of OM by sorption onto sedimentary mineral surfaces, and sustained anoxic conditions in the sediments (Hedges and Keil, 1995; Henrichs, 1995). Coastal marine sediments show large differences in their organic matter content and composition, reflecting variations in the supply of marine and terrigenous OM, the environmental conditions under which the OM was produced, and the transport and transformati on processes controlling OM fate.

Application of Geochemical Data to Studies of Marine Organic Carbon Cycles

Sediments represent a repository for organic matter entering and produced within the ocean and as such, form a basis from which the organic carbon cycle may be studied over various space and time scales (Altabet and Francois,7994; Calvert et aL., 1992;

Hedges et aL,7982; Macko, 1989; Muzuka and Hillaire-Marcel, 1999; Schubert and

Stein, 1996; Stein, 7997a). Of principal importance in applying sediment core data to the study of organic carbon cycling is an understanding of the sedimentary context of the cores and in particular, estimation of sedimentation rate. Sedimentation rates provide a basis for estimating dates associated with various sediment core sequences and thereby obtaining a record of changes in the OM within the system and they also allow estimation of rates of other processes (e.g., organic carbon burial).

t'OPb The atmospherically-delivered natural radionuclide (half-life 22.26 years) has been widely used to estimate accumulation rates in recent sediments (e.g., Robbins,

7978), including sediments from the Arctic Ocean (Baskaran and Naidu, 1995; Huh et al.,

1997; Smith et al., 2003) and a few cores from Hudson Bay (d'Anglejan and Biksham,

1988; Lockhart et al., I 995; 1998). The possible influence of surface mixing processes

(e.g., bioturbation) on the depth distribution of tracer substances is an important consideration in marine environments (Andelson et al., 1988). The influence of mixing

2lOPb processes may be addressed through modelling approaches that incorp orate a surface mixed layer (SML) and a deeper layer where the sediments are relatively unmixed (Carpenter et al., i 985; Lavelle et al, 1986; Macdonald et al, 1992).

2rOPb Sedimentation rates are derived from the portions of the profiles below the SML.

2lOPb-derived Additional sedimentary tracers improve confidence in rates (Boer et al., t"Cs 2006; Smith, 2001). thalf-life 30.0 years), which has a well-known input history associated with fallout from atomic weapons testing (largely between 1952 and 1964), has been widely used as an independent time marker to verify accumulation parameters

2loPb derived from (Baskaran and Naidu, 1995; Fuller et al., 1999). Contaminants with known input histories (e.g., Pb) provide possible additional tracers.

Suspended and sedimentary OM in coastal areas is generally extremely heterogeneous, containing various particles resulting from a wide range of inputs and numerous transport and transformation processes. Characterizing OM through the application of geochemical proxies provides a means of querying the sediment record for information of relevance to organic carbon cycling (Hedges and Oades, 1997; Meyers,

1994; Stein and Macdonald, 2004b). Elemental (carbon and nitrogen) ratios and stable carbon and nitrogen isotope ratios (õr3C and ôr5N¡ are among the most widely-used

'bulk' proxies providing information about organic matter origin (Meyers, 1994; Stein,

1991b). Low carbon/nitrogen (Cn$ ratios and high (or enriched) ô'3C values in marine sediments both distinguish OM derived from marine sources (phytoplankton or ice algae) from contributions of terrigenous OM, which is characteristically nitrogen-poor and isotopically depleted (Fernandes and Sicre,2000; Hedges et al., 1988; Meyers, 1994;

Ruttenberg and Goni, 1997). Use of these proxies together helps confirm that differences are indeed related to source and not post-production processes (e.g., microbial

9 degradation, zooplankton scavenging). Elemental and isotopic compositions of various source materials (e.g., terrestrial plant matter, marine phytoplankton) are well characterized in a general sense for Arctic waters (Gaye et a1.,2001; Goni et a1.,2000;

Lobbes et a1.,2000; Naidu et a1.,2000; Schubert and Calvert,200l; Schubert and Stein,

1996) and additional site-specific samples may be analyzed to increase confidence in regional source signatures. Stable nitrogen isotope ratios (ôttN) of sedimentary OM have been used in some environments also as a tracer of marine vs. terrigenous OM sources

(enriched and depleted, respectively; Naidu et al., 2000) but in other environments, these values provide a proxy for surface water nutrient conditions at the time marine OM was being produced (Altabet, 2004; Altabet and Francois,7994; Francois et al., 1992).

Although geochemical bulk parameters such as CÆ..1ratios and ðl3C values are important proxies for identifuing general origins of organic matter, specific biomarkers often provide more precise information about the specific sources contributing to complex mixtures of OM, like tliose commonly found in coastal sediments (Meyers, 1994; Stein,

1991b; Stein and Macdonald,2004b). The detailed identification of tenigenous sources is clearly imporlant for linking variation to changes in specific watershed sources (e.g., vegetation types) or processes (overland runoff vs. river bank erosion). One of the best proxies for terrestrial (vascular plant) carbon input is lignin because it is a major component of terrestrial vascular plants but essentially absent from marine organisms

(Goni and Hedges, 1995; Hedges and Mann, 1979). Lignin components differentiate angiosperm and gyrnnospelm sources, woody vs. non-woody tissues, and fresh vs. degraded material. Nevertheless, biomarkers like lignin complement rather than replace bulk proxies because they generally only directly represent very small proportions of the

10 total organic carbon content of sedimentS, €.9., lignin commonly composes only a few per cent of the total organic carbon content of the sediments (Goni et al., 2000; Goni et al., 2005). Bulk proxies and specific biomarkers used in combination generally yield the most confident OM source apportionments and hence greater insight into the influence of transporl and transformation processes.

Features of Arctic Shelves and Implications for OC Cycling

Arctic seas are unique environments for OC cycling because of the presence of sea ice cover, which affects both primary production and transport of terrigenous OM, the exceptionally strong influence of freshwater, coming from both large river inflows and seasonal ice melt, and the seasonality of forcing (runoff, ice formation, light) and consequently biological production (Carmack et al., 2006; Macdonald eL al.,2004b).

Arctic Ocean waters, which supply most of the inflow to Hudson Bay as well as all the marginal seas around the Arctic Ocean (Figure i-1), are characterized by cold surface temperatures, low surface salinities, and strong summer stratification. The

'meditemanean' character of the Arctic Ocean and marginal seas also means that land- ocean connections are strong and estuarine, coastal, and even river and watershed processes relatively influential for marine biogeochemical cycling (cf. Macdonald et al.,

2004b). Sea ice modifies coastal area processes, for instance by imposing horizontal structure (landfast ice, rubble zones, flaw leads), which influence the paths of river plumes, and by producing negative estuarine circulation in winter (as brine is released from growing sea ice) in areas where river plumes and positive estuarine circulation dominates in summer (Macdonald and Carmack, 1991). Moving shoreward, Arctic coastlines are distinctive morphodynamic environments, generally undergoing relatively

11 rapid submergence and retreat, with consequences for sediment and terrigenous organic carbon input into the coastal area. The coastal evolution is a result of rising relative sea level (RSL), which affects most shorelines in the world, but in the Arctic, enhances erosion of ice-rich cliffs and shorelines, already prone to thermal erosion (Dallimore et a1.,7996; Rachold et al., 2000). Longer ice-free seasons, increased wave energy, and warmer air temperatures associated with climate change contribute to the rapid coastal evolution.

Although the Arctic shelf seas share many common features and seasonal processes, they differ strongly in terms of properties of the inflowing seawater, freshwater and sediment supply, circulation, rates of ice production and transport, sources of nutrients, and coastal properties, and consequently show large variation in primary production, the contribution of terrigenous sources, the exports and transformations of the organic matter in the water column, and burial of marine and terrigenous materials (Chen et al., 2003; Macdonald and Anderson,2008; Stein and Macdonald,2004a). Considerable progress has been made recently toward understanding the basic mechanisms controlling

OC cycling on the various shelf seas (cf. Carmack et al., 2006; Carmack and Wassman,

2006; Stein and Macdonald,2004a), thus providing a conceptual framework for identiffing the processes controlling OC cycling in Hudson Bay.

Nutrient regimes and consequently marine primary production vary widely among the Arctic shelves, with greatest annual production in the Barents and Chukchi Seas

(Figure 1-1), which receive relatively nutrient-rich inflows from the Atlantic and Pacific

Oceans, respectively (Carmack et al., 2006). Annual primary production is roughly 5- to

10-fold lower in parts of the Canadian Arctic Archipelago (e.g., Lancaster Sound, Michel

t2 et a1.,2006), one of the Arctic Ocean outflow areas, and in the interior shelf seas

(Beaufort, East Siberian, Laptev and Kara Seas), where the oceanographic conditions are strongly influenced by inflows from the major Arctic rivers (Macdonald et a1., 2004b).

The nutrient regimes of the interior shelves depend on a complex combination of factors including the direct inputs of nutrients from the rivers, the stratification of the water column that results from the river inflows combined with sea ice melt, wind-driven vertical mixing and upwelling, and seasonal replenishment of nutrients in surface waters due to winter vertical mixing, which is in turn a function of winds, rate of sea ice formation and winter river runoff distribution. Surface water phytoplankton, ice algae, and the subsurface chlorophyll maximum layer contribute to total primary production in varying proportions among the shelves depending on the nutrient regime as well as the duration of the sea ice cover (Carmack et a1., 2006; Sakshaug, 2004).In Hudson Bay, the nutrient regime and annual primary production by phytoplankton and ice algae will similarly be influenced by the nutrients provided by the mostly Arctic Ocean surface- water inflow (Ingram and Prinsenberg, 1998; Jones and Anderson,7994; Prinsenberg,

1986), the nutrient inputs of the rivers (cf. Hudon et al., 1996), the frequency of upwelling driven by winds and other processes, and the replenishment by winter mixing, which reaches to 60-95 m depth (Prinsenberg,1987). The contributions of the various nutrient sources, the spatial patterns of primary production, and the relative contribution of ice algae and phytoplankton are not well known.

Because of relatively low annual primary production and large river influence, inputs of terrigenous organic matter generally exceed those of marine prirnary production in the interior Arctic Ocean shelf seas (Gebhardt et a1.,2005; Goni et a1.,2005;

l3 Macdonald et al., 1998; Stein and Macdonald,2004c). Coastal erosion represents a second large source of terrigenous sediment and organic matter and indeed, contributes more than river inputs in the Laptev and Kara Seas, where there are strong 'marginal filters' around the rivers (Gebhardt et al., 2005; Rachold et a1.,2004; Stein and

Macdonald,2004a). River inputs may be expected to dominate the OM supply to Hudson

Bay, considering that the Bay receives a proportionally larger river inflow (on the basis of surface area) than any of the interior Arctic shelves, including the Beaufort Sea

(Prinsenberg,l977). However, the quantity and composition of terrigenous OM associated with the river inflow into Hudson Bay are poorly known, with the exception of the Great Whale River in southeast Hudson Bay (Hudon et al., 1996). River inputs of terrigenous OM likely reflect the large watershed (vegetation) differences around Hudson

Bay, with tundra the predominant terrain in the north, Hudson Bay Lowlands in the southwest, and boreal forest in the southeast (Stewart and Lockhart, 2005). Studies of river influence on the coastal zone (e.g., plume dynamics, particulate matter transport, marginal filter effects) are limited almost exclusively to rivers discharging in the southeastern part of the Bay (cf. d'Anglejan, 1980; d'Anglejan and Biksham, 1988;

Ingram, 1981). The processes affecting terrigenous OC transport through and reworking in the coastal zone may be quite different between the western and eastern sides of the

Bay, considering the earlier and larger ice production, larger tides, and generally shallower bathymetry in the west (Saucier et al., 2004).

In terms of coastal erosion, Hudson Bay is exceptional among the Arctic shelf seas in that it has been under the influence of post-glacial isostatic rebound since the deglaciation that took place in the early Holocene (Peltier,2004). Present rates of relative

t4 sea-level fall (coastal emergence) are estimated at-1 crnlaaround the southern part of

Hudson Bay (Hillaire-Marcel and Fairbridge, 1978; Tushingham, 1992). Coastal processes and contributions of sediment and terrigenous OC to adjacent marine areas from emerging coastlines affected by discontinuous permafrost are not as well documented as processes along Arctic coastlines affected by relative sea-level rise (e.g.,

Beaulieu and Allard, 2003).

Deposition and burial of OC in sediments is generally larger (more efficient) and predominantly more terrigenous in composition in the Arctic Ocean than in the other world oceans (Stein and Macdonald,2004c). However, the quantity and composition of

OM buried in sediments in the various Arctic shelf seas varies widely depending on the quantity ofinputs, efficiency oftransports beyond the coastal area, frequency of resuspension, and other characteristics of the regional sedimentation processes. Shelf- basin exchanges, including offshore surface water advection and dense shelf overflows

(convection), transporl both tenigenous and marine OM (and breakdown products like nutrients) offshore to the Arctic Ocean basins from the interior shelves (cf. Goldner,

1999). Resuspension processes likewise transport or at least enhance sedimentation of both types of OM offshore (cf. Forest et a1.,2001; Forest et al., 2008; O'Brien et al.,

2006).In some of the interior shelf seas, buoyancy-boundary cunents transport particulate OM primarily alongshore (Carmack et al., 2006). Sea ice rafting of terigenous OM is a relatively important transport mechanism (relative to marine currents) on the shelves that produce more sea ice and also for longer range transport, e.g., to the interior Arctic Ocean (Eicken, 2004).

l5 Transport and sedimentation of OM in Hudson Bay is likely complex due to multiple possible transport processes, including large-scale estuarine circulation, a strong coastal cur¡ent in the eastern part of the Bay, relatively strong tidal currents, and possibly dense shelf overflows from the northwest shelf,, where sea ice formation rates are very high in a large coastal flaw lead (Saucier et al., 2004). Sea ice rafting may transport OM as well, although the importance of this mechanism for sediment transport is a subject of some debate (Henderson, 1989; Pelletier, 1986) and ice exchanges with Foxe Basin and

Hudson Strait are likely low (Prinsenberg, 1988; Tan and Strain, 1996). The shallowness of the Bay's water column (mean depth -125 m) is conducive to relatively strong vertical carbon flux to sediments from marine OM produced in surface waters (Lapoussière et al.,

2009). Furthetmore, regional studies suggest that waves and tidal cuffents support resuspension and transport of silts and clays frorn the shorelines and coastal areas into the marine basins (Beaulieu and Allard, 2003; Henderson, 1989; Josenhans et al., 1988;

Lavoie eta1.,2002,2008;Zevenhuizen et al., 1994), which should promote sedimentation and burial of OM. Resuspension and transporl of coastal sediments is enhanced in places througlr a combination of ice scour and marine currents (Hequette eI a1.,7999). The impact of these processes for sediment transport and deposition on a larger scale and their consequences for the burial of OM remain open questions.

OUTLINE OF THESIS CHAPTERS

For a region like Hudson Bay, which is so infrequently and sparsely surveyed, a comprehensive assessment of the sources, pathways and sinks of marine and terrigenous organic matter and the processes controlling cycling is difficult to achieve. However, in

2005, the Canadian ArcticNet initiative (lrtretwork of Centres of Excellence,

t6 http://www.arcticnet-ulaval.ca/) provided an opportunity to collect samples of suspended

particulate matter from riverine, estuarine and marine areas, and seafloor sediments in

Hudson Bay. These data are used here to develop a first-order assessment of the inputs

and fates of OM from various sources and the controlling processes in Hudson Bay.

A time series of suspended particulate matter samples was collected from an

estuary in westem Hudson Bay (Churchill River estuary) during March-May 2005 and

organically characterized (chlorophyll a, carbon and nitrogen content, ôr3C, ôrsN¡. These

data were used together with hydrological and meteorological data, physical observations

of the ice cover, models of ice growth, and measurements of salinity, oxygen isotopes

(ô'tO), and nutrients, to study the coupled physical and biological processes occurring in

estuarine and coastal areas of western Hudson Bay during the winter-spring transition

period, an intense and rapidly-evolving time in the OC cycle. A conceptual model for the

functioning of a small Arctic estuary during this transition period was developed, which provides a framework for comparing estuarine effects on plimary production and river plume dynamics within Hudson Bay and between Hudson Bay and other Arctic areas

(Chapter 2 in the thesis).

During September-October 2005, suspended particulate matter samples from rivers and marine waters and 13 sediment boxcores were collected from aboard the

2r0Pb CCGS Amundsen (cruise 0502). The sediment cores were dated using techniques

(see Chapter 5 for details) and the particulate and sediment samples characterized organically (CÆlIratios, ô''C, ðlsN, lignin). Sedimentaly lignin data were exploited to evaluate the origin, character and distribution of terrigenous organic matter in Hudson

Bay (Chapter 3). Bulk proxies were applied to assess the relative contributions of

17 terrigenous and marine organic matter to the sediments (Chapter 4, 5). Degradation rates and burial rates of marine and terrigenous organic matter in the Bay's sediments were estimated by combining the organic compositional data with sedimentation rates (Chapter

4,5).

The new data allow, for the first time, construction of a sediment and organic carbon budget for the Hudson Bay system (Chapter 4). The preliminary budget quantifies the organic carbon sources, reservoirs and fluxes and highlights major controlling processes in the OC cycle and potential sensitivities to change (Chapter 4). An application of the budget as a framework for interpreting the distribution of contaminants

(specifically mercury) in the Hudson Bay system (Hare et al., 2008) supporls the relevance and usefulness of this broad-scale synthesis, despite important data gaps that remain.

Finally, to address major data gaps concerning patterns and controls on marine primary production, the bulk proxies were used to evaluate the environmental (nutrient) conditions under which the marine component of sedimentary organic matter was produced (Chapter 5). A simple biogeochemical box-model budget was constructed following LOICZ modelling guidelines (Gordon et al., 1996) to test the plausibility of the proposed nitrate scheme.

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30 Chapter 2: Sea lce, Hydrological, and Biological Processes in the

Churchill River Estuary Region, Hudson Bay

ABSTRACT

A conceptual scheme for the transition from winter to spring is developed for a small arctic estuary (Churchill River, Hudson Bay) using hydrological, meteorological and oceanographic data together with models of the landfast ice. Observations within the

Churchill River estuary and away from the direct influence of the river plume (Button

Bay), between March and May 2005, show that both sea ice (production and melt) and river water influence the region's freshwater budget. In Button Bay, ice production in the flaw lead or polynya of NW Hudson Bay result in salinization through winter until the end of March, followed by a gradual freshening of the water column through April-May.

In the Churchill Estuary, conditions varied abruptly throughout winter-spring depending on the physical interaction between river discharge, the seasonal landfast ice, and the rubble zone along the seaward margin of the landfast ice. Until late May, the rubble zone partially impounded river discharge, influencing the surface salinity, stratification, flushing time, and distribution and abundance of nutrients in the estuary. The river discharge, in tum, advanced and enhanced sea ice ablation in the estuary by delivering sensible heat. Weak stratification, the supply of riverine nitrogen and silicate, and a relatively long flushing time (- 6 days) in the period preceding melt may have briefly favoured phytoplankton production in the estuary when conditions were still poor in the surounding coastal envilonment. However, in late May, the peak flow and breakdown of the ice rubble zone around the estuary brought abrupt changes, including increased

3T stratification and turbidity, reduced marine and freshwater nutrient supply, a shorter flushing time, and the release of the freshwater pool into the interior ocean. These conditions suppressed phytoplankton productivity while enhancing the inventory of particulate organic matter delivered by the river. The physical and biological changes observed in this study highlight the variability and instability of small frozen estuaries during winter-spring transition, which implies sensitivity to climate change.

INTRODUCTION

Estuaries that are ice covered for substantial periods of the year differ from those that experience no freezing (Macdonald et al., 1999). Ice cover prevents direct wind and wave mixing, which in some instances allows freshwater plumes to spread under the ice over a larger area than under open-water conditions (Freeman et al., 1982;lngram and

Larouche, 1987; Macdonald and Carmack, 1991; Macdonald et al., 1995). Ice cover may also reduce the turbulence associated with tidal currents due to frictional damping at the ice-ocean boundary (Prinsenberg,19S6; Saucier et al., 2004). Within some estuaries, the underwater topography of the ice (ridges) may present a vertical barrier to the spreading of freshwater runoff into the ocean (Macdonald and Carmack, I 991 ; Macdonald et al.,

1995). During spring, when peak flow precedes ice breakup in the estuary, these barriers may force the freshwater to underflood and overflood the coastal ice, often violently

(Reimnitz, 2002). However, heat associated with river discharge in spring may advance ice melt in the coastal zone, perhaps by two or more weeks (Dean et al., 7994; Searcy et al.,1996). The ice growth/melt cycle interacts with the hydrological cycle to control an estuary's salinity, stratification and mixing (Macdonald, 2000; Strain and Tan, 1993).

During wintet, ice formation withdraws freshwater from the underlying water column

32 while leaving behind the brine, promoting mixing and convection, whereas during spring, ice melt adds brackish water, enhancing stratification (Tan and Strain, 1980). Buoyancy associated with river discharge in winter may in some cases preclude convection

(Macdonald, 2000). Collectively, these sea ice processes lend a unique character to high- latitude estuaries and to the biological processes therein (Ingram et al., 1996; Legendre et a1.,1992; Legendre et al., 1996; Macdonald et al., 1987).

Much of our present understanding of estuarine processes at high latitudes comes from studies of large rivers such as the Mackenzie (Macdonald, 2000; Macdonald eL a1.,

1995; Macdonald and Yu, 2006) and Lena Rivers (Eicken et al., 2005;Lara et al., 1998;

Yang et a1.,2002). Studies of small northern estuaries, particularly during the winter months, are relatively scarce despite their collective importance (e.g., Dery et a1.,2005).

A few medium- to small-sized rivers (e.g.,La Grande River in James Bay, Great Whale

River in Hudson Bay, River Svartån in Finland) have been investigated partly because of their biological importance (Fortier et al., 1996; Freeman et al., 1982; Granskog et al.,

2005a; Legendre et al., 1981; Legendre et a1.,1996) and partly because of hydrological changes in some of the drainage basins (damming; see Freeman et al., 1982; LeBlond et a1.,1996).

Hudson Bay is one of the southernmost seas having a complete cryogenic cycle

(Markham, 1986). It receives a cumulative annual discharge of -710 km3 fi'om a large nunrber of rivers (Dery et al., 2005; Prinsenberg,1977). The freshwater yield from rivers, 65-80 crì ovel'the total area of the Bay, is about half the seasonal freshwater input from ice melt, estimated at i40 cm (Prinsenberg,1977; Prinsenberg, 1988). Hydrological alterations (dams, diversions and reservoirs) have motivated a number of estuarine

aa JJ studies, especially in southeastern Hudson Bay and James Bay (Freeman et al., 1982;

LeBlond et al., 1996). However, climate change and its effect on ice cover is becoming an urgent concem (ACIA,2005) and it is likely that northern estuaries will be among the first and most sensitive locations for such change (Ingram et al., i996; Macdonald,2000;

Searcy et al., 1996). Indeed, in Hudson Bay, changes in river discharge (Dery et al.,

2005; Dery and Wood, 2005; Lammers et a1.,2001), ice conditions (Wu et al., 2005), and ice-dependent components of the ecosystem (ACIA, 2005; Stirling et aL.,1999) have already been documented with further dramatic changes projected for the end of this century (Gagnon and Gough,2005; Gough and'Wolfe, 2001; Westmacott and Burn, reeT).

Here, we combine hydrological and meteorological data, physical obser-vations of the ice cover and models of ice growth, and measurements of salinity, oxygen isotopes

(ô'tO), nutrients, and suspended particulate material, to develop a conceptual scheme of the Churchill River estuary, western Hudson Bay, during the winter-spring transition. We follow the temporal evolution of the freshwater budget at two coastal locations, one directly influenced by Churchill River discharge (the estuary), and the other outside the apparent river plume (Button Bay). We describe 1) the effects of the Churchill River discharge on surface salinity and stratification in the coastal environment,2) the physical effect of the ice cover on the extent of the river plume, 3) the effect of the river's sensible heat on sea ice growth/melt, and 4) the distribution and abundance of nutrients and chlorophyll a, as influenced by the estuary's physical structure. This study is the first to investigate afrozen estuary in western Hudson Bay, an envirorunent that differs widely

34 from the better studied eastern coast of the Bay (Ingram, 1981 ; Ingram and Larouche,

1987; Ingram et al., 1996;Larouche and Galbraith, 1989; Legendre et al., 1981).

MATERIALS AND METHODS

Study area

The Churchill River discharges into the coastal waters of western Hudson Bay at the town of Churchill (59"N, 94"W; Figure 2-1) where the mean annual air temperature is

-6.9"C and daily average temperatures range from -27"C in January to 12"C in July.

Although the river's headwaters are in the northern prairies, about 90%o of the river flow upstream of Southern Indian Lake (-320 km above the Bay) is diverted into the Nelson

River to generate power. The lower end of the Churchill River forms an enclosed estuary approximately 13 km long and up to 3 km wide, with a weir located at the upstream end.

In summer, seawater intrudes into the lower section to create a highly stratified estuary

and a brackish (5-10) plume extends for a distance of 4 km or more at low tide. The river plume is typically deflected to the east because of the cyclonic circulation in the bay

(Baker et al., 1994). Previous winter data include heat budget parameters from 1 960-6 1

(Schwerdtfe ger, 7962) and a few measurements related to ice algal production (W.

Bernhardt, pers. comm., 2005). The seasonal ice cover begins to form in October-

November following high sensible heat loss due to strong, cold winds (Barber,1967;

Danielson, 1969; Saucier and Dionne, 1998; Saucier et aI.,2004). The energy flux remains negative through the winter until there is open water to absorb solar radiation in

May-June.

35 Winter-spring sample collection

Winter-spring sampling was conducted between 14 March and 30 May 2005

(Table 2-1). Sampling focussed primarily on two stations (81, 58o 48'N, 94o 11'W and

Button Bay, 58o 48.5'N, 94" 17.2'W) located ca. 1 and 5 km, respectively, from the mouth of the Churchill River estuary but supporting data were collected at sites within the river above the weir (58" 40.6'N,94" 10.7''W ), near the estuary (82,83), within the extended landfast zone (T1-T5) and in the open flaw lead (L1-L5) (Figure 2-l) as detailed in Table 2-1.

t t lDUi"",

'# l,F rnr.¡orRsr tce [--l enou¡roeo rce 7 nuesutc: l++l N¡osn-E PAcK rcc ^\-/\ FLAW LEAD

Figure 2-1. Layout of sampling sites (circles) in the Churchill study area and general ice conditions (March-May 2005).

36 Table 2-1. Summary of sampling

Date Site Station(s) Measurement and samples Winter,2005 14 March Button Bay Water samples 16 March Button Bay Water samples 17 March Estuary EI Water samples 19 March Button Bay Water samples Pre-melt,2005 6 April Button Bay CTD, water samples, ice core 8 April Estuary E1-3 CTD, water samples, ice core l0 April Estuary EI Water samples 11 April River Vy'ater samples 14 April Estuary EI CTD, water samples 14 April Button Bay CTD Early rnelt,2005 27 April Button Bay CTD, water samples 28 April Estuary E1 CTD, water samples 29 April River Water samples 2May Estuary EI,TI-T3 CTD, water sarnples 3 May River Water samples 5 May Button Bay Ice core Peak flow, 2005 19 May Estuary El,T2-T5 CTD, water sarnples, ice core 20 May Estuary L1_L5 CTD, water samples 21May Button Bay CTD, water samples 23 May Estuary E1 CTD 23 May Button Bay CTD Break up, 2005 29 May Estuary Near El 'Water samples 30 May River Water samples Fall,2005 1l Oct 59.03oN, 87.57'W 23 CTD, water samples 12OcT 58.78oN,91.51"W AN02-05 CTD, water sarnples l3 Oct 59.98oN,91.96'W AN0t-0s CTD, water sarnples 13 Oct 59.04oN,94.04'W 25 CTD, water samples 14 Oct Button Bay CTD, water samples 14 Oct Estuary EI CTD, water samples 14 Oct River Water samples I6 Oct 60.45oN, 89.37'W 26 CTD, water samples 16 Oct 61.05oN, 86.21"W 27 CTD, water samples

37 Sampling was conducted either from the landfast ice surface, through a hole

drilled with an ice auger, or from a boat launched from the ice edge (L1-L5). Depths were

measured from the top of the ice. A Sea-Bird SBE 19-plus CTD (conductivity,

temperature and depth) profiler equipped with a transmissometer (C-Star, WET Labs

Inc., OR) was used for water-column profiling. Water samples were usually drawn

through a Teflon-lined stainless steel hose (Swagelok Company, Solon, oH) using a

submersible pump (Shurflo 9300 Series, SHURflo,LLC, Cypress, CA) to fill sample

bottles at the surface. A Kemmerer type water sampler was used on May 2,19 and20

(Table 2-1).

Samples for salinity and oxygen isotope analysis were collected into plastic and

glass bottles, respectively, which were capped tightly and sealed around the lid with

Parafilm to minimize evaporation. Samples were stored in the dark in coolers, taking care

to prevent freezing. Samples for total dissolved nitrogen (TDN), total dissolved phosphorous (TDP), and dissolved organic carbon (DOC) were collected by passing an

aliquot of water through pre-combusted Whatman25 mmGF/F filters, mounted on acid- rinsed polypropylene syringe filter holders. The filtrate was collected into 125-mL

Nalgene bottles and stored cold and in the dark. For nitrate, phosphate and silicic acid analysis by marine methods (JGOFS, 1994), whole, unfiltered samples were placed in the same type of bottle and frozen on dry ice. For dissolved reactive silicate (DRSi) analysis, an aliquot was passed through Nuclepore (0.6 um) filters mounted on polypropylene syringe filter holders and collected in25-mL plastic vials. Samples with low salinity (<5) were stored cold and in the dark (to minimize losses on thawing; (Macdonald and

Mclaughlin,1982)), while others were frozen on dry ice. samples of suspended

38 particulate material were collected on Whatman 25 mm GF/F filters by low pressure (<10 psi) vacuum filtration. Filters were folded, wrapped in pre-combusted foil, frozen on dry ice, and then stored in -30"C freezers until analysis.

Ice cores were collected from Button Bay, E1 and T3 using a MARK II ice corer

(9-cm internal diameter; Kovacs Enterprises, Lebanon NH, USA), immediately cut into

5-10 cm sections with a stainless steel saw. Sections were placed in ziplock bags and melted at room temperature, then subsampled for oxygen isotope analysis.

Fall sample collection

Fall sampling of the Churchill estuary and western Hudson Bay (offshore) was carried out in October 2005 from the Canadian Coast Guard Ship Amundsen and azodiac launch. Offshore water samples for salinity and oxygen isotopes were collected from 5-7 depths at 6 stations (55-244 m, Table 2-1) using the Niskin bottles on the shipboard rosette. Water samples and CTD profiles were collected from Button Bay, E1 and the

Churchill River using the same equipment that we used during the winter-spring program.

Analytical methods

The salinity of samples from the winter-spring program was measured at the

Institute of Ocean Sciences in July 2005 using a Guildline Autosal (model 84008) salinometer and samples from fall 2005 aboard the Amundsen usinga Guildline Autosal

(model 84004). Samples were standardized against Standard Sea Water and repeated determinations indicate a precision of +0.003. Overall uncertainty based on field duplicates was better than +0.02.

39 Oxygen isotope composition, expressed as ôl80 referenced to the V-SMOW

standard, was determined for water samples at the University of Ottawa (G.G. Hatch

Isotope Laboratories) using a Gasbench flushed with a gas mixture of 2Yo CO2 in helium

and a DeltaPlus XP isotope ratio mass spectrometer (ThermoFinnigan, Germany). The

analytical precision was +0.15%o with two pairs of fìeld duplicates showing a precision of

+0.4%o. ðl80 for melted ice samples was determined at the Dating Laboratory, University

of Helsinki, Finland, using a Finnigan-MAT Delta-E mass-spectrometer (Thermo

Electron Corporation, Waltham, Ma, USA) with an overall precision of +0.2o/oo.

TDN, TDP, DOC, and DRSi were measured at the Freshwater Institute (FWI).

TDN and TDP were measured on a Technicon autoanalyzer following photo-combustion

and acidification of the sample, whereas DRSi was measured using a Lachat Flow

Injection (unsegmented) continuous flow analyzer (Strickland and Parsons, 1972). DOC

analysis was carried out on acid-decarbonated subsamples using an OI Model 700

Automated DIC/DOC Analyzer (Stainton et al., 1977). Nitrate CtrO¡ + NOz), silicic acid

(Si(OH)4) and phosphate (POa) were measured at IOS, using a Technicon three channel

autoanalyzer. A conection was applied to all samples for chromophoric dissolved organic matter (CDOM). DRSi (FWI) and silicic acid (lOS) results were in excellent agreement

at concentrations less than 35 pM (average relative standard deviation of 5o/o, r:0.998, n:17) and slightly lower at IOS than at FWI at higher concentrations (RSD:15olo,

=0.9I4, n:8). Chlorophyll a and phaeopigment concentrations (Strickland and Parsons,

1972) were determined on a Turner Designs (1O-AU) fluorometer at FWI in June 2005,

after 24-hr extraction (90%o acetone at 4oC). Organic carbon (OC) was measured on acid- decarbonated filter samples at the University of British Columbia (UBC) using a Carlo

40 Erba NA-1500 Elemental Analyzer (Verardo et al., 1990) and its stable isotope

composition (reported as ôl3C relative to Pee Dee Belemnite) determined by an in-line

isotope ratio mass spectrometer. Total nitrogen (TN) and its isotopic composition

(reported as ô1sN relative to N2 in air) was determined on separate untreated subsamples.

Replicate samples indicate a precision of +0.3o/o and +1 .6%o for OC and TN, respectively.

Other data sources and modeling

Churchill River discharge at the estuary was estimated by adding together the mainstem flow at Swallow Rapids (150 km above the Bay) and two gauged tributaries, the Little Beaver River, which joins just below Swallow Rapids, and the Deer River, which joins about 45 km above the Bay. Discharge data were provided by Manitoba

Hydro (Swallow Rapids) and the Water Survey of Canada (Little Beaver River and Deer

River). Based on advice from Manitoba Hydro (M. Drouin, pers. comm., 2005), we allowed a two-week delay between Swallow Rapids and the Little Beaver and the mouth of the Bay. Water temperature data from a thermistor at the Churchill pumphouse, 18 km upstream from the Bay, were provided by Manitoba Hydro (K. Dobson, pers. comm.,

2005). Meteorological data were collected at the weather station in Churchill by

Environment Canada. Tidal elevations were recorded in Churchill Harbour by the Marine

Environmental Data Services (MEDS), Fisheries & Oceans Canada.

The extent of open water and flooded ice in the lower reaches of the river and estuary area (-60 km2) was estimated from the MODIS "MODO9GQK" BAND 2 surface reflectance image product (near infrared 841-876 nm, 23 L7 m pixel resolution). Images that were cloud-free over the Churchill River were density sliced based on subjectively selected end-member reflectance properties. Four characteristic classes were identified:

4l 1) open water, reflectance

3) flooded snow/ice, mixed pixel (e.g., classes 1 and 4), reflectance 32-12%; and 4)

snow/ice, reflectance >72yo. Open water for heat budget calculations was estimated as the

class 1 area plus 50% of the class 2 area.

Two models of ice growth were considered: the theoretical (quadratic) equation

developed by Anderson (1961), and the one-dimensional thermodynamic model

developed by Flato and Brown (1996) and later modified by Hanesiak et al. (1999). The

Anderson (1961) equation was used to assign a growth timeline to the ice cores for the purposes of reconstructing surface water salinities through the period of ice growth.

Difficulties and uncertainties in this approach are discussed in Macdonald et al. (1995).

The equation relates ice depth (H) to the accumulation of freezing degree days (0) as:

H2 + 5.1H :6.70, e : Ilff-fa)dt, where Tf is the freezing point of the water, Ta is the air temperature, and t is time. The freezing point of the water was estimated assuming a salinity of 32-33 for Button Bay and 15 for E1 (Tf : -1.8"C and -0.8oC, respectively)

(Fofonoff and Millard, 1983). Air temperature data collected at Churchill were used to estimate start date (freeze-up) and the freezing-degree-day record. To estimate error, we applied the same method to data frorn 1960-61 and compared the results to measured ice thicknesses in Button Bay from that period (Schwerdtfeger,lg62). The comparison suggests errors of 1-2 weeks early in the growth season and as much as 3 weeks by the end ofthe season.

The thermodynamic model (Flato and Brown,7996; Hanesiak et al., i999) was used to evaluate onset and progression of melt and to plovide a check on the simple theoretical ice growth curve. The model uses hourly input data (cloud amount (tenths), air

42 temperature ("C), wind speed (- r-'), relative humidity (%) andsnowfall (mm day-')) and a resolution of 50 vertical layers (1 snow and 49 sea ice, (Hanesiak et al, 1999)). The model was run with hourly air temperature, relative humidity, wind speed and cloud amount data for Churchill. Snowfall was taken to be the precipitation (mm day-1, same source) recorded between 13 October 2004 and 31 March 2005 (the period during which maximum daily air temperatures were below OoC). Because the model uses hourly input data, the daily snowfall value was applied to the first hour of each day. The total accumulated snowfall (-0.1 m) in the model run is at the low end of the range measured in late March 2005. The thermodynamic model also requires specif,rcation of the

'effective (ocean) mixed layer depth', which is a measure of the heat storage available during ice-free periods (Flato and Brown, 1996). The value (here 10 m) is chosen which most closely reproduces the probable freeze-up date (-November 1).

RESULTS

Hydrographic conditions

Distinct þdrometeorological conditions characterized the five sampling periods over March-May 2005, leading us to designate these periods sequentially as winter, pr€- melt, early melt, peakflow and break up (Figure2-2).Winter (mid-March) had mean daily temperatures of -20"C and low river discharge. Pre-melt (early April) saw daytime temperatures rise occasionally above zero; there were several heavy snowfalls and two rainfall events (April 3-4 and 10-1 l). Early melt (late April - early May) was a period of rapidly rising river discharge driven by melt in the upper half of the Churchill River basin. Peak flow was reached by mid-May, supported by local melt (Figure 2-2) followed

43 by break up after May 24 with river water temperatures climbing abruptly to more than

6.6"C.

Sampling dates spanned stages of the semi-diurnal tidal cycle (Figure2-2).The variability associated with the tidal cycle was assessed by conducting time series CTD casts on April 14, April27, April 28, May 21, and May 23.In Button Bay, the variation in surface salinity readings between high and low tide was generally in the second decimal place. In the estuary, the variation was about 0.5 during pre-melt and about 5.0 during early melt and peak flow.

Peak Break Pre-melt Early mel¡ flow uP E 4 -c .9)I 0) _Ë õ ! tr 0 20 0) L 3 10 (o L u-- 0 Õ\J Fo E v- o L

O o O) L 600l (c)(( o ñ^ tr¡ : 105 _c r: :iJ ooo.l I $ fia ,ì 1.0 b =(') t: o- LL ,: c ov ,ol- ,: 0.1b : ü. .þ oL 14 21 28 04 11 18 25 09 16 23 o March April May E

Figure 2-2.Tidal elevations in Churchill Harbour, air temperatures in Churchill, and Churchill River discharge (solid line) and temperature (dashed line) measured at the weir over the fTve main sampling periods (winter, pre-melt, early melt, peak flow and break up). Black symbols show specific sampling dates.

44 Observations of the ice

Initially, landfast ice covered Button Bay and formed a2-3 km wide coastal margin (Figure 2-1, -3A). The landfast ice was smooth except along the shore where it was bottomfast or tidally grounded and at the outer edge where it formed ice ridges and rubble. The mobile pack ice, blown about by the wind, was sometimes close to the outer edge of the landfast ice and sometimes several kilometers ofßhore. The level fast ice was about 0.7 to 1.0-m thick with a partial snow cover, whereas in Button Bay the ice was

1.2-1.6m thick with a 0.1-0.3 m snow cover. The rubble was up to 7 m thick (Galley et a1.,2005). The lower Churchill River and enclosed estuary were also completely ice covered (-1.5 m thick at the mouth).

Through pre-melt and early melt the rubble zone gradually grew larger, particularly near the estuary where parls of it appeared to be grounded at low tide. The snow and ice surfaces were affected by rain and new snowfall, decreasing the overall albedo (Ehn et al., 2005). When river flow increased in early melt, visibly coloured, turbid water overflooded at the landfast ice-rubble zone boundary. Above-zero daytime temperatures during the same period were accompanied by the appearance of melt ponds on the ice surface (ENVISAT-ASAR imagery, 12.5 m resolution). The lower Churchill

River (below the Hydro thermistor) and enclosed estuary (60 km2), began to open up, forming initially 2-3 kmz of open water below the weir (Figure 2-38) and increasing to

-11 km2 during peak flow and -38 km2 at break up (Figure 2-38).

Coastal ice conditions around the estuary changed abruptly on May 12-13, when parts of the rubble ice broke off and drifted away (Figure 2-34). The timing coincided with SW winds up to 12 ms-] that lasted more than 5 hours, followed by a3.93 m tide

45 (Figure 2-2).The fast ice in the estuary stayed in place for about 7 days after the loss of the rubble. A return to predominantly freezing temperatures (Figure 2-2) slowed the surface melt but melt water moved from the ice surface to a subsurface layer, suggesting that the ice interior continued to warm and become more porous (Grenfell and Perovich,

2003). Above-zero temperatures retumed on May l9 and the fast ice around the estuary began to decay rapidly. By May 28,the ice at the mouth of the estuary had largely ablated (Figure 2-3A, B). In Button Bay, in situ melting was slower as evidenced by melt ponds remaining on the surface as late as May 27 and a much-dampened retreat of the ice edge (not until June).

46 Ëarly melt Freshet Break up

50 Reflectance (%)

Figure 2-3. Evolution of the (A) landfast ice edge off the Churchill River, showing where rubble ice was lost during the peak flow (white arrorvs), and the fast ice then ablated in the break up (black arrow) and (B) surface reflectance, largely a function of open water, in the Churchill River estuary. Images are modified from A) Quickbird RGB composites, May 16 and27, and B) MODIS "MOD09GQK" BAND 2, April2l,N4ay 16,Nday 29.

47 Models of ice growth

Ice growth curves derived from Anderson's (1961) equation (see section 2.5) and a l-D thermodynamic model (Flato and Brown,1996; Hanesiak et al., 1999) reproduce the observed mid-April ice thicknesses in Button Bay (1 .4-1.6 m) assuming continuous growth after November 1 . Whereas Anderson's ( 1 96 1 ) model indicates growth continuing into early May, the thermodynamic model indicates maximum ice thicknesses in mid-April, followed by about 0.12 m of melt in late April and 0.13 m in the first half of

May. The modeled ice thinning accelerates sharply in the second half of May, leading to break up by June 3, which is about a week and a half earlier than observed break up. The difference between model and observations might arise from using air temperature at the

Environment Canada weather station located slightly inland at Churchill which is likely warmer than over the ice surface in Button Bay. Alternatively, the layer of superimposed ice that formed as a result of rainfall and repeated freezelthaw cycles in late April-early

May, might have played a role in preserving ice volume.

For the estuary, thin ice at the time of coring (May 19) together with Anderson's formulation imply that ice began to grow here in early February. Satellite images confirm that ice cover was lost and new ice began to grow at that time. Anderson's equation and the thermodynamic model yield similar growth curves for the estuary ice from February onward, with maximum thicknesses (0.8-0.9 m) reached in early April.

The thermodynamic model indicates about 0.15 m of melt in the second half of April,

0.27 m of melt in the first half of May, and cornplete melt by }i1ray 20. Similar to the model predictions for Button Bay, modeled melting in the second half of May appears to

48 be l-2 weeks earlier than observations probably because the model does not account for the heat flux associated with river water.

'Water column structure and salinity

Strong changes occured in the water columns of both Button Bay and the estuary in concert with the changes in the ice. Initially, the water column in Button Bay was well- mixed, cold (minimum -1.86'C), and more saline (maximum 33.8) than typical fall conditions, which, in October 2005, were 29-30 in Button Bay and 25 .4-33 .7 offshore

(n:34, depths 0-230 m).By pre-melt, the water column was fresher (33.6) but still at the freezingpoint (-1.84'C) (Figure 2-$.By early melt, the water column had freshened to

32.8, which implies the addition of 0.24 m of freshwater to a 10 m water column (roughly the depth at low tide). By peak flow, the water column reflected the addition of a further

1.29 m of freshwater and had become weakly stratified, with bottom water of 31.5 salinity overlain by a brackish surface layer (salinity 20-25), slightly above the freezing point (0.1-0.3"C).

The water column in the estuary was fresher and more stratified than Button Bay, and had larger seasonal changes (Figure 2-4).Infall2005, the water column at E1 was weakly stratified with a 1.5 m surface layer (salinity 5-15) overlying a gradual halocline down to -5 m depth, and then bottom water of uniform salinity (29).In winter and pre- melt, the surface layer had salinity 17-18 but the bottom water was more saline than in the fall (maximum 33.5). Between winter and pre-melt, the thickness of the brackish layer increased from 4 m to about 6 m. Temperatures were consistently close to the freezing point. By early melt, dramatic freshening had produced a surface layer of salinity 3 to 4 and there was a sharp halocline at 5-6 m. The low salinity plume extended

49 out under the landfast ice to the east (T1-T5; Figure 2-1) with little change in salinity for a distance of at least 8.5 km. By peak flow, after the breach in the rubble zone, the surface layer was almost completely fresh (surface salinity 0-3) but thinner (3-a m).

Similar conditions persisted at E1 and T1-T5 through break up but the low salinity (

(L2-L5), away from the breach, showed relatively high salinity (n-zQ. The freshening in the estuary between pre-melt and early melt implies an addition of 1.73 m of freshwater to the (10 m) water column, with slightly less (0.a9 m) persisting after the breach in the rubble zone.

0 -¡ I Button Bay -5 --t Pre-melt ---- Peak flow I 't'-,\\ -10 I Early melt --- Fall a ¡ v -lC -c ão o -5

-10 E1 -15 0 5 10 15 20 25 30 35 Salinity

Figure 2-4. Salinity profiles for Button Bay and El under the ice in the pre-melt, early melt and peak flow periods (spring 2005) and in open water in the fall of 2005.

Oxygen isotope 1ôt80¡ data 80 The ôl values in Button Bay were similar to offshore waters (-4.1o/oo to -l .7o/oo) except during peak flow/break up. The river and estuary varied seasonally with river water decreasing from -12.3o/oo in the pre-melt to -19.5o/oo during peak flow/break up (Table 2-2).The water ðl80 values show strong, seasonally varying linear relationships with salinity (Figure 2-5A).

Oxygen isotope values in ice averaged -2.2o/oo for Button Bay and --60/oo for the estuary. ôr80 increased with ice depth in Button Bay except for the bottom few centimeters (Figure 2-58). In the estuary, the ice ôl80 was more variable, with a decrease to -8%o or lower at about 0.3-0.4 m, and then a decrease to a minimum of -7I.3o/oo below

0.7 m. Extremely low ðr80 values at the ice surface likely reflected contamination by snow and were neglected.

Seasonal trends in salinity derived from the record in the ice

The ôr8O values in ice (Figure 2-58) were used together with the linear ôr8O - salinity relationship in water and the isotopic fractionation between water and ice determined during this study (2.2 o/oo) (Figure 2-5A) to reconstruct the under-ice salinity throughout winter (Figure 2-6; see Macdonald et al., 1995; 1999 for methodology). To assign an approximate timeline to the derived salinity values (Figure 2-6),we used the ice thicknesses modeled by the Anderson ( I 961) growth curves, with accuracy estimated at

-1-3 weeks.

Table 2-2. Oxygen isotope (ðttO) values (oÁo) for western Hudson Bay (offshore), Button Bay, the estuary surface waters, and the Churchill River, in fall2005 and over the winter-spring sampling periods Offshore Button Bay Estuary Churchill River Fall 2005 -4.1 to -1.7 -3.1 -1 1.4 -13.6 Winter -3.7 to -3.2 -8.2 Pre-melt -2.6 -8.5 -12.9 to -12.3 Early melt -3.3 -17.4 to -15.8 -18.6 Peak flow/Break up -6.1 -17.1to -14.9 -19.5

51 (a) water samples o P -10 ça N Fall F Winter & pre-melt I Early melt I Peakflow& break up

Estuary Button Bay

(b) tce core sections 5 -0.¿ I _c o- -0.8 c)

{)(J -1.2

- t.o -20 -15 -10 -5 0 ô18 o Figure 2-5. ôr80 - salinity relationships for water samples from the Churchill estuary region (a) and õr8O profiles for ice cores from the estuary area (April 8 at El, solid square, May 19 at T3, open square) and Button Bay (April6, solid triangle and May 5, open triangle) (b).

.à so .E ñ U> b20 o E o L810

CN= 01 15 29 13 27 10 24 07 21 07 21 04 18 02'16 Nov Dec Jan Feb Mar Apr May

Figure 2-6. Seasonal trends in surface water salinity in Button Bay and the estuary region (El and T3) derived from ice core records of õr8O.

52 The ðr80 record in the Button Bay ice cores suggests that the salinity at the time of freeze up was 28-30 (Figure 2-6), which is consistent with the fall 2005 observations

(Figure 2-4).The surface salinity in Button Bay then varied between 28 and 31 until the end of January, at which point it linearly increased to about 33-34 in early March. The bottom sections of the cores suggest decreasing salinities in the early melt and peak flow periods. The õr80 record in the estuary ice cores (E1 and T3) suggests variable surface salinity (Figure 2-6).Initially, in early February, the salinity was about 15-71 , but within the first two weeks of growth decreased to 6 at El and 11 at T3. Decreases in salinity

(especially at E1) recur on roughly two-week intervals, perhaps related to tides. A more significant decrease in salinity starts in April and persists until the end of the record.

Freshwater and heat budgets for river water and ice

The contributions of sea ice melt and river water to salinity trends were estimated by treating water samples as mixtures of seawater (SW), river water (RW) and sea ice melt (SIM), and determining the proportions of these water types algebraically from measured ôr80 and salinity (see Macdonald et al. (1995) for methodological details).

Solutions for sea ice melt (SIM) can be positive (additions of sea ice melt) and negative

(additions of brine due to ice formation). Salinity and õr80 values for the seawater (SW) end member (3:32.4| ðl8O: -2.24o/oo) were assigned based on average measurements for western Hudson Bay (offshore) samples in the 30-80 m depth range. For the river water

(RW) end-member, we used the average ôl80 values of the samples collected at the weir during winter (-12.60/oo), spring (-18.7o/oo) and fall2005 (-13.6%o). For SIM, we used the

ôr8O value of the seawater end member plus an isotopic fi'actionation of 2.2o/oo (ice being heavier). We also applied a correction to account for the incorporation of river water into

53 the landfast ice. After Macdonald et al. (1995), we partitioned the ice cores into sea ice

and river ice fractions using the equation: R: (ôttOu -2.2 - ðl8Os¡¡y1)/( ôttOo* - ôt8Os,r)

x H x density of ice/density of water, where H represents the total (liquid) height of an ice

core, R the equivalent height of river water, and ôl8Ou the average ôl80 value within the

core. This approach yields 0.57 m of river ice and 0.33 m of sea ice for the cores from the

estuary and0.49 m and 0.91 m, respectively, for Button Bay. We estimated a õr80 value

for river ice of -11.4o/oo (2.2o/oo greater than the river water). Algebraically, these properties resolve the river ice composition as an apparent mixture of 85Yo river water

(RV/), 79Yo sea ice melt (SIM) and negative 4Yo sea water (SW). The water excluded

during river ice formation, which is slightly lighter, would thus contain an apparent extra

15% RW and 79Yo negative SIM. In the estuary, where the ice comprises 0.57 m of river

ice, the effect would be an overestimate of the fraction of river water in the water column by 0.085 m and an overestimate of ice growth by 0.11 m. In Button Bay, the 0.49 m of river ice-equivalent would lead to overestimates of 0.07 4 m RV/ and 0.09 m of ice

growth.

Button Bay had negative SIM inventories throughout the study implying net addition of brine to water from 0.14 to 1.51 m of sea ice formation (Table 2-3). The estuary started with a zero SIM inventory in the fall and reached a minimum of -1.01 m in pre-melt. The overall sea ice production estimated from the difference between the fall

SIM values and the minimum SIM values is 1.01 m in the estuary and I.37 m in Button

Bay. The ice observed in these areas contains the equivalent of only 0.33 m and 0.91 m of sea ice (respectively), the rest being river ice. Clearly, the local production of sea ice appears inadequate to account for the water column inventories of brine.

54 After winter, SIM values increased both in Button Bay and the estuary, consistent with either melting ice or the loss of brine through advection. In Button Bay, SIM increased 0.91 m from winter to pre-melt and then another 0.17 m by peak flow (Table 2-

3). In the estuary, SIM gained 1.15 m between pre-melt and early melt and then an additional 0.08 m by breakup. The increases during peak flow and breakup were most marked ìn the plume (calculated as the top 4 m, Table 2-3).

The thermodynamic ice model predicted that in situ meh. of the sea ice would start in mid-April and produce 0.25 m by mid-May. In the estuary, melt may also be driven by the heat content of river water. A simple heat budget (Table 2-4) suggests that sea ice melt would begin to appear in the estuary in late April and become very signihcant by the break up (May). Heat (calculated relative to 0'C) from the Churchill River was estimated from discharge and water temperature data (Figure 2-2).We also estimated the potential radiative heat flux to the open water downstream of the weir, which gradually enlarged

(from 2.6km2 in the early melt to 1l km2 in the peak flow and 38 km2 in the break up;

Figure 2-38). Solar radiation was adapted from Danielson's (1969) monthly mean values for Churchill, which agreed well with a short record collected in Button Bay in March-

April 2005 (Galley et a1.,2005). Heat from the Churchill River in the months of April-

May had the potential to melt about 55 km2 of ice 1.5-m thick (Table 2-4). If the calculated ice-melt volumes were distributed over the l5 kmz estuary area they amount to about 0.013 m in the early melt, 0.06 m in the peak flow and 0.73 m by the break up.

Considering that the ice in the estuary was only 37Yo sea ice (tlie remainder being resolved as river ice), these quantities translate into about 0.0045 m,0.022 m, and 0.27 m of SIM, respectively. The latter value (0.27 m) is similar to what the freshwater budget

55 (Table 2-3) shows as the increase (0.33 m) in SIM in the plume between the peak flow and break up periods. In situ melt driven by the surface energy balance (as calculated in the thermodynamic model) would contribute -0.04 m of SIM. The melting of the river ice fraction of the in situ ice cover would make an additional small contribution to the apparent height of SIM. Thus, there is general agreement between the amount of ice melt projected for the estuary by the heat budget and the amount accounted for by changes in

SIM in the water column in the early melt and peak flow periods.

The river water (RW) component of the freshwater budget (Table 2-3) reflects seasonal changes in river discharge (Figure 2-2) as well as the impounding of the plume by the rubble ice zone. Despite a decrease in river discharge, RW increased in the estuary from the winter (1 .95 m) to the pre-melt (3.02 m), while decreasing (1 .06 m to 0.15 m) in

Button Bay. Collectively, these data imply that the rubble zone became more effective at impounding the plume as winter progressed. Taking the area of the estuary as 15 km2, we estimate flushing times (RW inventory/river inflow) of about 4 days in the winter and 6-l days in pre-melt and early melt (Table 2-3).Inpeak flow, RW inventories are lower and flushing times shorter (1-3 days), which suggest that the capacity of the area to hold fi'eshwater had been reached. The break up period had an even lower inventory (4.33 m), presumably because the rubble ice barrier had been breached. For comparison, we can estimate the effect of the tide on the circulation of water through the coastal area.

Assuming a 15 km2 area, the volume of water exchanged with the ocean during each spring tidal cycle (-0.06 km3¡ is larger than the daily river discharge at any point during the study. The volume exchanged in a tidal cycle during the neap phase (-0.02km3¡ is

56 larger than the daily river discharge in the winter and pre-melt periods but only about

40%o of the daily discharge during later periods.

Table 2-3. Equivalent heights (m) of river water (RW) and sea ice melt (SIM) in a 10 m water column in the Estuary (El) and Button Bay and the top 4 m (the plume) at El, and estimated estuary flushing times based on Churchill River discharge Water Column Inventorv Flushins Time' Button Bav Esl :uary l0m l0m Top4m SIM RW SIM RW SIM RW Davs Fall -0.14 0.76 0.00 4.11 +0.07 3.26 0.3 Winter -1.51 1.08 -0.93 r.95 -0.35 1.64 3.8

Pre-melt -0.60 0.r5 1.01 3.02 -0.3 r 1.91 6.4 Early melt -0.50 0.32 Aor 28. hish tide -0.1 6 1_ ) J +0.1I 3.04 4.1 Apr 28, low tide +0.14 5.04 +0.23 3.19 1.1 Mav 2. hieh tide +0.1 3 4.90 +0.19 3.10 2.8 Peak flow -0.42 1.63 +0.1 3 4.55 +0.38 3.65 1.3 Break up +0.21 4.33 +0.11 3.39 1.2 calculated as (i5 km'Area of Estuary x m RV/)/ (Discharge x 86400 s day-').

Table 2-4. Estimated heat budget elements contributing to sea ice melt in the estuary region a Water Riverine Solar Open Radiative Sea lce Melt Templ Sensible Radiation2 Water Heat Flux Equivalenta Heat Area to Open Water' m' s-' deg C J day-r J m-2day km2 J day-' m'day-' m day-l Winter 86 0.22 6.9F+12 1.28+07 0 0 Pre-melt 82 0.05 1.68+12 1.lE+01 0 0 Early melt 28-Apr 123 0.05 2.28+12 2.08+01 2.6 4.68+13 t.8E+05 0.012 02-May 301 0.10 l.lE+13 2.08+07 2.6 4.68+13 2.lE+05 0.014 Peak flow 608 0.18 3.9E+13 2.18+07 ll 2.08+14 9.08+05 0.06 Break up 630 9.62 2.28+15 2.28+01 38 1.58+14 1.1E+07 0.13 'River temperatures measured at CR-30 Pumphouse by Manitoba Hydro 2adapted from monthly mean values, Danielson (1969) 3assuming oassumes albedo : 0.1 ; latent heat of fusion of 334 kJ kg-r, sea ice density of 900 kg m-3; m day-l represents volume distributed over 15 km2 area.

57 Dissolved nutrients and particulate matter

The wintertime coastal seawater in Button Bay was relatively rich in dissolved

phosphorous (1.6 ¡rM), moderately rich in dissolved nitrogen (16.6 pM), and low in

dissolved organic carbon (<50 pM) and silicic acid (9.5 pM). Concentrations of DOC and

silicic acid rose and dissolved phosphorous (TDP) declined during peak flow and break

up, whereas dissolved nitrogen (TDN) remained relatively constant (Table 2-5).

Much of the TDP and TDN was organic. Nitrate represented about 11% of TDN

(n:4) in the seawater; unpublished results for samples (n:5) collected from the same

region in March 1994 show the same nitrate proportion, whereas ammonium represented

43% (W . Bernhardt, pers. comm ., 2005). Phosphate represented 22%-53% (n:5) of TDP

in our samples.

Nutrient concentrations in Churchill River water were in sharp contrast to

seawater, with high concentrations of silicic acid (e.g., 58.5 pM in the pre-melt), TDN

(27.3 ¡t}l4) and DOC (1640 pM in the early melt), and low concentrations of TDP (0.1

prM). River properties changed abruptly during early melt including a30Yo decline in

silicic acid and a more than ten-fold decrease in nitrate (Table 2-5). This was mirrored by

an abrupt loss of the nutrient-salinity relationship in the estuary (Figwe2-7).

Chlorophyll ø concentrations in surface waters (2 m from the top of the ice) also

followed different temporal patterns in the river, estuary and Button Bay (Figure 2-8). In

Button Bay, chlorophyll ø concentrations were consistently low (<0.3 mg m-3; from the winter to the early melt period; they then rose sharply to I .4 ^g*-' in the peak flow. Chlorophyll a: phaeopigment ratios followed the same pattern (Figure 2-8). Li the river,

chlorophyll a concentrations and chlorophyll a: phaeopigment ratios increased between

s8 pre-melt and early melt. In the estuary, chlorophyll a concentrations and chlorophyll a:

phaeopigment ratios were already higher than winter values by pre-melt, and the

concentrations remained high (1.0-2.3 mg m-3) through to peak flow.

Table 2-5. Concentrations (mean (range) in uM) of dissolved nutrients in Button Bay, the estuary surface waters, and the Churchill River, over the winter-spring sampling periods Winter Pre-melt Early melt Peak flow & Break up Total Dissolved Nitrogen (TDN) Button Bay 16.6 17.7 19.2 (1s.1-18.1) (17.0-18.4) Estuary 15.2 23.8 19.1 14.1 (17 .4-41.0) (r7 .e-20.3) (1 1.7-18.s) River 27 .3 31.8 24.8 (2s.3-29.2) (30.s-33.2) Total Dissolved Phosphorous (TDP)

Button Bay 1.6 1.73 1 .1 (1.5-i.7) (1.7r-1.74) Estuary 1.0 0.9 0.6 0.3 (0.8-1.0) (0.5-0.7) (0.2-0.1) River 0.1 0.4 (0.0e-0.13) (0.3s-0.3e) Dissolved Organic Carbon (DOC) Button Bay <50 53 I95 Estuary 101 144 1090 1490 (1 13-215) (1000-1 170) River 756 1640 (1460-1 8 I 0) Dissolved Reactive Silicate (DRSi) Button Bay 9.5 8.3 I .9 1 1.6 (8.e-10.0) (8.1-8.5) Estuary 30.8 31 .5 35.9 26.9 (26.8-3s.8) (28.2-40.1) (24.e-28.1) River 58.5 45.3 (58.4-s8.6) (3e.s-sl.2)

s9 (a) o

o o

f o {D3 (U -= lr É z. I I t 2 t I O Winter & pre-melt I Early melt & peak flow

{b) t LI o

>30 I a lr % o o =..o f -g .9 o Ø20 I o 10 16

010203040 Salinity Figure 2-7. Nitrate - salinity and silicic acid - salinity in Churchill Estuary surface water samples. Note the linear relationships in the winter & pre-melt periods (open circles) and departures from linearity in the early melt & peak flow periods (closed squares).

60 I (a) I Brtt"^€r;l E o D I esruary | o) I O River I É¿ tr o ED n ! * g1o. H ro üu O

(b) c E Ëso) A 'õ" E U 8zo n Ic HË (fI An ^E A ()=1 åo ^

(d) lra 6 '1J o o ÊE (J nun -¿o -30 g

$..ttt o.".*Ñ ***"o **"-* g..p,"oç Period

Figure 2-8. Seasonal trends in surface water chlorophyll ¿ concentrations and ratios of chlorophyll a to phaeopigment, ancl the õrsN and õr3C values of suspended particulate material in Button Bay, the estuary region (81, T1-T5), and the Churchill River.

61 The stable isotopic composition 1ôr3C, õ'tN) of the suspended particulate material in the estuary closely matched the composition of river-bome material throughout the study (Figure 2-8). Both regions show lower ôr5N values (4-5%o in the winter and pre-

melt,2.5-3%o in the early melt and peak flow) and ôr3C values (-29 to -30%oin the winter and pre-melf, -28o/oo in the early melt and peak flow) than Button Bay (6.2-7 .2%o and -

25.4 to -23.5%o, respectively). During peak flow, the Button Bay ôr5N and ôr3C values became closer to the values of river and estuarine material (- 4%o and -28%o, respectively).

The river became an increasingly significant source of particulate material as river discharge increased between pre-melt and early melt. Particulate organic carbon (POC) concentrations (0.2-0.3 mg L-') in the winter and pre-melt were similar in all regions, whereas in early melt and peak flow, POC reached 1.8 mg L-l in the river and the estuary surface water (not shown). These data are consistent with our observations and transmissometer readings of increasing colour and turbidity as river flows increased at the end of April/early May.

DISCUSSION

Conceptual model of the Churchill estuary area during winter-spring

The Churchill River estuary and the coastal area away from the direct influence of the river plume (Button Bay) show a coupled seasonal evolution of the sea ice and the watel column. Both sea ice production and rnelt and river water influence the area's freshwater budget and both local and regional processes modulate the effects. We propose a general scheme to explain the physical evolution of the Churchill estuary region as winter tt'ansforms to spring (Figure 2-9). At the end of winter Button Bay and

62 adjacent regions are covered with stable landfast sea ice extending out to a water depth of about 15 m. The ice, which has been growing since November, has almost reached its maximum thickness (1.2-1.6 m) and is covered with 0.1-0.3 m snow. The seaward margin of the landfast ice is partly anchored and protected by rubble ice that has built up from repeated collisions with the mobile pack. In the lead beyond the rubble zone, wind-driven advection sporadically moves the pack ice away from the shore (Figure2-I,3) after which the flaw lead rapidly refreezes. The water column is at its most saline (33.8), having accumulated brine rejected from in situ ice growth and also by brine advected from the lead and possibly the large polynya to the north (Saucier et al., 2004). Because local production of sea ice appears inadequate to account for the water column inventories of brine (see section 3.7),we infer that ice production in the lead and brine advected shoreward is important for the salinization of coastal water. Intermittently, the lead is up to 10s of kilometers wide which, combined with low air temperatures, supports significant ice and brine formation (Saucier et al., 2004; Schwerdtfeger, 1962). The increasing salinization of Button Bay until March, as recorded by ðl80 in the ice cores irnplies that advection from the latent heat polynya in NW sector of the Bay is more important than local ice production because the latter slows down at the end of winter due to self-insulation of the landfast ice and snow cover, whereas the former produces icelbrine throughout winter (-2.6 mice day-r (Saucier et a1.,2004)).

In the estuary, the low winterlime flow of the Churchill River creates a brackish plume overlying the cold, saline bottom water. The plume spreads under the nearshore landfast ice, which is lelatively thin (-70-90 crn) because ice cover was forced out of the

63 kilomelres 0612 13 14 15 16 17 fo ¡ I ¡ WEIR EST#3 EST#2 1. WINTER <-wlND-> 'zo"c new ¡ce orowth landfast Ìce & rubble /ln flaw'lead river ice & partial snov,/ zone - l mobìle +2 I Þack ice 0 ^^ 1 2 J 4 5 7.5 10 15 20 2. PRE.MELT +-WIND-> -5'C sutace affected by raìn. n"ru ¡." nrÎf,Îåa snow. day-time metr 'ô;órth-. -l Ê ¡1oo¡ru +2 river ice ¡ lãndfãst ice .^¡- \ I oack ice 0 " 1 2 ? 4 5 7.5 l0 15 20

. 3, EARLY MELT over-ftoodino <-WIND-> I -4 "c with 'river 'j waier""-' ^^.^ Çå water betow weir Î,t"i mob¡re +2 -¡iver landfâst ice I .,¡r¡^ralead pa,ck ice 0 ice 1 3days c^ ì*-O.C ^ &4J A5 o 7.5 10 15 20 4. PEAK FLOW breach <-wtND-> -o v rn ooen 6FER llargn rub.ble fiaw mobite +2 tendfesr ìce I tead pàC[ ice 0 I 2 J 4

7.5 10 15 20

5. BREAK UP <-WIND-> :tô?-e'c ',îr patches mobite +2 of ice remaining pack ice

1 2

4 5 7.5 t0 i5 2t

Figure 2-9. Conceptual model of the winter-spring transition in the frozen Churchill estuary area. Transect extends from the weir, at the upper end of the estuary, across the landfast ice zone, and out into the flaw lead. Contours show isohalines (dashed portions speculative).

64 area in early February. Although some rubble lies along the seaward margin of the

landfast ice, there is more in Button Bay where the ice has been stable all winter. The ice

continues to grow through March and early April and the rubble zone around the estuary

builds up; eventually it presents a more effective barrier to lateral spreading and results in

a thicker plume, allowing little river water to disperse to the lead or Button Bay.

Conditions change abruptly in Late April - early May, particularly in the estuary.

Salinization ends and spring freshening begins. As discharge increases, river water

accumulates in the estuary, forming a fresh surface layer several metres thick.

Eventually, freshwater storage inside the rubble zone is overwhelmed by persistent high

discharge and river water leaks through/under the rubble to disperse into the surrounding

coastal zone. River water also overfloods the landfast ice near the rubble zone,

contributing to an overall reduction in surface albedo. The surface energy balance

becomes net positive for the first time, initiating in situ ice melt. The water column still

contains brine from earlier sea ice growth but it is being flushed gradually from the entire

coastal area and much more rapidly from the surface layer of the estuary by the large

volume of river water passing through. The heat of the river water is starting to enhance

local melt, which then contributes a new source of fi'eshwater to the estuary surface

plume.

In the peak flow period (May), there is almost complete freshening of the estuary

surface waters and rapid flushing of the area (< 2 days.) As air temperatures increase and

ice melt progresses, the rubble zone barier weakens, allowing more river water to

disperse into the sunounding coastal area. Heat from river discharge in addition to net positive radiative flux has probably increased the core temperature of the rubble and fast

65 ice, increasing the brine volume (Barber ef al., 1994), and thus weakening the ice. With

the large sails and keels enhancing the rubble's susceptibility to tidal and wind forcing,

the rubble zone eventually breaches, allowing the accumulated freshwater to flow freely

through the gap. In late May, warm air temperatures produce (surface) melt ponds

throughout the landfast ice zone and increase the open water area in the enclosed estuary.

Melt advances particularly fast in the estuary because of rapidly-increasing river water

temperatures, especially after break up. Ice ablation in Button Bay follows a couple of

weeks later.

Biological implications of estuary structure

Ice cover, freshwater residence time, stratification, and the distribution of river-

borne and entrained (marine) nutrients ultimately provide important controls on primary

production in a frozen estuary and the surrounding coastal system. River discharge plays

an important role in the availability of light for primary production, improving it by promoting stratification, or detracting from it because of turbidity and CDOM (Granskog

et a1.,2005b; Legendre et al., 1981 ; Legendre et al., 1996; Welch et al., I 991). Where river water and seawater are complementary in their nutrient composition, a mixture of water masses can provide greater nutrient availability than either water mass alone

(Gosselin et al., 1985; Legendre et al., 1981). On the other hand, sources of freshwater that are impoverished in one or more nutrients can dilute nutrient concentrations in estuary surface waters with possible consequences for algal growth.

We observed relatively high levels of chlorophyll a (0.7-1.7 mgm-3; in the

Churchill Estuary beginning in pre-melt, whereas levels remained low in Button Bay and the Churchill fuver (<0.25 mg m-3). These observations imply that there was a brief

66 period during pre-melt when stratification, light, and nutrient supply within the estuary supported phytoplankton production, while conditions in other areas did not. Pre-melt was characterized by weak stratification, brackish surface water, and a relatively long estuary flushing time. At this time, surface waters were supplied by both marine (P) and riverine Q.{ and Si) nutrient sources (Table 2-5, Figure 2-7). Nutrient ratios calculated from our data (TDN:TDP : 20) or previous (March 1994) measurements of inorganic nitrogenous nutrients (I.{O:+NH¿:TDP : 16) (V/. Bernhardt, pers. comm.,2005) suggest that the limiting nutrient for phytoplankton growth in the estuary would have been phosphorous (critical value 15). However, this marine nutrient would probably be replenished frequently by tidal vertical mixing. In contrast, nutrient ratios calculated for

Button Bay suggest limiting levels of silicate for diatoms (critical value of 1.1 ;

TDN:DRSi:2.0;NO3+NH4:DRSi :2.1; W. Bernhardt, pers. conun.,2005), which could not be alleviated by the tide. The nitrogenous nutrient supply in Button Bay is also more limited than in the estuary, where the river delivers considerable ammonium and organic nitrogen (DON), in addition to nitrate.

Alternatively, the elevated concentrations of chlorophyll a inthe estuary may simply be accumulation of organic material delivered by the river. The variability and the high relative abundance of photosynthetic pigments other than chlorophyll a in the estuary might reflect contributions from sources such as fresh terrestrial detrital material, freshwater algal cells in poor physiological condition, and ice algae. The similarity between river and estuary ðr5N and ôl3C values (Figure 2-8) certainly suggests a similar source. The relatively long (- 1 week) flushing time of the estuary in the pre-melt period implies that there would be an opportunity for river-borne materials to accumulate, even

67 if they were being delivered at relatively low concentrations. Thus, the dominant source of organic matter may have been the river, rather than material produced in the estuary itself, a conclusion also reached by Legendre et al. (1981) for the Great Whale River estuary in southeastern Hudson Bay. Nevertheless, at least in some years, depending on the way that the estuary seasonally evolves, there may be periods of phytoplankton production there in advance ofother areas.

Comparison to other frozen estuaries and broader implications

Our conceptual model of the physical and biological functioning of the Churchill

Estuary presents a strong contrast to the previously studied estuaries of southeastern

Hudson Bay. For instance, the Great Whale River estuary develops a winter plume that covers an aÍea of I 00-600 km2 (Ingram, 1981 ; Ingram and Larouche, 1987; Ingram et a1.,

1996), predictable solely from river discharge, largely because the plume can spread without intenuption for a great distance under the region's great expanse of stable landfast ice (Ingram and Larouche,19871, Ingram et al., 1996). The area also has small tides (Larouche and Galbraith, 1989). In addition to its large size, the wintertime Great

Whale plume is characterizedby a long freshwater residence time and a consistent and continuous salinity gradient along its length. Thus, nutrient concentrations are enhanced by entrainment of seawater into the gradually-spreading freshwater plume (Legendre et al., 1981; Legendre et al., 7996) and ice algae and phytoplankton are predictably distributed in relation to the hydrodynamic controls (Legendre et al., 1981; Legendre et al., 1992; Legendre et al., 7996). Once solar energy increases in May, there is an abrupt spring bloom (Legendre et al., 1981), something that we did not observe in 2005 in the

Churchill Estuary. Another contrast to the Churchill system is the month long delay

68 between freshet in the Great Whale River and the break up of the surrounding sea ice.

The plume stays intact during this period provided the sea ice covers at least 50% of the surface area (Lepage and Ingram, 1991). In the Churchill area, melt may be advanced by the positive feedback inherent in replacing ice/snow cover with low albedo open water.

Alternatively, with its headwaters in southern Canada, the Churchill may deliver a larger sensible heat flux following break up. Regardless, it is clear the Churchill and Great

Whale River estuaries differ in their wintertime structure and function, and would probably respond very differently to change.

The Nelson River estuary in southwestem Hudson Bay seems to be a more extreme analog of the Churchill. It does not cunently form or retain a landfast ice cover but rather discharges its winter flow directly into the coastal flaw lead system. We speculate that the winter estuary must be well mixed due to the lack of ice cover (Baker,

1989; Baker et a1., 1993; Schneider-Vieira et al., 1994), and nutrient rich from marine and riverine contributions. However, primary production or nutrients supporting it would be rapidly exported offshore. If the freshwater export creates stratification in the leads downstream, conditions appropriate for biological production could occur there by early

April. Alternatively, the heat content of the river water and accelerated solar heating due to all the open water could advance regional sea ice melt, which would then promote stratification and provide sufficient light conditions to initiate production, possibly by the end of April. However, positive conditions could be offset by turbidity or CDOM derived from the same river inflow.

Both the Nelson and Chulcliill estuaries export (rather than store) wintertime river discharge, meaning there is buoyancy flux into the surounding coastal flaw leads

69 throughout the winter. In the Arctic Ocean, large river estuaries store winter inflow which

reduces the buoyancy flux into the shelf flaw leads and these then retain the capacity to

support brine-driven convection (Dmitrenko et al., 2005; Eicken et al., 2005; Macdonald, 'Winter 2000). buoyancy flux may in part explain the lack of evidence for penetrative

convection in Hudson Bay, despite circumstances (extensive flaw leads over shallow

shelves) that would favour it. Brine rejected from growing sea ice is also an important

control on the depth of the winter mixed layer in Hudson Bay, driving it down to a depth

of 90 m or more by the end of winter (Prinsenberg, 1987). It is unlikely that the runoff

from one or two rivers could significantly affect Hudson Bay's oceanographic conditions

(LeBlond et al., 1996; Prinsenberg, 1983) but the 1976 diversion of freshwater from the

Churchill to the Nelson River -200 km to the south, and a shift to winter rather than

spring discharge, may alter where penetrative convection occurs, thus affecting Hudson

Bay bottom water renewal, and/or change the depth of the surface mixed layer through the winter. We note that Button Bay had a wintertime salinity of 29.6 + 0.3 (n:5) before the diversion (Schwerdtfeger, 1962), compared to as much as 33.8 in March 2005. The diversion also has shofiened the transit time of river runoff as it moves around the bay's perimeter and reduced the buoyancy in the boundary current between the two locations.

Shifts like these may alter the distribution of terrestlial CDOM and nutrients as well as the retum of marine nutrients to the euphotic zone for this section of coast.

A recent model developed for Hudson Bay (Saucier et al., 2004) produces a significant drift towards less saline and warmer bottom waters which, if one considers

Hudson Bay bottom water to be stable, ffiây indicate an as yet unidentified location/process that produces cold, saline water during winter. Our water column data

70 suggest that the Churchill estuary presently supports about 1 m of sea ice growth. Away

from the tiver's influence, the top l5 m of the water column contains brine reflecting

-2.25 m net ice growth, a number that could be considerably larger if brine has been lost

by convection or advection. Thus, our data suggest that the coastal flaw lead is a particularly important area of ice growth.

The variability in landfast ice conditions, river discharges, and tides in Hudson

Bay is mirrored throughout marginal seas of the Arctic Ocean (Eicken et a1.,2005;

Macdonald et al., 1995; Macdonald and Yu, 2006: Mahoney et al., 2001). Nevertheless, our current understanding of high-latitude estuarine processes relies heavily on studies of a few large estuaries like those of the Mackenzie and Lena Rivers (Dmitrenko et al.,

2005; Eicken et al., 2005) which do not necessarily store their runoff components in the same way as smaller rivers. With climate change projected to be greater and perhaps faster in Hudson Bay, owing to its more southerly latitude and strong terrestrial influence

(ACIA, 2005; Gagnon and Gough,2005; Gough and Wolfe, 2001; Westmacott and Burn,

1997), further study of Hudson Bay estuaries may provide early insight into the effects of change.

CONCLUSIONS

The extent of the landfast sea ice and the rubble ice at the seaward margin of the landfast ice place important controls on the distribution of freshwater in the Churchill

Estuary. Winter river discharge is partially impounded and diverted, forming a distinct, alongshore, under-ice plume, well inshore of where the plume is located in open water.

The extent of uninterrupted landfast ice determines the size and shape of the plume, whereas river discharge dictates the amount of leakage through the rubble ice boundary.

71 Tides probably play an important role in dispersing the river discharge into the

surrounding coastal area but we were unable to quantiff their effects in this study. During

peak flow, the rjver water flushes the estuary, rapidly exporting properties of the river to

the surrounding coastal area. The resulting flushing and circulation patterns provide what

appears to be an important physical control on nutrient distribution and springtime

biological production in the frozen estuary.

The disposition of the sea ice is vulnerable to physical forcing; for instance,

sections of the rubble and fast ice can be abruptly lost during the winter-spring period.

The sensible heat of river water also melts local ice, and contributes to rapid ablation of

the estuary ice cover in the week or two following river break up.

Macdonald (2000) suggested that two important changes face arctic estuaries and

river plumes, one being the quantity of runoff and the other being temperature rise and its

effect on sea ice. While the Churchill River discharge has probably undergone greater

modification recently through hydrological diversion than it faces through future

hydrological change, the ice climates of the lower river and estuary still face significant

change due to direct warming and/or changes in the quantity or timing of runoff (ACIA,

2005). These changes could affect the production and melting of river ice and sea ice, the

impoundment of freshwater and strength and location of estuarine circulation, and the timing and distribution of nutrients and biological production.

All of the world estuaries face change, forced by various factors including altered hydrology, water temperature, species distributions, water diversions and nutrient or organic loadings. For the Arctic's estuaries we may add change in the ice climate to this list. As shown here, sea ice acts as a seasonal control on freshwater pathways, in some

72 cases retaining winter inflow in the nearshore while in other cases allowing most of it to

be exported into flaw leads. The timing of ice cover relative to river hydrology (freshet)

is pivotal for determining how light climate in the estuary interacts with the supply of

nutrients delivered by the river or by entrainment. We suggest that the important role of

ice in arctic estuaries is as a modulator, operating on other sources of forcing including

mixing, gas exchange, light penetration, and secondarily as a habitat (ice algae and

marine mammals). In the Arctic, river diversions and reservoirs have the added

complexity (compared to temperate systems) of altering the interaction between inflow

and ice cover, and thus, for example, shifting flow from freshet, where it strongly impacts

the demise of estuarine ice cover, to winter, where it can spread out under the ice to

deliver material and buoyancy to the offshore. In this study we provide in Figure 2-9 a

schematic diagram upon which realistic models might be constructed to evaluate changes

in forcing and their combined effects.

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79 Chapter 3: Sources, Pathways and Sinks of Particulate Organic Matter

in Hudson Bay: Evidence from Lignin Distributions

ABSTRACT

Hudson Bay is alarge, estuarine, shelf-like sea at the southern margin of the

Arctic, where changes in seasonal ice cover and river discharge appear already to be underway. Here we present lignin data for dated sediments from eleven box cores and evaluate sources of terrigenous carbon, transport pathways, and whether terrigenous organic matter has been influenced by recent environmental change. Lignin yields (0.04 to 1.46 mg/l00 mg organic carbon) decreased from the margin to the interior and from south to north, broadly reflecting the distribution of river inputs. Lignin compositional patterns indicated distinct regional sources with boreal forest (woody gymnosperm) vegetation an imporlant source in the south, vs. tundra (non-woody angiosperm) in the north. Lignin patterns suggest redistribution of a fine-grained, mineral-associated fraction of the southern-derived terrigenous carbon to the nortl'reast part of the Bay and ultimately into west Hudson Strait with the Bay's cyclonic coastal circulation. A small component of the carbon makes it to the central basins of Hudson Bay but most of the terrigenous organic material in that area appears to derive from resuspension of older, isostatically- rebounding coastal and inner shelf deposits. Most modern plant debris appears to be retained near river mouths due to liydrodynamic sorting, with the exception of the southwest inner shelf,, where these matelials extend > 30 km from shore. Temporal changes in the composition of terrigenous organic carbon recorded in most of the southern Hudson Bay cores perhaps reflects increases in erosion and cross-shelf transport

80 from coastal deposits, possibly mediated by change in ice climate. In contrast, temporal changes in the nofthwest may relate to changes in the supply of modem plant debris under recent warlner conditions. On the westem shelf, changes may relate to ice climate and the distribution of northern coastal water andlor changes in the delivery of materials by the Churchill River due to water diversion. Although the cores show evidence of change related to the ice climate, there is little evidence that ice itself transports terrigenous organic carbon within the system.

INTRODUCTION

Terrigenous organic matter plays an important role in global marine systems (Liu et al., 2000) with a specially marked significance in Arctic systems. The Arctic Ocean, which contains 1% of the global ocean volume, receives -11% of the global river discharge (Rachold et al., 2004). Accompanying this fresh water are significant fluxes of terrigenous dissolved (Opsahl et al., 1999) and particulate organic matter (POM), estimated at about 8.1% of the global organic carbon flux fi'om rivers (Rachold et al.,

2004). Coastal erosion supplies an estimated additional one-fifth to one-quarter as much terrigenous parliculate organic carbon (POCr.,r) to the Arctic Ocean and indeed dominates terrigenous carbon fluxes on some shelves (Rachold et al., 2000). The abundant supply of terrigenous matter translates into a uniquely prominent role for teruestrial POM in Arctic

Ocean sediments. Terrigenous (allochthonous) organic matter predominates over marine

(autochthonous) organic matter in sediments from many Arctic continental shelves such as those from the Beaufoú (Goni et al., 2000; 2005), Laptev (Peulve et al., 1996;

Zegouagh et a1.,1996), and Kara Seas (Fernandes and Sicre, 2000) as well as from the

81 interior Arctic Ocean basins (Stein and Macdonald,2004). This latter situation contrasts

sharply with other ocean basins where sediments preserve mainly marine carbon.

Factors controlling the distribution and fate of POC,.,. in marine systems include the nature of the organic matter (e.g., plant debris or mineral-associated materials, fresh vs. degraded), the timing and means of its entry to the marine system, and the efficiency of transport processes therein (e.g., Hedges and Keil, 1995; Henrichs, 1995). For example, ancient and highly altered POM, generally associated with mineral surfaces, may be less vulnerable to remineralization than modern POM or plant debris, thus facilitating its wider transport and burial (Goni et al., 2005). POM discharged from rivers in the winter, when shelves are ice-covered, may experience a vastly different transport and remineralization regime from POM discharged during the open-water and peak marine production period. The mechanism of transport, for instance ice rafting (Reimnitz eT al., 1993) vs. marine currents (Bianchi et al., 2007; Gordon and Goni, 2004), may influence whether terrigenous POM moves from the coast to the interior essentially unsorted, or altered by parlicle removal processes.

Hudson Bay is one of the most southerly extensions of Arctic marine waters, experiencing a complete amual sea ice cover, and, like the Arctic Ocean, receiving very high river discharge (about 30% of the total Canadian runoff.¡. Its coastline spans about

14 latitudinal degrees and its watershed extends west to the Pacifìc watershed and south to the Mississippi watershed of the northern United States (Figure 1). On a background of gradual (geologic) change, including falling relative sea level due to post-glacial isostatic rebound, Hudson Bay and its watershed are now rapidly changing as part of global climate change (ACIA, 2005; Dery et a1.,2005; Tynan and DeMaster, 1997). It is

82 proposed that the consequences of climate change may occur sooner and more strongly in

Hudson Bay than in the Arctic Ocean, owing to its more southerly latitude and close

association with terrestrial systems (ACIA, 2005; Gagnon and Gough,2005; Gough and

Wolfe, 2001; Westmacott and Burn, 1997). Many of the observed and predicted changes,

including altered timing and quantity of river runoff, altered sea ice conditions, and permafrost degradation, are signif,rcant to the supply and fate of terrigenous POM within the system. However, little is known about the composition and quantity of riverine particulate inputs, the importance of coastal erosion, or the freshwater and sediment transport pathways within Hudson Bay. This knowledge is necessary to detect, understand and project the consequences ofchange. There is also uncertainty about the effects of human disturbances in the watershed, including massive diversions and dams

(Messier et al., 1986; Prinsenberg, 1980), to the Hudson Bay system.

In 2005, as part of the Canadian ArcticNet initiative (http://www.arcticnet-

ulaval.cal), we began a study of the recent sedimentary record in Hudson Bay to investigate its modern organic carbon cycle, with the intent of deciphering the impoftance and nature of terrigenous organic matter and how (or ifl it has been influenced by recent environmental change. Here, we present a first assessment of the sources, composition, and distribution of land-derived organic matter in the surface sediments of the Bay. We apply organic geochemical tracers, specifically including eight characteristic lignin phenols and other distinct organic matter products (e.g., 3,5-dihydroxybenzoic acid) produced by alkaline CuO oxidation. Lignin is arnong the most specific tracels for terrigenous organic matter because it is a major component of terrestrial vascular plants but essentially absent from marine organisms (Goni and Hedges,1995; Hedges and

83 Mann, 1979).In general, lignin compounds are intrinsically stable and relatively resistant

to microbial degradation, and are therefore ubiquitous in environments with inputs of

organic matter from vascular plants. Other compounds like 3,5-dihydroxybenzoic acid

are complementary in that they are more general products of terrestrial organic matter

degradation in soils (Christman and Oglesby,l97l; Goni and Hedges, 1995; Houel et al.,

2006; Ugolini et al., 1981). Lignin products in sediments, together with 3,5-

dihydroxybenzoic acid, have thus been widely used to reconstruct terrestrial organic

matter sources and transport pathways in both freshwater (Bianchi et al,2007; Farella et

a1.,2007; Hedges et al., 1982; Houel et al., 2006 Hu et a1.,1999; Lobbes et al., 2000;

Onstad et al., 2000) and marine systems (Goni et al., 1997;2000; Gough et al., 1993;

Hedges and Parker,1979; Miltner and Emeis,2001).

Here, we reconstruct the distribution and character of the major sources of

terrestrial organic matter to Hudson Bay from the recent sedimentary record, and propose

a conceptual scheme for the transport and sedimentation of these organic compounds

within Hudson Bay. We also examine the sedimentary record seeking evidence of

environmental change over the last century. Previous studies of Hudson Bay sediments have focused on the importance of glacial and post-glacial processes using stratigraphic, microfaunal, and palynological analyses from a limited number of sample sites (Bilodeau

et al., 1990), surface sediment texture analyses (Henderson, 1989), or seismic data

(Josenhans et al., 1988). We provide here the first information on composition,

distribution and trends of modern organic biomarkers in Hudson Bay sediments from 11

dated box cores.

84 METHODS

Sediment sampling

Sediment cores were collected during the 0502 cruise of the Canadian Coast

Guard Ship Amundsen (September-October 2005). Along the cruise track, coring sites

(Figure 1) were selected from bathymetric and sub-bottom data gathered by the 8M300

&. 3 . 5kHz transducer array aboar d the Amundsen (http : //www. omg. unb. ca).

Figure 3-1. Maps showing the Hudson Bay watershed, major rivers and the position of the tree line (after Rouse et al., 1997; Stewart and Lockhart, 2006) and the Iocation of the l1 sediment cores examined in this study (filled circles labelled with core number).

85 Cores were retrieved using a20 cm x 30 cm box corer, which penetrates the seafloor to a maximum of 50 centimeters. The cores were sectioned aboard the ship, generally into 1 cm intervals for the top l0 cm,2 cm intervals for the next 10 cm, and 5 cm interuals for the remainder of the core. One core (i 1) was sectioned in2.5 cm intervals between 5 and 15 cm and 5 cm intervals below that. Each section (sample) was homogenized in a thoroughly pre-cleaned 500 mL l-Chem glass jar with a Teflon-lined lid. After homogenization, subsamples intended for elemental and isotopic analyses were sealed in Whirlpac bags or 20-mL plastic vials. Samples were stored in a freezer (-20"C) on board the Amundsenuntll the end of the cruise, then shipped to the Freshwater

Institute (FS¡I), where they were maintained in storage at -30'C.

Laboratory analyses

At FWI, samples were subsampled for moisture determination, then freeze dried and rehomogenized. Radioisotope analyses were conducted at the Environmental

Radiochemistry Laboratory, Soil Science Department, of the University of Manitoba. t"Cs was counted on a gamma spectrometer using a hyper-pure germanium crystal. The counting efficiency was determined using standard reference materials distributed by the

U.S. Environmental Measuretnents Laboratory as part of the Quality Assurance Program

t'oPb 2rOPo (QAP50-QAP60, soil) and reference soil samples. was determined via its daughter. The samples were prepared according to the procedure detailed in Flynn

(Flynn, 1968) and counted on an alpha spectrometer using a silicon surface barrier

20ePo t'OPo detector. A tracer, calibrated against a NIST standard (Isotope Product

Laboratories, #6310) was employed for quantitation. The enor in counting replicate samples was under 7yo.226kawas determined using the radon de-emanation method

86 226Rastandard (Mathieu, 1977) and counting efficiencies checked using NIST NBS

SRM 495 3 . C). Replicat e "íRameasurements were generally within 20%o.

Organic carbon content was measured at the University of British Columbia flJBC) using a Carlo ErbaNA-1500 Elemental Analyzer (Verardo et al., 1990), following acidification to remove inorganic carbon. The relative precision of the method is +0.3%.

Lignin-derived phenols and 3,5-dihydroxybenzoic acid yields were measured at

Oregon State University using alkaline CuO oxidation in a microwave digestion system

(Goni and Montgomery, 2000). Freeze-dried sediment (300-500 mg) was reacted with

CuO and NaOH (2M) in N2-pressurized Teflon vessels for 90 minutes at a temperature of 150'C. After the oxidation, recovery standards (ethyt vanillin, trans-cinnamic acid) were added, solid residues were separated by centrifugation, and the aqueous hydrolysates were acidified with concentrated HCI to pH 1. Oxidation products were then extracted twice into ethyl acetate, dried under N2, and redissolved in pyridine.

Immediately prior to injection, samples were derivafized with bis trimetþlsilyl

oC. trifluoroacetamide (BSTFA) +l%otrimethylchlorosilane (TMCS) for 10 min at 60

Quantification of oxidation products was by gas chromatography-mass spectrometry

(GC-MS) with the MS operating in selective ion monitoring mode. Chromatographic separation was achieved on a 30 m x 250 ¡rm DB | (0.25 ¡rm film thickness) capillary oClmin column, using an initial temperature of 100 "C, a temperature ramp of 4 and a

oC. final temperature of 300 External calibrations were performed to test the response of the GC-MS and were highly linear ß2:0.99) over the concentration ranges measured in the samples. Quantified reaction products included eight derived from lignin: vanillyl (V- series) phenols (vanillin, acetovanillone, vanillic acid), syringyl (S-series) phenols

87 (syringealdehyde, acetosyringone, syringic acid), and cinnamyl (C-series) phenols þ-

coumaric acid, ferulic acid), and one (3,S-dihydroxybenzoic acid) derived from non-

lignin sources. On average, the error in the yields of CuO products is 5-20Yo of measured

value, with the highest errors associated with the compounds with lowest yields. The

detection limit for individual CuO oxidation products was about 0.001 mg/l0 g sediment

and total lignin yields were calculated as the sum of the detected vanillyl, syringyl and

cinnamyl phenols. Vanillin (Vl), vanillic acid (Vd), syringealdehyde (Sl), and 3,5-

dihydroxybenzoic acid (3,5-Bd) were detected in all samples, andp-coumaric acid þ-Cd)

and syringic acid (Sd) were detected in the majority (all but 6%o and I4o/o, respectively).

Acetovanillone (Vn) and acetosyringone (Sn) were below detection limits in about a quafer (23%-27%) of the samples, particularly those from the central basins (cores l0 and 14) and northwest slope (core 13), which had trace total yields. Ferulic acid (Fd) was below detection limits in the majority of samples (7I%), including core 7 and the cores from the interior (10, 14), west (11), and northwest (12,13). Among the samples with the lowest lignin yields, especially cores 12, 13 and 14, it is possible that the amino acid tyrosine was present in sufficiently high concentrations (relative to lignin) to contribute to the measured p-Cd yields (e.g., Hemes and Benner, 2002). Poor correlation between p-

Cd and a well-established tyrosine product,p-hydroxyglyoxalic acid (Goni and Hedges,

ßgs),suggests that the influence of any tyrosine-derived p-Cd is not widespread among the Hudson Bay sediments. Nevertheless, the p-Cd values for cores 12,73 and 14 should be considered maximum values for lignin-derived p-Cd.

88 Geochronologies

2loPb Geocluonologies were constructed for the cores using profìles of activity,

which generally decreased logarithmically beneath well defined surface mixed layers

2rOPb (SML) (Figure 2). Excess or unsupported was estimated from the 2rOPb activities

226Ra deep in the cores and activity measurements on several sections from each core.

The cores were then modeled using a simple two-layer advection-diffusion model

(Macdonald et a1., 1992), with a rapidly diffusive top layer overlying a much less

diffusive deep layer. Analogous to the 'simple' model (Robbins, 1978), sedimentation

velocities (w) in units of cm a-' were calculated from the slope of the linear regression of

2r0Pb the ln (excess activity (dpm cm-3)) in the deep layer vs. sediment depth (cm) (Figure

2).Dry mass sedimentation rates were calculated from the average porosity within a core

and an assumed mean density of sediment solids of 2.65 g cm-3. Sediment-mixing rates

for the two layers (Kbl and Kb2, respectively) in units of cm2 a-l and the excess2r0Pb

activity of depositing sediment parlicles (dpm cm-3) were chosen to minim ize the mean-

2lOPb squared eror between the measured and modelled activities. Estimated

l37Cs sedimentation rates were verifìed against the profile for each core, except for cores

l37Cs 12 and 15, which contained only in the l-2 cm interval. Sedimentation rates were also checked by comparing apparent surface fluxes (calculated as the product of sedimentation rate (crn a-r) and specific activity (dpm cm-3)) with those implied by the

t'OPb excess inventories in each core (product of inventory (dpm cm-2) and the decay constant (0.031 t+ a-l¡¡. The major uncertainties associated with the geochronologies are

226Ra 2roPb 1) variation in activity measurements and hence supported activities, 2) low

89 2loPb total activities in coarse-grained sediments, which affects cores 9, 11 and especially

7, and is also a source of variation in cores 6, 13, 14 and 15.

a 10 / a

15 a I 15 20

ol = 0.06 o = 0.09 ro = 0.05 or = 0.18 25 40 0 0 10 11 a

5 4

()E 10 10 .c I oí) l5 15 10 20 20

o = 0.10 al = 0.15 al = 0.15 25

1 15 ,I 2 ) 4 ,/ a/ ./ 6 ,/ I a I :

12 o = û-14 o = 0.06 14 -3 -2 -1 0 r 2 .3 -2 ..1 0 1 2 -3 -2 -1 0 1 2 Ln (excess 21oPb) 2rOPb Figure 3-2. Profiles of the natural log of excess activity in the secliment cores. Points represent measured values while lines represent the model results. Supported 2lOPb levels, surface mixed layer (SML) depths and other parameters are listed in Table 3-1.

90 RESULTS

Sediment properties and accumulation rates

The eleven sediment core sites were widely distributed, including west Hudson

Strait (core 15, 430 m water depth), Hudson Bay's pseudo-shelves (eight cores from water depths of 34 to 153 m) and the main central basin of the Bay (cores 10 and 14,200-

244 m water depth; Figure 3-1). Two cores were collected from what we consider to be the inner shelf (<30 km from shore) and the others from the mid- to outer shelf (-100 to

290 km from shore). The cores varied texturally from a mixture of sand and silt (core 9, from the southwest inner shelf) to almost exclusively silt and clay (cores 10 and 14, from the central basins). The average organic carbon (OC) content of the cores varied from

0.53%to l.l4Yo (Table 3-1) and increased with water depth (R:0.77, p:0.010).

2roPb Sedimentation rates determined from profiles (Table 3-l; Figure 3-2) were highest (0.25 gc--'u-'¡ along the shallow southwest inner shelf,, followed by the western shelf (core 1l). The lowest sediment accumulation rates (0.04-0.07 g cm-2 a-r) occurred in the central basins (cores 10 and l4), northeast (core 4), and south, near the mouth of

James Bay (core 7), and also in Hudson Strait (core 15). Shelf cores in the southeast

(cores 6 and 8) had sedimentation rates of 0.07-0.13 g cm-2 a-r and those in the northwest

(cores 72 and 13) 0.14 g cm-2 a-t lTable 3-l). The analyzed samples within each core thus represent sediments that have accumulated over periods of 56 years (core 9) to more than

200 years (core 7), while 'surface' samples (0-1 cm) represent periods of 6 to 22years.

Particle residence times in surface mixed layers and hence 'intrinsic time resolutions' of each core, calculated as the depths of the mixed layers divided by the sedimentation

91 velocities (Robbins, 1978), varied from about 6 years in the southwest inner shelf to 20-

60 years in other shelf areas and the central basins (Table 3-1).

Total Iignin yields and accumulation rates

Total lignin yields varied from 0.0043 to 0.0590 mg/l0 g sediment expressed on a salt-free dry mass basis (X8; Hedges and Mann, 1979) and from 0.042 to 1.460 mg/100 mg OC on an OC-normalized basis (r\8; Hedges and Mann, ï979). Variation within each core was relatively low (9%o-33% RSD), except for core 9, which had twice the lignin yield at surface, compared to the subsurface (Figure 3-3). Mass-based yields increased slightly toward the surface (more recent) sediments of cores 13 and 14 (p<0.027), whereas OC-normalized yields decreased significantly toward the surface in core 8

(p:0.025). The top two sections of core 6 had 37%o lower yields than the underlying sections (Figure 3-3).

Spatially, the cores with the greatest lignin yields were those in the south, especially southwest, while those with the lowest yields were in the centre of the Bay

(Figure 3-3). Among the shelf sediments, the highest yields were in the southeast (core

8), followed by the nofthwest (cores 12 and l3), norlheast (core 4), and Hudson Strait

(core 15). Overall, the average yields from each core decreased with increasing distance from shore and from south to north (Figure 3-3); yields were not correlated with sediment accumulation rates.

92 Table 3-1. Core properties and sedimentation parameters

Core Water Distance Sediment texture Organic SML Range of Supp. ''uPb depth offshore Carbon (crn) "uRa ldpm (dp. g-') (rn) (k-) (%oC) s-') 4 153 100 Silt and clay 1.14 3 2.2-2.9 2.58

6 119 25 Silt, clay, some sand 0.70 5 0 .43-1.2 0.90 1 106 100 Sand-silrclay 0.58 I 0.92-1.6 0.90

8 I 50 200 Silt and clay I .10 6 1.9-2.2 I .86 9 34 25 Sand-silt mix 0.53 I 0.33-0.57 0.60

l0 200 260 Silt and clay 1.02 3 0 .58-'12.3 4.62 I I 86 100 Sand-silt-clay 0.66 2 0.16-0.63 0.54

12 I 16 160 Silt, clay, some sand 1.08 5 0 .62-1.3 1.02 13 145 290 Silt, clay, some sand 0.80 3 0.45-2.3 1.014 14 244 410 Silt and clay 1.14 2 3.5-1 1.3 9.30

15 430 10 Silt, clay, sand 0.83 I 0 .14-0.51 0.90

Core Specific Sediment. Sediment accum. rate Excess '''Pb Deepest Time activity velocity (g cm-2 yr-r) inventory section with resol'n (dprn g-') (cm yr r) (dpm cm-2) r37cs (years) 4 20.4 0.06 0.04 17.0 3-4 cm 50 (0.0s-0.08) (0.03-0.05) 6 32.2 0.09 0.07 46.5 7-8 cm 5l (0.06-0.r r) (0.0s-0.0e) 1 11.5 0.05 0.05 20.9 2-3 cm 22 (0.02-0.31) (0.03-0.38) 8 32.3 0.18 0.13 66.2 14-16 cm 34 (0.16-0.20) (0.r r-0.20) 9 3.2 0.18 0.25 24.2 8-9 cm 6 (0.1s-0.22) (0.2r-0.3r) l0 19.8 0.10 0.06 27.2 3-4 crn 30 (0.06-0.12) (0.04-0.08) I I 8.1 0.15 0.20 43.4 7.5-10 cm 13 (0.12-0.2r) (0.11-0.2e) 12 7 .2 0.15 0.14 l9.l only at l-2 34 (0.12-0.19) (0.12-0.18) cm 13 8.4 0.14 0.14 18.6 4-5 crn 2l (0.0e-0.16) (0.0e-0.16) 14 9.0 0.06 0.03 9.8 3-4 cm 35 (0.04-0.06) (0.02-0.04) 15 4.5 0.06 0.06 11.4 only at I -2 18 (0.04-0.09) (0.04-0.09) cm Cores are saìt-corrected for bottom water conditions at the time of sampling. Surface mixed layer (SML) 2'oPb was determined by eye from the profile. Sedimentation velocities (and 95% confidence limits) were 2roPb¡ estimated from the slopes of Ln (Excess vs. depth; accurnulation rates were determined from sedimentation velocities (see methods for details).

93 I L L L L J,,, I þ¡ l¡ HUDSON I åto BAY

Lignin yield oE -o Õ 14

1.2 B C o 1.0 o Ëot O I P o.e cr) I,l 5 0.4 ro f õ ¡ I F 0.2 It ri T I r ll I rìr 0.0 0 100 200 300 400 500 55 56 57 58 59 60 61 62 Distance from shore (km) Latitude ('N)

Figure 3-3. Total lignin (z\8) (yields (mg/100 mg OC) in sediment core sections (A) and relationships with distance from shore (B) and latitude (C) (mean *SD).

94 Table 3-2. Yields (mg/100 mg OC) of lignin phenols, total lignin (.t\8) and 3,5- dihydroxybenzoic acid (3,5-Bd)

Core Sec'n %oc VI Vn Vd SI Sn Sd p-cd Fd 3,5-Bd (cm) ^8

0-r r. r8 0.023 0.008 0.0r 2 0.0r 7 0.0r 3 0.02 t 0.09 0.025

t-2 r.18 0.019 0.009 0.0r4 0.022 0.012 0.0t I 0.025 0. r 1 0.029 3-4 1.14 0.016. 0.008 0.013 0.026 0.013 0.01I 0.027 0.13 0.0 r3 0.032 5-6 t.l4 0.0r9 0.010 0.0r4 0.034 0.0r4 0.0r4 0.029 0.13 0.034 7-8 r.r0 0.016 0.009 0.012 0.033 0.014 0.012 0.026 0.12 0.030 9- 10 t.09 0.0r4 0.009 0.0r r 0.024 0.012 0.012 0.028 0.1 I 0.035 0-r 0.79 0.092 0.029 0.059 0.030 0.016 0.016 0.069 0.33 0-0 r 9 0.055 1-2 0.79 0.093 0.029 0.069 0.045 0.020 0.078 0.33 0.066 3-4 0.70 0.098 0.040 0.086 0.088 0.026 0.034 0.r05 0.5I 0.030 0.080 5-6 0.68 0.r09 0-039 0.088 0.085 0.027 0.035 0. r 04 0.s2 0.033 0.076 7-8 0.65 0.r06 0.034 0.082 0. r 59 0.024 0.032 0. r 07 0.58 0.033 0.070 9-r0 0.60 0.091 0.034 0.071 0. r 02 0.026 0.03 t 0.091 0.48 0.031 0.068 0-r 0.58 0.054 0.026 0.044 0.070 0.029 0.033 0.054 0.31 0.046 1-2 0.67 0.046 0.023 0.038 0.070 0.028 0.048 0.25 0.038 3-4 0.58 0.056 0.022 0.038 0.074 0.028 0.048 0.27 0.034 5-6 0.74 0.044 0.018 0.028 0.093 0.023 0.040 0.2s 0.027 7-8 0.s7 0.042 0.022 0.028 0.065 0.022 0.028 0.048 0.26 0.032 9-t 0 0.36 0.086 0.05| 0.r68 0.055 0.095 0.46 0.049 0-r 1.28 0.033 0.0r 3 0.023 0.028 0.0 r s 0.014 0.032 0. r 6 0.034 l-2 1.16 0.035 0.0 r 5 0.026 0.039 0.013 0.0r6 0.038 0.20 0.0r 6 0.037 3-4 L12 0.035 0.0r 5 0.024 0.036 0.012 0.0r6 0.040 0.r9 0.0r5 0.034 5-6 r.09 0-039 0.017 0.029 0.043 0.012 0.0r5 0.04r 0.21 0.017 0.037 7-8 1.06 0.044 0.021 0.033 0.049 0.022 0.026 0.052 0.27 0.028 0.045 9- l0 0.98 0.045 0.0 r 9 0.033 0.052 0.01 7 0.021 0.049 0.26 0.022 0.045 t2-14 0.99 0.043 0.018 0.032 0.05 r 0.016 0.02 r 0.046 0.25 0.021 0.043 0-r 0.40 0.608 0.144 0.259 0. r 95 0.052 0.053 0.092 1.46 0.057 0.068 1-2 0.'70 0.127 0.042 0.083 0.092 0.032 0.037 0.062 0.5 r 0.036 0.046 3-4 0.62 0.r38 0.042 0.07 r 0.rr0 0.033 0.04 r 0.079 0.55 0.039 0.049 5-6 0.49 0.r87 0.055 0.086 0. r28 0.037 0.047 0.077 0.66 0.042 0.049 7-8 0.48 0.r40 0.042 0.065 0.212 0.046 0.044 0.075 0.67 0.046 0.038 9-r0 0.48 0.r28 0.043 0.075 0.187 0.044 0.053 0.085 0.67 0.055 0.046 r0 0-r 1.12 0.013 0.008 0.0 r 3 0.0r8 0.0r 6 0.07 0.020 1-2 1.07 0.021 0.007 0.0 r 6 0.0r6 0.0 r 3 0.07 0.0r9 3-4 r.0r 0.027 0.008 0.0r6 0.05 0.0r4 5-6 1.02 0.014 0.009 0.026 0.0r5 0.0r4 0.018 0. r 0 0.022 7-8 0.95 0.012 0.008 0.029 0.014 0.0t 5 0.0r8 0.10 0.022 9-r0 0.95 0.0 t2 0.009 0.029 0.0r6 0.0r5 0.019 0.10 0.022 lt 0-l 0.66 0.029 0.014 0.018 0.067 0.025 0.02 r 0.052 0.23 0.046 1-2 0.026 0.0r4 0.0r6 0.05I 0.02 t 0.020 0.050 0.20 0.044 z-J 0.027 0.0r 6 0.0 r 8 0.049 0.024 0.05r 0.r9 0.046 3-4 0.03 r 0.0r7 0.02 r 0.049 0.025 0.050 0. r 9 0.046 5-7.5 0.028 0.0r6 0.020 0.054 0.021 0.022 0.046 0.2 r 0.04 r 7.5- r0 0.035 0.0r7 0.026 0.0s3 0.021 0.040 0. r9 0.041

95 Table 3-2. (continued)

Core Sect'n VI Vn Vd SI Sn p-cd Fd 3,5-Bd (cm)

12 0-r r.3 r 0.014 0.008 0.009 0.025 0.0 r 9 0.012 0.033 0.12 0.042 l-2 0.96 0.01'7 0.010 0.0r0 0.032 0.018 0.0r2 0.042 0. 14 0.048 2-4 0.99 0.0r4 0.008 0.009 0.025 0.014 0.012 0.035 0.12 0.042 4-6 1.04 0.0r4 0.008 0.01 0 0.028 0.0r r 0.013 0.037 0.12 0.0s2 6-8 r.r0 0.013 0.008 0.0r0 0.022 0.0r5 0.0r3 0.035 0.12 0.045 8-r 0 r.05 0.014 0.009 0.012 0.020 0.021 0.0 r 4 0.035 0. r 3 0.052 t3 0-r 0.95 0.0r3 0.008 0.029 0.0r3 0.0r3 0.02I 0.r0 0.022 t-2 0.98 0.017 0.010 0.034 0.0 r 5 0.024 0. 10 0.027 3-4 0.87 0.0r 0 0.008 0.021 0.0 t 3 0.019 0.07 0.020 5-6 0.75 0.0r 0 0.009 0-020 0.01 4 0.02 I 0.07 0.020 7-8 0.69 0.017 0.013 0.037 0.02 r 0.09 0.020 9-t0 0.56 0.0r 8 0.014 0.044 0.08 0.023 14 0-l r.l8 0.020 0.0 r 0 0.020 0.019 0.020 0.09 0.024 1-2 1.14 0.0r 5 0.009 0.022 0.0r9 0.013 0.0t 9 0.l0 0.025 3-4 r.09 0.016 0.009 0.0r9 0.0r 9 0.01 7 0.08 0.022 5-6 1.02 0.0r 4 0.008 0.024 0.014 0.01 I 0.01 5 0.09 0.0r7 7-8 l.28 0.012 0.006 0.01 2 0.012 0.04 0.0r 5 t5 0-t 0.85 0.020 0.0 r 2 0.0r 6 0.029 0.021 0.017 0.023 0. r6 0.020 0.0s2 1-2 0.86 0.017 0.010 0.012 0.029 0.017 0.0 r4 0.020 0.t2 0.048 3-4 0.73 0.02r 0.013 0.0r6 0.03r 0.0r9 0.019 0.028 0. 15 0.064 5-6 0.87 0.017 0.0r3 0.014 0.023 0.0r 8 0.01 7 0.023 0.l3 0.062 7-8 0.86 0.019 0.012 0.01 7 0.024 0.020 0.017 0.024 0.r3 0.059 9-r 0 0.80 0.018 0.01 I 0.013 0.02'7 0.01 3 0.01 6 0.021 0.12 0.060

Lignin accumulation rates within the cores are shown in Table 3-3. To account for differences in sedimentation among the cores (i.e., surface 0-l cm sections representing sedimentation over 6 years in some cores and 22years in others), the rates (mg m-2 a-l and g m-' a'',respectively) were calculated for thlee different time period s (1875-1925,

7925-1975,1915-2005), using the lignin yields (mg/10 g sediment) in sediment intervals coresponding to these periods, multiplied by sediment accumulation rates for each core

(g "--t a-r: Table 3-1).

96 Table 3-3. Lignin accumulation rates (-g --t a-t¡ fo. three time periods (1875-1925, 1925-1975,1975'2005) calculated from tignin yields (mg/10 g sediment) in sediment intervals corresponding to these periods multiplied by the sediment accumulation rate within each core (Table 3-1)

Core 1875-1925 1925-1975 1975-2005 4 0.59 (n.a.) 0.s4 (0.03) 0.44 (n.a.) 6 2.48 (0.47) 2.6e (0.01) 2.03 (0.00) 7 0.74 (n.a.) 0.83 (n,a.) 0.87 (n.a.) 8 3 .37 (0.33) 2.93 (0.18) e 8.12 (0.0s) 10.8 (3.40) 10 0.58 (n.a.) 0.s I (0.16) 0.50 (0.01) 1l 2.6s (0.14) 2.10 (0.28) l2 1.84 (0.04) 2.00 (0.3 l) 13 0.80 (0.13) t.23 (0.29) 14 0.31(n.a.) 0.36 (0.06) 0.38 (n.a.) ls 0.6s (0.02) 0.61 (0.03) 0,77 (n.a)

Bracketed values reflect one standard deviation about the depth-averaged values.

About 640/o of the variation in lignin accumulation rates was explained by variation in sedimentation rates. The southwest inner shelf (core 9) had disproportionately large lignin accumulation rates (maximum 10.8 mg ^-' al;Table 3-3). Other shelf samples, both from the southeast (excluding core 7) and northwest, had the next highest accumulation rates (2.0-2.9 mg m-2 a-r: Table 3-3). The lowest accumulation rates were in the central basins (cores 10 and 14) and northeast (core 4). The rates were generally quite similar over the three time periods, decreasing slightly (18%-25%) in cores 4 and 6 and increasing slightly (20%) in cores 14 and 15. Between the two more recent periods (1925-1975 and 1975-2005), there were larger increases in accumulation rates in the northwest (core 13: 55%) and the southwest (core 9: 33%).

Lignin composition

Five lignin compositional ratios were calculated for each sediment sample: syringyl to vanillyl (S/v), cinnamyl to vanillyl (c/v), 3,5-Bd to vanillyi phenols (3,5- Bd/V) and the ratios of acidic to aldehydic moieties for the vanillyl ([AdiAl]v) and

saingyl families ([Ad/Al]s). Some samples displayed individual ketone (Sn and Vn)

yields that were below detection limits, including selected horizons from cores 7,I1,70,

13 and 14. For these samples, in order to maintain consistency in the way S/V and C/V

ratios were computed, we assigned a value of 20% of the sum of the other measured

lignin phenols of the same family on the assumption that ketones are generally

conservative (20%) relative to the acid and aldehyde products of their respective families

(Hedges et al., 1988; Hedges and Weliky, 1989). To check rhis approximation, we

compared ratios calculated in this manner to the actual ratios calculated for sediments in

which we were able to quantify the ketones. There was generally good agreement (within

4%) but for several samples, including the surfaces of cores 4, 6, 8, 72 and 15, the

approximations underestimate S/V (>i0%). C was set equivalent to p-Cd alone because

Fd yields were frequently below detection.

Downcore profiles for S/V, C/V, ([Ad/Al]v) and ([Ad/Al]s) are shown in Figure

2roPb 3-4 using the data to estimate time of deposition (Table 3-1; Figure 3-2). The cores

from south Hudson Bay, especially 6 and 9, and the noftheast shelf (core 4) all show

decreasing S/V and C/V ratios in recent sediments (Figure 3-44). In core 9, which has an

intrinsic time resolution of about 6 years (Table 3-l), the largest decrease seems to have

occured at about the 1960s. The decrease seems more gradual in cores 4 and 6 but may be an artefact of their time resolution being 50-60 years. Cores in the central basin (e.g.,

10) also liad their lowest S/V ratios in the surface or near-surface sediments

(conesponding to deposition over the last-27 years). Contrasting S/V and C/V profiles occur in cores from the west part of the basin (11,12, and 13; Figure 3-4A), with

98 increasing (rather than decreasing) S/V and C/V ratios in recent sediments, especially post- 1980.

[Ad/Al]v and/or [Ad/Al]s ratios have decreased fairly abruptly in several cores (4,

6, 9) during the last 20 years or so (Figure 3-48). In core 10, a large decrease in [Ad/Al]v occurred in the 1930s, followed by an increase post-1980. In core 7, [Ad/Al]v ratios gradually increased between the 1920s and 1960s. In cores 71,72, and 13, [Ad/Al]v and

[Ad/Al]s ratios generally decreased from about 1940 onwards, and in core 11 (like cores

4 and 9), the [Ad/Al]s ratio has decreased fairly abruptly post-1980 (Figure 3-48).

Spatial patterns in the compositional ratios generally paralleled total lignin yields, with higher ratios occurring as total yields decreased (R< -0.49, p<0.0001). S/V and C/V ratios increased significantly from north to south (R: 0.79 and 0.69), respectively; Figure

3-54). S/V ratios were lowest (minimum 0.30) in cores from south Hudson Bay and highest (maximum 2.3) in cores from the northwest (cores 72 and 13) and central basins

(cores 10 and 14) (Figure 3-6). C/V ratios similarly were lowest (minimum 0.09) in cores from south Hudson Bay and highest (maximum 1.1) in the northwest (especially core l2).

99 2000 1980

1S40

11 19201 2.O 0.2 0.4 il! 1.0 1.5 2.0 0. 0 0.5 1.0 2000r 1'i\ 10 1 880+- 0.0 0.5 0.5 1.0 1.5 2.0 2000 2000r 13 15

1 980

1920

1880 1920{- 0.0 0.5 1.0 1.5 2.0 2.5 ---r.---+1860#J1.0 1^5 2.0 2.5 0.0 0.5 1.0 1.5 2.o 2.5 Raiio Figure 3-4. Profiles of lignin compositional ratios in the sediment cores, including ratios of syringyl, cinnamyl, and 3,5-Bd to vanillyl phenols (S/v, c/v, and 3,5-Bd/v, respectively) (A) and acid to aldehyde ratios for vanillyl ([Ad/Al]v) and syringyl ([Ad/Al]s) phenols (B). Dates of sediment deposition estimated from 210Pb data are shown on the y-axis, however the intrinsic time resolutions of the cores varied from 6-57 years (shown as vertical lines in each plot).

100 1 980

't94

1 1900

1 900 1 1gg0J- -{1 0.6 0.8 1.0 0.0 0.2 0.4 0.6 0.8 1.0 0.4 0.6 0.8 1.0 0.2 0.4 0.6 0.8 1.0 1.2 2000 2000 11

1 1980 1 980 1

1 1960 1 960

1 1940 I

1880 1920-{- 0.2 0.4 0.6 0.8 o.2 0.4 0.6 0.8 2000 2000 13 14 15 1980 1980 '1980 1960 ì 1960 1940 1940

1920 1 920 1 1 "l 1900 1880 1880

1860 -.r 1860+.---_-l..1L--.| 0,4 0.6 0.8 0.0 0.2 0.4 0,6 0_8 1.0 0.0 0.2 0.4 0.6 0.8 1.0 Ratio Figure 3-4. (continued)

101 Ratios of 3,5-Bd/V generally followed CA/ quite closely (R:0.96), with lowest values

(minimum 0.07) along the southwest inner shelf (core 9) and the highest (maximum 1.62)

on the northwest shelf (core i2) (Figure 3-7). Core 15 in Hudson Strait was an exception,

with high ratios of 3,5-Bd/V but low ratios of C/V.

Similar to the other ratios, [Ad/Al]v and [Ad/Al]s were generally lowest in core 9

(<0.65 and <0.40, respectively) and highest in cores 12 or 13 (maximum 0.88 and 0.73,

respectively; Figure 3-8). Low [Ad/41]v values also occurred in a few samples from

cores 10 and14 and low [Ad/Al]s values occuned in a few samples from southeast and

northeast Hudson Bay. [Ad/Al]s ratios showed a relatively strong correlation with

distance from shore (R: 0.85; Figure 3-5D).

2^5 1.2 3^0 1.0

2.4 1.0 2.5 TrB ò TY 0.8 0.8 2 1.5 å¿ ã +{ I 0.6 u) + 0.6 1, 1.0 Ao 0 'ü i= i 0.4 %r fi o.4 1 { .1 0. t.T 0.e 0.5 4"2

o_0 0.0 0.0 0.0 0.9 C D 0,8 0. 0.7 0.7 I * {=TI o 0;6 0.6 T (õ ¡. ïTI À É 0.5 { 0.5 I 0,4 +l Tð1 0.3 E rAd/Arj;l r teor¡uv {;{+ff I I 0.2 55 56 57 58 53 60 61 62 0 100 200 300 400 500 Latitude ('N) Distance from shore (km)

Figure 3-5. Ratios (mean ÈSD) of syringyl and cinnamyl phenols to vanillyl phenols (S/V and CA/, respectively) vs. latitude (A) and distance from shore (B) and ratios of acidic to aldehydic syringyl phenols ([Ad/Al]s) plotted against the same x-variables (C and D).

102 &*n^v<Þ¡øûoCORE 4 67 B 9101112'131415

2.O

1.5

1.0

0.5

0.6 0,s 'l,2 Woody gymnosperm dv Figure 3-6. Plot of CA/ vs. S/V ratios. Core numbers identifu the surface sample of each core. Filled boxes show ranges for major vascular plant types (Goni et al., 1998; Hedges and Mann ,19791' Hu et al., 1999). Open box shows range for suspended particulate organic matter (POM) collected from l3 northern rivers, including the Mackenzie River (Goni et al., 2000) and 12 Russian rivers (Lobbes et al.,2000).

@*¡^Y<Þrø*ôCORE 4 6 7 8 9 101112131415 1.8 ) 1.5 " " -"--"-""'l--"'-' "" t:l ÔI 1.2 d 15o IE cofìo rf) ç"t

r";ffi%off$fuo Surface soil g gtvvY '-"v'"-"-"-"r'- "-"-" -"------ìÞrãñiirebris t 0.0 0.3 0.6 0.9 1.2 C/V

Figure 3-7. Plot of C/V vs. 3,5-Bd/V ratios. Core numbers identify the surface sample of each core. Dashed boxes show approximate ranges for plant debris and peat (Goni et al., 2000; Louchouarn,l99T),, surface soils (Louchouarn, 1997; Prahl et al., 1994), and subsurface soils (Houel et al., 2006; Ugolini et al., 19Sl).

103 @*ft^Y{ÞI9ù,ÔCORE 4 6 7 8 9 10 1112131415

tt, É s,

o'2 o'4 o'8 1'o roJ,lu" Figure 3-8. Plot of [Ad/Ad]v vs. syringyl [Ad/At]s ratios. Filled box shows typical range for fresh plant tissues (Hedges and Mann r 1979; Hedges et al., 1982; Goni et al., 1993). Sd yields were below detection in samples shown on the x-axis.

DISCUSSION

The lignin and 3,5-Bd data reveal clear spatial and temporal variations in the quantity and composition of terrestrial organic material deposited over the last century or so in the sediments of Hudson Bay. Here we relate these variations to the distribution of major sources of terrigenous POM to the Bay, the character and magnitude of these sources and the marine transport and deposition processes that ultimately distribute the material within the Bay.

Spatial patterns of sediment and lignin accumulation

The very low sedimentationrates (ca. 0.04 g cm-z a'1: Figure 3-2;Table 3-1) in the central part of Hudson Bay and variable rates (0.04-0.25 g cm-z a-t¡ on the shelves agree with regional sedimentation maps developed from seismic data (Figure 3-9). Based

t04 on the current surveys, only about 4Yo of the total area in Hudson Bay shows significant

postglacial sediment accumulation, with wide areas of exposed glaciomarine sediments

and glacial till dominating most of the seafloor (Fig. 3-9), especially in the central part of

the Bay (Henderson, 1989; Josenhans et al., 1988).

62"N

60'N

58'N

94'W So"W B6'W 82"W 78.W

l---l con'n,ous posrgraciar n 3f.ï;:'"q:'ifn1'Iffi:', | | Marine Sodiments | | Overlyìng Glaciomarine Sediments, Bedrock or Till ;;d,fi: il sm:*Graciomarine N r'no"aGlacialÏll Exposed Bedrock

2l0Pb-derived Figure 3-9. sediment accumulation rates (black bars: g cm-2 a-l) and modern lignin accumulation rates (white bars: mg --t u-t), shown as an overlay on a preliminary sediment classification map for the Bay, based on interpretation of Huntec D.T.S. and 3.5kHz seismic data (Josenhans et al., 1988).

105 Our cores were generally sited in areas where the maps show at least pockets of

2tOPb sediment accumulation (Figure 3-9) and the majority of the profiles display log-

2rOPb linear profiles (Figure 3-2). There is a possibility that the prof,rles of some of our cores reflect surface biomixing (<10 cm) rather than sedimentation. This may be the case

2l0Pb; for cores 7 , 14, and i 5, which have low activities and shallow vertical profiles for in these cases we consider the profiles to provide only maximum sedimentation rates.

The high sedimentation rates in cores 9 and i 1, which were sited near the Winisk and Churchill Rivers, respectively (Figure 3-7), are consistent with current thinking that modern fluvial sediments are deposited primarily in nearshore areas near major rivers

(Henderson,1989; Josenhans et al., 1988). About 30% of the Bay's total yearly river inflow is discharged into the region between Churchill and James Bay (Dery et al., 2005), and Hudson Bay Lowland rivers like the Winisk also transpoft greater loads of suspended solids compared to the rivers around the rest of the Bay (Henderson, 1989;Newbury et

al., 1984; Stichling, I974). We suspect that the southwest (Lowland) rivers supply greater quantities of both fresh and preserved plant biomass, in addition to sediment, considering the boreal forest and extensive peatlands within their watersheds, compared to tundra vegetation and mineral soils further north (Stewart and Lockhart, 2005). The much greater lignin accumulation rate in the southwest (core 9, 10.8 mg --t a ': Table 3-3), compared to the west (core II,2.7 mg m-2 a-r¡, despite fairly similar sediment accumulation rates (0.25 and 0.20 gcm-2 a-l¡ supports this hypothesis. The disproportionate lignin accumulation in the southwest probably relates to the presence of woody debris and other plant fragments, which contain higher levels of lignin than non- woody and mineral-associated organic matter forms (Goni and Montgomery, 2000;

106 Hedges et al., 1986) and can comprise part of the heterogeneous organic matter load

delivered by rivers, especially large rivers (Onstad et aL,2000). The presence of plant

debris in occasional layers of core 9 would explain the rather high lignin yields from

some of the sediment sections (e.g., 1.46 mg/l00 mg OC at 0-1 cm: Figure 3-3). These

yields equal or exceed yields from northern river materials (Goni et al., 2000; Lobbes et

aI.,2000).

The prevalence of lignin-rich plant debris across the southwest shelf probably

relates partly to enhanced riverine supply and partly to enhanced coastal erosion and

cross-shelf transport in this part of the Bay. In many systems, coarse debris tends to be

deposited very close to shore (Bianchi et aL,2002; de Haas et al., 2002; Gordon and

Goni, 2003;2004) and this seems also to be the case for much of Hudson Bay. For

example, the coarse-grained material supplied to James Bay via its large rivers (e.g., La

Grande, Eastmain) is beiieved to be retained near the river mouths and not transpofted

into southeastern Hudson Bay (d'Anglejan,1982; Kranck and Ruffman,1982; Leslie,

1963). Coarse materials delivered by large southeast Hudson Bay rivers (e.g., Great

Whale River) appear to be trapped in localized depressions near the coast (Zevenhuizen

et al., 1994). Thus, although coastal materials and tidal flats are constantly subjected to

erosion and resuspension by waves, storm surges and tidal currents, supported in the long

term by isostatic rebound, predominantly fine-grained components of coastal deposits

(e.g., clays from the post-glacial Tyrrell Sea) undergo transport offshore (Zevenhuizen et

aL.,1994). Several factors may enhance resuspension and lateral transport of coarse-

grained materials, including plant debris, along the wide, shallow SW shelf (only 35 m deep about 30 km from shore): This shelf contains few localized depressions to trap

t07 material, and is subject to relatively strong tidal currents (Prinsenberg and Freeman,

1986) and wind-induced currents due to its long northwesterly fetch (Henderson, 1989).

This coast may also have alonger ice-free season b.ruur. it does not (always) sustain a

landfast ice cover throughout the winter. Sea-ice rafting may also contribute more

particulate transport along the southwest coast than it does in other parts of the Bay when

sediment-laden ice is blown ofßhore by the prevailing westerly winds (Maxwell, 1986).

In eastem Hudson Bay, although substantial volumes of sediment are incorporated into

the ice (average 29%by volume), the bulk of the ice melts in situ and thus does not

contribute to wider redistribution of coastal materials (Zevenhuizen et al., 1994).

Spatial patterns of lignin compounds

Two major trends emerge from the lignin compositions in Hudson Bay sediments

(Figure 3-5, 6,7,8): a regional variation, expressed as a latitudinal trend, and a gradient with distance from shore. The latitudinal variation in organic composition consists of relative enrichment of syringyl (Sl and Sd) and cinnamyl phenols (p-Cd) in northern sediments, compared to southern sediments, in which vanillyl phenols (Vl and Vd) predominate (Figure 3-5, -6). This latitudinal pattern corresponds directly with differences in the types of vascular plants supplying the terrestrial organic matter in the respective watersheds. 'Ultimate' lignin sources may be identifred by their fingerprints using the ratios of syringyl and cinnamyl phenols to vanillyl phenols (S/V and C/V, respectively) (Goni and Hedges,1992; Hedges and Mann, 1979;Huet al., 1999). S/v ratios reflect the relative abundance of lignin from gymnosperms such as conifers (SfV :

0) vs. angiosperms such as grasses and hardwoods (S/V> 0), whereas C/V ratios reflect the relative abundance of woody tissue (C/V:0) vs. non-woody tissue such as leaves,

108 needles, and bark (C/V>O). Thus, the very low SA/ and CA/ ratios of the southwest inner shelf sediments (core 9) indicate that woody/gymnosperm sources are most important

(Figure 3-6). The very low S/V ratios but higher C/V ratios of the southeast inner shelf sediments (core 6) indicate a mixture of woody and non-woody sources from a predominant gymnosperm origin. At the other extreme, the very high S/V and C/V ratios for the northwest shelf sediments (cores 72 and 13) indicate that non-woody angiosperms are the major, if not exclusive, sources of lignin for that region (Figure 3-6). These differences in lignin composition are consistent with boreal forest (spruce) as the important terrigenous source of POC to the southem Bay and tundra (dwarf birch and willow, sedges, and wildflowers; Stewart and Lockhart, 2005) to the northern Bay.

The intermediate S/V and C/V ratios of the sediments from the northeast, central basins and west Hudson Strait (Figure 3-6) suggest that the major sources to these regions are non-woody angiospefins mixed with some gymnosperm-derived organic matter. The most likely source of gymnosperm organics is redistribution of modern fluvial material from southern Hudson Bay. This direction of dispersal - from the south, northward up the east coast, east into Hudson Strait and also west into the interior of the Bay - is consistent with the general cyclonic (counter-clockwise) circulation pattern of surface water masses within Hudson Bay (Ingram and Prinsenberg,1998; Prinsenberg, 1986b; Saucier et al.,

2004; Wang et al., 1994). However, it opposes the ice transport pathway, which is generally from nofth to south (Markham, 1986; Prinsenberg, 1988; Saucier et al., 2004).

Relatively rapid northward transport of terrestrial organic matter as far as Hudson Strait may occur by means of a coastal current on the eastern side of the Bay (Saucier et al.,

2004).

109 The inshore-offshore gradient in lignin composition is related to marine transport

processes that are involved in redistributing nearshore materials. From the inner shelf to

the outer shelf and the central basins, there is a strong enrichment of syringyl and

ciruramyl phenols relative to vanillyl phenols (i.e. higher S/V and C/V ratios) (Figure 3-5,

-6). We can clearly see these trends in S/V and CA/ by comparing the inner shelf cores in

the southern part of the Bay (cores 9 e, q to those slightly farther offshore in the same

region (cores 7 &,8; Figure 3-6); we presume all of these cores share a common

terrigenous southern source. Further evidence for compositional change related to marine transport is provided by a comparison of the S/V and CA/ ratios of the Hudson Bay

sediments with the properties of suspended materials collected from 13 northem rivers

(Goni et a1.,2000; Lobbes et aL,2000). Only inner shelf sediments from Hudson Bay

(cores 6 and 9) have S/V and C/V ratios within the range occupied by the northern river materials (open box in Figure 3-6), whereas outer shelf and basin sediments, regardless of location, have greater ratios.

Previous studies have identified hydrodynamic sorting, which favours fine- particulate transport, as a means of spatially segregating both lignin yields and signatures

(Bianchi et al., 2002; Gordon and Goni, 2004; Keil et a1.,1994: Prahl, 1985; Tesi et al.,

2007). The segregation arises because lignin components of the dense, coarse plant debris, which is preferentially retained in the coastal zone, have different signatures than those in the fine sediment fractions, which are widely dispersed to more distal locations in the Bay. Our data suggest that fine-grained materials bearing a non-woody angiosperm signature are preferentially dispersed away from the coast to the shelf and interior

110 (Bergamaschi et al., 1997; Goni et al., 1998), while coarse-grained materials bearing a woody gymnospem signature are retained close to the coast.

There are several possible explanations for the association of non-woody angiosperm organic matter with fine sediments. It may derive partly from the watershed source (e.g., clays associated with grassland areas, sands with boreal forest), and/or partly from diagenetic alteration during storage and transport (Benner et al., 1990; Haddad et al., 1992; Hedges et al., 1988). Syringyl and cinnamyl phenols are more vulnerable to physical and chemical breakdown and leaching than their vanillyl counterparts

(Grunewald et al., 2006; Kaiser and Guggenberger, 2000) and thus tend to dominate fine, mineral-associated organic matter (Hemes ef a1.,2001; Keil et al., 1994;1998). The S/V- and C/V-rich signatures of dissolved organic matter (DOM) (Hagedorn et al., 2000;

Houel et a1.,2006; Kaiser and Guggenberger, 2000) also may be incorporated into the signatures of POM through interactions when whitewater rivers (rich in suspended sediments) and blackwater rivers (rich in humic acids) meet (Ertel et a1.,1986) as occurs, for example, where the Nelson (whitewater) and Hayes (blackwater) Rivers meet. In addition to fine grain size, mineral association offers protection from further (complete) degradation, which may aid in wider dispersal and especially persistence through many cycles of deposition and resuspension (Ertel and Hedges, 1984; Filley et aL,2002; Goni et a1.,1993; Nelson et a1.,1995). Indeed, it has been shown that, unlike modern carbon, fossil carbon exhibits a relatively constant loading on mineral surfaces during transport and deposition in marine systems. The most likely source of the non-woody angiosperm signal, therefore, is erosion of geologically older soils or sediments, perhaps glaciolacustrine and glaciomarine materials exposed on bathymetric highs or in nearshore

111 and tidal flat areas undergoing scour. Rivers incised into old deposits onshore, including

marine clays (related to the post-glacial Tyrrell Sea), and fluvial deposits from the early

to mid-Holocene period when emergence was very rapid (>0.1 m a-l; Hillaire-Marcel and

Fairbridge, 1978), present possible sources (Dredge and Nixon, 1992;Lavoie et al.,

2002).

The importance of hydrodynamic sorting for lignin compositional variation in the

Hudson Bay sediments is supported by the increase in 3,5-Bd/V ratios from the irurer

shelf to the outer shelf and interior, parallel to the enrichment of S/V and C/V (Figure 3-

7). 3,5-Bd accumulates during degradation and humification processes in soils

(Christman and Oglesby,lgTl; Houel et al., 2006) and has been widely used to trace the

transport of fine-grained, mineral associated soil organic matter as it disperses through

marine systems (Gordon and Goni, 2004; Louchouarn,7997; Prahl et a1.,7994; Tesi et

aL,2007).3,5-Bd/v ratios are extremely low in fresh plant debris (<0.02; Goni and

Hedges, 1992; Goni and Hedges, 1995; Louchouarn, 1997), reach moderate levels in peat

(0.126; Goni et a1.,2000) and surface soils (<0.4) and very high levels in subsurface or mineral soils (>0.4 to at least 1.5; Houel et a1.,2006; Prahl et al., 1994; Ugolini et al.,

1981). Among the Hudson Bay sediments, the low 3,5-Bd/V ratios (0.07-0.20) in the southwest inner shelf sediments indicate the presence of relatively undegraded plant debris, while the more elevated ratios of the shelf and basin sediments reflect progressively larger contributions from soils, as would be expected from preferential transport and offshore accumulation of fine particles. A complicating factor in using 3,5-

Bd/V ratios as indicators of mineral-associated soil organic matter is that there are marine sources of 3,5-Bd such as kelp (Goni and Hedges, 1995). In the Hudson Bay sediment

TT2 samples (Hudson Strait excluded), the close correlation between 3,5-Bd and p-Cd

(R:0.89) or 3,5-Bd/V and C/V (Figure 3-7) suggests that for most of the sediments, the

source of the 3,5-Bd is tenestrial soil. The very high (>0.a) 3,5-Bd/V ratios of most of

the Hudson Bay sediments (all but cores 6,9 and 7) (Figure 3-7) suggest not only that

soil is an important source but, more specifically, subsurface or mineral soils. Riverbank

erosion and melting of permafrost have both been identified as important sources of

subsurface soils to northern and Arctic rivers, including those entering Hudson Bay

(Adshead, 1983;DredgeandNixon, 1992;Gonietal.,2005;GuoandMacdonald,2006;

Lavoie et al., 2002).

Acid to aldehyde ratios for vanillyl ([AdiAl]v) and syringyl phenols (tAd/Alls)

provide additional indicators of the degree of degradation of terrestrial organic matter,

especially oxidative lignin degradation, which may be extensive in the terrestrial

environment (Filley eta1.,2002; Goni et a1.,1993; Hedges etaL,1988;Nelson et al.,

1995) or well-oxygenated waters (Haddad et al., 1992; opsahl and Benner, 1995). The

low [Ad/Al]v and [Ad/41]s values of core 9 are within the range for fresh plant materials

(Figure 3-8) and thus support our inference that relatively undegraded plant debris is

supplied by rivers to the southwestem part of the Bay and eroded from the coastal zone to be deposited on the inner shelf. All other cores have high (mostly >0.6) [Ad/Al]v ratios indicative of highly degraded materials, and thus consistent with our hypothesis that only the highly degraded fraction of terrigenous organic matter, associated with fine sediments, undergoes transport to other areas. The increase in [Ad/Al]s ratios with distance from shore (Figure 3-5D) suggests some degradation of non-woody angiosperm materials during transport offshore. This implies that there is probably a modern

i13 component to the fine-grained non-woody angiosperm material introduced to the system,

which gets partially degraded, and an older component less subject to degradation.

Although wind-blown pollen can contribute significantly to the POCI",, in more

distal areas of northern marine systems, especially regions without a significant riverine

or coastal source (e.g., (Miltner and Emeis,200l)), none of the lignin compositional

patterns we see in Hudson Bay suggest that this source is imporlant here. The only

exception is the northwest, where the C/V ratios are relatively high. Pollen is

characterized by elevated yields of cinnamyl phenols, which are thought to be important

constituents of pollen-specific polymers such as sporopollenin (Wehlingef aL.,1989), and

thus, samples with high pollen content yield very high C/V ratios (Hu et a1.,7999; Keil et

al., 1998). However, even here the pollen source is likely minor, given the correlation between C/V and 3,5-Bd/V, which is not characteristic of pollen. The pollen expected to be important in this region, primarily from conifers (Pinus or Picea; Bilodeau et al.,

1990), has very low S/V ratios (Hu et al.,1999), which is also inconsistent with the

sediment signatures in the northwest. Previous studies have concluded that aeolian processes are not significant for the transport of inorganic terrigenous material to offshore

Hudson Bay (Henderson, 1989).

Western Hudson Strait is exceptional as the only region where hydrodynamic sorting and transport does not appear to drive biornarker compositions. For example, the unusually elevated 3,5-Bd/V ratios (e.g., relative to C/V ratios; Figure 3-7) suggest significant inputs derived from kelp. The transport of kelp-derived 3,5-Bd would require ice rafting of kelp macerals from Foxe Basin, which is both a region of kelp production and a source of particle-laden ice for Hudson Strait (Campbell and Collin, 1958;

t14 Prinsenberg, i986a). Ice from Foxe Basin disperses into Hudson Strait in July-August

(Prinsenberg, 1986a), where it melts and loses its particulate load.

Downcore patterns of lignin in Hudson Bay sediments

There are clear differences between nofthwest and southern cores in the

compositional shift occuning with depth in sediments, consistent with the distinct

regional sources proposed. In the northwest, the major trend is a decrease in [Ad/Al]

ratios over at least the last 60 years (Figure 3-48), which reflects an increasing relative

impoftance of less degraded plant material. This is accompanied by an increase in lignin

accumulation rates (core l3: Table 3-3). In the south, the major trend is a decrease in S/V

and C/V ratios over the last 40-50 years (Figure 3-44), which reflects an increasing

importance of woody gymnosperm relative to non-woody angiosperm organic matter.

Similar trends in core 4 on the nofiheast shelf and possibly core 10 in central Hudson Bay

are consistent with these regions being linked (through particle redistribution) to southern

Hudson Bay sources. The compositional change is accompanied by an increase in lignin

accumulation rate in core 9 but a decrease in cores 4 and 6 (Table 3-3). The former trend

is consistent with the higher levels of lignin in woody debris vs. non-woody and mineral-

associated OM (Goni and Montgomery, 2000; Hedges et al., 1986); the latter may result

from an increase in sedimentation rate, effectively diluting the sedimentary lignin

content.

Before assigning lignin trends in the cores to variations in source material with time, we must first evaluate the potential for in situ diagenetic alteration. Sediment

diagenesis elicits composition changes similar to those induced by degradation on land

(i.e., increases in [Ad/Al]v and [Ad/Al]s ratios, generally accompanied by decreases in

115 SA/ and C/V (Benner et al., 1990; Opsahl and Benner, 1995)). The progressively

decreasing [Ad/Al] values with depth in cores 6 and 7 and increasing S/V and C/V ratios

with depth in numerous cores (4, 6,7 ,8, 9, 10) (Figure 3-4) suggest that in situ diagenetic

effects cannot explain the trends. Previous studies have found lignin phenols to be

relatively unmodified by diagenesis within accumulating sediments for periods of

hundreds or even thousands of years (Hedges et al., 1982; Louchouarn, 1997).

Many changes including natural large-scale climatic shifts, geomorphological

changes related to post-glacial isostatic rebound, and recent hydroelectric developments

and accelerated global climate change may contribute to the downcore variability in the

Hudson Bay cores. Major vegetation changes such as shifts in the position of the tree line

Q.,liclrols, 1974), while too slow (100s-1000s of years) to be directly registered in the

cores, mean that there are distinct reservoirs of POC with different angiosperm- and

gymnosperm-derived signatures in the watershed and thus considerable capacity for

variation in source material with time. Mobilization of the different materials may hinge

upon the relative importance of surface vs. subsurface or river bank erosion, which in

turn, depends on river ice conditions and the timing and quantity of runoff. Over the last

40 years, river discharge to Hudson Bay has decreased (ca. 13 o/o) and its annual peak

advanced (by 8 days) and diminished in intensity (by 0.036 km3 day-r; Dery et al., 2005),

probably resulting from both water diversion and changes in permafrost, wetland processes and the hydlological cycle. Partial diversions of rivers like the Churchill and

Koksoak into the Nelson and La Grande Rivière systems, respectively, have resulted in

larger shifts in freshwater discharge to those (west and southeast) regions of the Bay

(Dery etaL.,2005). Thus, the combination of natural processes (vegetation shifts) and

116 recent human-induced change (altered river discharge) has considerable leverage for

altering the provenance of materials transported from land to Hudson Bay. This kind of

change may be registered in the compositional shifts from non-woody angiosperm- to

woody gymnosperm organic matter toward the surface of most southern and northeastern

Hudson Bay cores. Similarly, the decrease in the degradation state of the terrestrial

material (Figure 3-4) recorded in the sediments of northwest Hudson Bay (core 12) could

be a manifestation of increased erosion of surface organic material in the watershed

caused by increased rates of spring warming (Rouse et al., 1997; Seneze et a1.,2002) or

increased overland flow (Lammers et al., 2001). The clear record in core 11 from the

west shelf of an increase in non-woody angiospenn (S/V- and C/V-rich) POC over the

last 20-30 years (Figure 3-4) is interesting in that the partially diverted Churchill River is

close and upstream. Within the intrinsic time resolution of the record (about 13 years;

Table 3-1), the onset of the compositional shift roughly matches the onset of low flows in the Churchill River (ca.1977; Prinsenberg,1994). The increasingly nofthern signature of

core 11 could derive fi'om a reduction in the organic material supplied by the upper (more

southerly) portions of the river's watershed.

There is clearly also capacity for major compositional change within the Hudson

Bay marine system itself. The variation in S/V ratio within core 9 alone is as large as the variation in S/V ratio among the other cores of southern Hudson Bay or the variation in core 4; indeed, the S/V ratio in the deepest section of core 9 is similar to the values in the surficial sections of cores 4 and 7 (Figure 3-6). This suggests a compositional change propagating slowly fiom coastal reservoirs to central basin sinks, bringing about a kind of time-lagged compositional shift. Considering that the angiosperm-gymnosperm balance

I17 in the marine system seems to be controlled by redistribution of coarse plant debris from

its initial site of deposition in the coastal zone, a likely explanation is change in the

transport pathways and hence sorting of the terrestrial organic carbon within the marine

environment. A more rapid transport from the coast to the outer shelf is also a plausible

mechanism for the trend toward less-degraded lignin towards the surface of core 12. The

altered patterns of coastal erosion may relate to the longer ice-free season in the Bay

(Gough and Wolfe, 2001; Wu et al., 2005).

The high degree of spatial and temporal coherence in the sediment core records

suggests that large-scale and long-term processes underlie many of the observed trends.

The coherency of the data set instils confìdence that lignins may be used as a foundation to infer the types of terrestrial organic material reaching the Bay, where it comes from,

and how it is transported and deposited within the Bay itself. The downcore records strongly imply change in these processes during the past century or so and larger changes in more recent decades. However, these cores provide a sparse sample density considering spatial variations in the Bay's circulation (Saucier et al., 2004) and sedimentation patterns (Josenhans et al., 1988). Temporal reconstructions based on POM compositions alone are complicated by changes in particulate transport that lag isostatic rebound (Martini, 1986), that are responding to massive hydroelectric development, and are presently undergoing climate change.

SUMMARY AND IMPLICATIONS

Patterns of lignin in eleven box cores collected from the Hudson Bay region reveal clear spatial and temporal patterns in the quantity and composition of terrestrial organic matter deposited over approximately the last century in the sediments of Hudson

it8 Bay. These patterns reflect the sources of the terrigenous organic carbon and its transport

within the Bay. Sedimentary terrigenous organic matter in the southern part of the Bay is

derived from a mixture of angiosperm and gymnosperm, woody and non-woody tissue

sources, which reflect the boreal forest and wetland/prairie vegetation within the

watersheds. Possible specific sources include a mixture of peat, organic surface soil, deep

mineral soil eroded from river banks, and marine clays or old fluvial deposits associated

with the post-glacial Tynell Sea. In contrast, sediments from the northern regions of the

Bay contain terrigenous organic matter derived almost exclusively from angiospenn,

non-woody plants typical of tundra vegetation. A small portion of the southern-derived

material is redistributed to the northeast by the Bay's general cyclonic circulation

(Saucier et a1.,2005) and also into west Hudson Strait and the central basins of Hudson

Bay. However, there is generally very liule modern deposition in the central basins and

the materials here may be quite old and represent organic mafter inputs during the last

glacial period.

In addition to the regional watershed (vegetation) control oì tgnin composition, hydrodynamic sorting of terrestrial POC in the coastal zone is an important control on the quantity and composition of terrestrial organic matter distributed throughout the Bay. The sorting results in preferential dispersal of f,rne-grained sediments enriched in non-woody angiosperm organic carbon and, in general, retention of coarse-grained sediments and gymnosperm-derived POC in coastal zones close to their river sources. The shallow southwest shelf appears to be an important exception, with coarse fluvial materials apparently transpofted at least 30 km offshore. Enhanced supply from Hudson Bay

Lowland rivers combined with more pronounced coastal erosion has resulted in the

119 southern shelf being an important area for sediment and lignin accumulation. Using the

lignin yields in the cores, we estimate the proportion of terrigenous (vs. marine) organic

matter as 40Yo-50o/o or higher in the southwest inner shelf sediments, compared to 15Yo-

35%o in the southeast part of the Bay,I\Yo-I5o/o inthe northwest, and

basins. The proportions along the southwest shelf are similar to those on river-dominated

Arctic shelves like the Beaufort Shelf (Goni et a1.,2000).

Because mixing of organic matter from northern and southem sources and the

preferential transport of fine-grained materials, enriched in cinnamyl- and syringyl-

enriched components, may both contribute to the intermediate lignin signatures we find

in the northeast and interior of the Bay and west Hudson Strait, it is diffrcult to

distinguish the relative importance played by each of these processes. Another complication is the number of possible sources of fine-grained sediments with a non- woody angiosperm signature because of the marine clays in the watershed and coastal zone. Nevertheless, the lignin distributions underline the importance of heterogeneous terrestrial organic matter sources and hydrodynamic sorting during transport in Hudson

Bay. Altered patterns of coastal erosion, possibly mediated by change in ice climate, may underlie angiosperm-gymnosperm compositional shifts recorded in most of the southern

Hudson Bay cores and shifts in the degree of degradation of depositing terrigenous materials in the northwest. On the western shelf, compositional shifts may record changes in the organic transport by the Churchill River as a result of upstream diversion into the

Nelson River. The avaiiable evidence suggests that ice does not provide an important direct means of transport of terrigenous organic material, with the exception of Hudson

120 Strait, where ice exported from Foxe Basin caries kelp markers, but ice likely mediates change indirectly through control of open water and interaction with hydrodynamics.

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130 Chapter 4: Toward a Sediment and Organic Carbon Budget for

Hudson Bay

ABSTRACT

A preliminary sediment and organic carbon budget is constructed for the Hudson

Bay marine system based on published literature and new data collected from nine rivers

and 13 sediment boxcores. The budget considers the main inputs of terrestrial and marine

components of sediment and organic carbon and the main sinks (sediment burial) and

losses (oxidation, export to Hudson Strait). Sedimentation rates (0.032-0.23 g cm-' a-'¡

2lOPb were estimated from profiles and a sediment advection-diffusion model in our cores

and in two cores collected previously by Lockhart et al. (1998). Sedimentation rates

2lOPb l37Cs implied by the profiles were verified against profiles of and, for some of the

cores, contaminant Pb. The sedirnent sink in Hudson Bay was estimated at 138 (+64) x

10ó t a-lby applying the average sedimentation rate to the area of active sedimentation in the Bay (-15% of the seafloor), calculated according to a regional sedimentation map

developed from seismic data (Josenhans et al., 1988). The known sediment sources to

Hudson Bay include river input (1 1.6 (+6) x 106 t a-r), subaerial coastal erosion (8 (+a) x

106 t a-r), marine primary production (24.2 (tl7) x 106 t a-r), and atmospheric deposition

(0.5 x 106 t a-r). Based on the available data, the modern sediment input (44.3 (+1 8) x 106 t a-l) represents only about one-third of the apparent sediment sink in Hudson Bay.

Resuspension and lateral transpolt of shallow-water deposits (primarily glacigenic) and winnowing from topographic highs, likely consequences of the exceptional, continuous relative sea-level (RSL) fall in Hudson Bay, provides a plausible source for the -116 x

131 106 t a-' imbalance. Using organic carbon and ðl3C values in the cores, we estimate the

burial of particulate organic carbon (POC) in sediments at 1.3 (+0.6) x 106 t C ar, of

which - 80% is marine. The importance of reworked glaciomarine carbon to the burial

flux is difficult to estimate because modern (autochthonous) primary production provides

alarge source of POC to the system (16.1 (+7) x 106 t C a-') and its degradation is not

well-constrained. The 0.58 x 106 t C alterrestrial carbon, estimated to be transported as

part of the redistributed sediment load, accounts for almost one-half the total terrigenous

POC inputs to the Bay. If sediment and OC supply and burial in the Hudson Bay system

are presently responding to the history of isostatic rebound, as proposed here, the task of predicting and measuring the additional consequences of river diversions and climate

change (e.g., sea ice, permafrost, runoff) is rendered considerably more difficult.

INTRODUCTION

Sediment and organic carbon budgets provide ways of describing the overall functioning of a marine system by ranking the relative importance of diverse processes that control transfers between land and sea and determining the fate of the materials in the marine environment (de Haas et a1.,2002; Durrieu de Madron et al., 2000;

Johannessen et al., 2003). Budgets also provide a basis for comparing contemporary conditions to previous conditions and for identi$ing the aspects of sedimentation and carbon cycling most likely to be affected by future change (e.g., Gebhardt et a1.,2005;

Stein and Fahl,2004a). Constructing carbon budgets and quantiffing sedimentary carbon sinks is important for global shelves in general (Liu et al., 2000) and especially important for the Arctic's coastal seas, because they are a source of major uncertainty in global budgets, they have changed dramatically on geological time scales, and they face

132 anthropogenic change in the future. Budgets of inorganic and organic particulate matter also provide a crucial basis upon which budgets for other substances of interest, such as contaminants, may be built. Recent detailed analyses of sediment and carbon sources

(riverine and atmospheric input, coastal erosion, marine primary production), and losses to metabolism, burial and exporl, along various shelf seas of the Arctic Ocean (Gebhardt et a1.,2005; Stein and Fahl, 2004a;2004b; Stein and Macdonald, 2004b) have demonstrated a considerable variation among shelves. Efforts to constrain sediment and carbon budgets in different shelves and basins (Gobeil et al., 2001b) and to understand the respective fundamental controls are therefore important first steps toward predicting differences in the response of various shelves to environmental change.

Hudson Bay, at the southern margin of the Arctic cryosphere (Figure 4-1), has been identified as a region highly sensitive to climate variability (ACIA, 2005; Gough and Wolfe, 2001), where some ecosystem changes may already have begun to occur

(Stirling et a1.,7999). The system may therefore serve as a sentinel for changes to come in other Arctic areas, provided key functional similarities and differences are understood.

Like the Arctic Ocean, Hudson Bay receives a large volume of river run-off (-713 km3 a-l; from a massive drainage basin (3.1 x 10ó km2, Prinsenberg, 1986) and therefore is subject to both land-based and maline-based climate effects. It is also semi-enclosed by land, like the Arctic Ocean, which means its oceanography is strongly influenced by local factors, such as river run-off, precipitation, sea ice formation and melting, each of which is changing rapidly. However, Hudson Bay differs from the Arctic Ocean in ways that are likely to be important for sedimentation and carbon cycling including, for example, its shallowness throughout. In particular, a continuous history of isostatic rebound (relative

133 sea level fall) in Hudson Bay opposes the relative rising sea level characteristic of other

Arctic shelf seas (Forbes et al., 1995), implying a fundamental difference in the long-term control of the sediment sinks.

Recent multi-component, broad-scale, integrated studies (e.g., MERICA, études des MERs Intérieures du Canada, and ArcticNet programs) provide the first opportunity to construct a sediment and organic carbon budget for the Hudson Bay system. Here, we combine newly-acquired data for sedimentation based on thirteen widely-distributed, carefully dated sediment cores, and new data for rivers (organic carbon and particulate inputs) with previous data to construct a contemporary sediment and organic carbon budget. The preliminary budget presented here allows a first assessment of the relative importance of various sources and sinks and whether estimated sources (river and atmospheric input, coastal erosion and marine primary production) are in balance with the estimated sinks (sedimentation and advection). We consider the implications of the results for sensitivities to future change and we also highlight critical unceftainties and gaps in our understanding of the system to help focus future research.

OVERVIEW OF THE HUDSON BAY SYSTEM

Hudson Bay encompasses an area of about 841,000 km2 in central Canada

(Prinsenberg, 1986). It has an average depth of only 125 m and subdued seabed relief with slopes genelally less than 2 degrees (Josenhans et al., 1 988) (Figure 4-1, -2). A broad pseudo-shelf out to a depth of about 80 m extends anywhere from 20 to 100 km off the coast. A shallow, northeasterly{rending bank parlitions the shelf regions in the southern part of the Bay, and channels at the northem boundary of the Bay connect the system to Hudson Strait and Foxe Basin in the north. James Bay, about 150 lan wide by

134 400 km long, is an important sub-region which receives much of the Bay's river inflow

(ca.45%).

.13

Figure 4-1. Map of Hudson Bay showing major rivers sampled in 2005 and 13 sediment box-core sites (filled circles labelled with core number).

135 ,) Area (km

200x1 03 400x1 03

g -2oo Eo- q) o

0.00 0,05 0.10 0.15 0.20 0.25 0.30

Sedímentation rate 19 cm-2 a-1)

Figure 4-2. Hypsometry of Hudson Bay in 50-m depth interuals (bars) and as a curve showing the area (k-') and proportion (7o) of the seafloor above the depth 210Pb-rlerived indicated on the y axis. Horizontal dotted lines and circles show sedimentation rates by depth for the cores from this study (filled circles) and four previously published values (open circles) from d'Anglejan and Biksham (1988) and Lockhart et al. (1998).

The shoreline varies from coastal clifß in the noftheast to low-lying rocky coasts along much of Quebec and the northwest coast, and emergent tidal flats, due to isostatic uplift, in the southwest (Figure 4-1). Hudson Bay was the location of the thick (-4 km)

Laurentide ice sheet during the last (late Wisconsinan) glaciation and underwent deglaciation only about 8000 years ago (Josenhans and Zevenhuizen, 1990). As the ice retreated from the northern part of the Bay, marine waters flowed in through Hudson

136 Strait, forming the Tyrrell Sea, and flooded the isostatically-depressed lowlands adjacent

to present day Hudson Bay (Shilts, 1982). Emergence occurred rapidly (>10 cm.a-r) in

the early and mid-Holocene and continues today at the still relatively rapid rate of about 1

t cm a (Hillaire-Marcel and Fairbridge, 1978; Sella et a1.,2007). Glacial features remain

well-preserved and glacigenic sediments exposed across much of the Hudson Bay

seafloor, implying limited postglacial sedimentation (Josenhans et al., 1988).

Hudson Bay is one of the most southerly areas to experience complete annual sea-

ice cover. During most years, the shore is rimmed with landfast ice and the offshore area

is more than 9/10th'covered by more mobile ice (Prinsenberg, 1988). Ice formation

begins in late October, and coverage is virtually complete by December. Maximum ice

thickness and extent tend to be reached by April, melt progresses through May, June and

July, and open water conditions typically return by the first week of August (Markham,

I 986).

The physical oceanography ofthe Bay has been described by Prinsenberg (i986),

Jones and Anderson (1994) and Saucier et al. (1994,2004). Briefly, dense, cold, saline

water enters from Hudson Strait and Foxe Basin primarily in the northwest and at depth,

whereas relatively warm and fresh water exits the Bay in a surface outflow at the northeast corner. The general circulation pattern is cyclonic, with surface currents

averaging 5 cm s-l in summer and.2-3 cm q-l in winter, when the Bay is ice-covered

(Saucier and Dionne, 1998). Tides are semidiurnal and have a range of more than 4 m along the west coast and about 2 m alongthe east coast. Circulation is also influenced by the predominantly northwesterly winds (Prinsenbe rg, 1982).

r37 Freshwater has a profound influence on the physical, chemical and biological

properties of Hudson Bay. Freshwater stabilizes the surface layer, which suppresses

mixing and the upward transport of nutrients (Prinsenberg, 1988). The consequent steep

vertical and horizontal gradients in water properties limit the distribution of the biological

communities (Harvey et al., 7997; Roff and Legendre, 1986). Based on Prinsenberg's

(1977 , 1980, 1988) estimates for freshwater in Hudson Bay, precipitation and evaporation

result in an annual net loss of approximately 23 cm, runoff adds about 85 cm of

freshwater, and ice melt adds 140 cm in summer, while ice production presumably

withdraws this same amount in winter. The average annual river runoff into the system

diminished by about I 1 cm (13%) over the period 1964-2000 (Dery et al., 2005).

METHODS

General Approach

Sediment and organic carbon budgets were constructed from a combination of

previously published data and new data collected primarily during the ArcticNet 0502

expedition (http://www.arcticnet-ulaval.ca) aboard the Canadian Coast Guard Ship

Amundsen, between 15 September and 26 October 2005. Sedimentation rates were

2lOPb determined from profiles in 13 newly-collected sediment cores and two previously-

collected cores (Lockhart et al., 1998) and the total sediment sink estimated using these

rates in conjunction with a published regional sedimentation map (Josenhans et al., 1988;

Geological Survey of Canada Open File Report #2215). Surface fluxes of organic carbon

(OC) to the sediments, OC burial rates and oxidation rates in the sediments, were

estimated fi'om downcore prof,rles of OC content and core sedimentation rates. Marine

and terrestrial carbon components were differentiated using stable carbon isotope ratios

138 (ô"C) and CuO oxidation-derived lignin phenols, with end-member õr3C values constrained by data for suspended particulate matter collected from nine rivers discharging into the Bay. Other sinks (advection, ice rafting) and inputs from marine primary production, river runoff, and coastal erosion were estimated from the literature and from limited additional data for river suspended sediment characteristics collected in

2005. Uncertainty bounds for the various terms (bracketed) and brief descriptions of how these were estimated are provided throughout and taken into account when we discuss the

'Where data and, for example, compare the sizes of sources and sinks. two estimates are combined either by addition or multiplication, we have used propagation of error formulae to derive the new bounds.

Sample Collection

Suspended P articulate Matter

Samples of suspended particulate matter (SPM) were collected from nine Hudson

Bay rivers (Figure 4-7;Table a-l) by pumping water from 0.5-1 m depth and collecting the parliculate phase on stacked pre-combusted GF/D and GF/F glass fibre filters (142 or

293 mm diameter; nominal pore size 2.7 and 0.7 pm, respectively), mounted on stainless steel f,rlter holders. We used a submersible pump connected to the filter apparatus by

Teflon-lined stainless steel hose. Most of the samples were collected during September-

October 2005. Additional samples were collected from the Nelson and Hayes Rivers in

July 2005 and the Great Whale, Little Whale, and Nastapoca Rivers in June 2006 using the same pump and filter set up. Samples were collected from the Churchill River during

April-May 2005 by deploying tluough a hole in the surface of the ice. Filters were stored

r39 frozen (-20'C) and sub-samples (-21 mm diameter) were punched out with a stainless

steel punch and oven-dried at 60"C just prior to analysis.

Table 4-1. Suspended particulate matter properties and DOC concentrations for Hudson Bay rivers River Prov Discharge' Runoff Watershed Month ParticulateMatter DOC' Area km'a-' % mm a-' km2 c/N2 ôr3c mg L-l (o/oo\ Churchill MB 20.s1 3.6 71.2 288880 Oct. t 0.l -28.61 14.1 May 11.4 -21.90 April tt.] -27 .95 April t l.t -29.82 April r 1.3 -29.72 Nelson MB 94.24 16.3 83.1 1125520 Oct. 16.3 -27.65 9.8 July 24.1 -28.15 Hayes MB 18.62 3.2 r 80.7 103000 Oct. 21.1 -28.81 11.2 July 19.5 -29.42 Winisk ON 14.69 2.5 268.s 54110 Oct. 13.6 -28.92 2.4 Great Whale QC 19.71 3.4 451 .6 43200 Oct. 8.8 -26.10 2.3 July < -304 3.04 June lt .4 -28.98 Little'Whale QC 3.14 0.6 3 19.8 I 1700 Sept. 14.1 -21.60 4.2 June 10.3 -29.30 Nastapoca QC 1.86 1.4 629.0 12500 June 10.4 -29.80 Sept. 2.5 Innuksuac QC 3.2s 0.6 290.2 1 1200 Sept. 10.1 de QC I 1.63 2.0 415.4 28000 Sept. 10.0 -26.28 2.3 Povungnituk Various, QC Aug. < _2gs flowing to E James Bay Discharge-wei ghted average6 -28.2 * 0.3 'Discharge, runoff and area from Dery et al. (2005) 2Coruected for presence of inorganic N using positive N-intercept from regression of N vs. OC 'Granskog et al. (2007) oRetamal et al. (2007) sMontgomery et al. (2000)

Sediment Cores

Sedimentcoring sites (stations 3 to 15: Figure 4-1;Table 4-2)were selected along the cruise track from bathymetric and sub-bottom data gathered by an EM300 and 3.5kHz transducer ar.ray (http://www.omg.unb.ca). Sediment cores were collected using a box

t40 corer, which penetrates the seafloor to a maximum of 50 cm. An intact core surface was confirmed visually and overlying water carefully siphoned off with minimal disturbance.

The cores were sectioned immediately, generally into I cm interuals for the top 10 cm,2 cm intervals for the next l0 cm, and 5 cm intervals for the remainder of the core. One core (1 1) was sectioned in2.5 cm intervals between 5 and 15 cm and 5 cm intervals below that. Stainless steel tools were cleaned between samples, and sediment from the outermost 5 cm of the box was discarded. Each sample was homogenized in a pre- cleaned 500 mL l-Chem glass jar with a Teflon-lined lid. After homogenization, subsamples intended for elemental and isotopic analyses were sealed in Whirlpac bags or

20-mL plastic vials. Samples were stored in a freezer (-20'C) on board The Amundsen until the end of the cruise, then shipped to the Freshwater Institute (FV/I), where they were maintained in storage at -20"C.

Particle Size Analysis

Grain size analysis was performed on a near-surface (1-2 or 2-3 cm) section of each core at the Department of Geography, Queen's University, Kingston ON. Samples were pre-treated with repeated applications of 35o/o HzOz at 40"C for two weeks to completely remove organic material. After treatment with peroxide, 1-2 mL of sodium hexametaphosphate (38 g L-') and sodium carbonate (8 g L-') was added to disperse aggregates. Samples were then analyzed three successive times with a Beckman Coulter

L5200 laser diffraction grain size analyser, equipped with a fluid module, under continuous sonication. The third analysis was retained and repolted unaveraged for each sample.

t41 Carbon, Nitrogen and Stable Isotope Analysis

Organic carbon (OC) was measured on acid-decarbonated f,rlter samples at the

University of British Columbia (UBC) using a Carlo Erba NA-l500 Elemental Analyzer

(Verardo et al., 7990) and its stable isotope composition (reported as ôr3C relative to Pee

Dee Belemnite) determined by an in-line isotope ratio mass spectrometer. Total nitrogen

(TN) was determined on separate untreated subsamples. For sediment samples, total

carbon (C) was determined by elemental analyzer and total inorganic c by Co2

coulometer, and then organic C obtained by difference. Replicate samples indicate a

precision of +0.3yo and +1 .60/o for OC and TN, respectively. Atomic carbon to nitrogen

ratios (CntÐ were coffected for the presence of inorganic nitrogen using the positive N-

intercept from a regression of total N vs. organic C (not shown) (Ruttenberg and Goni,

1997; Schubert and Calvert, 2001).

Radioisotope Analysis

Radioisotope analyses were conducted on freeze-dried rehomogenized sediment

samples at the Environmental Radiochemistry Laboratory, Soil Science Department, of the University of Manitoba. The water content of each section was determined after fi'eeze-drying and results are reported on a dry-weight basis, salt-corrected for bottom

t37Cs water conditions at the time of sampling. was counted on a gamma spectrometer using a hyper-pure germanium crystal. The counting effïciency was determined using standard reference materials distributed by the U.S. Environmental Measurements

Laboratory as part of the Quality Assurance Program (QAP50-QAP60, soil) and

t'OPb 2lOPo reference soil samples. was detelmined via its daughter. The sarnples were prepared according to the procedure detailed by Flynn (1968) and counted on an alpha

142 20ePo spectrometer using a silicon surface barrier detector. A tracer, calibrated against a

''oPo NIST standard (Isotope Product Laboratories, #6310) was employed for tt6Ra quantitation. The error in counting replicate samples was less fhanT%o. was

determined using the radon de-emanation method (Mathieu, 1977) and counting

226Ra efficiencies checked using NIST standard QriBS SRM 4953.C). Precision of the

"6Ra measurements based on duplicate or triplicate analyses of 40 samples was l2o/o.

Calculation of Sedimentation Rates

2lOPb Sedimentation rates were estimated from the core profiles of excess activity

by f,rtting the data to output from a steady-state, advective-diffusive model that explicitly

accounts for both sedimentation and mixing (Kadko et al., 1987;Lavelle et al., 1985;

Macdonald et al., 1992):

rxÆ\=-)c ^{-ôAz ôz' ôz'

2l0Pb wlrere C is excess activity (dpm z (cm) the depth in sediments, w" (cm "--'), a-'¡ the

sedimentation velocity, Kb (cm2 a-'¡ the mixing coefficient, and À the decay constant

(0.031 14 a-t). Sediments were considered as having a rapidly mixed surface layer of mixing rate Kb1 overlying a more slowly-mixed deep layer with mixing rate Kb2.

2loPb 2rOPb Supported activity 12lOPbruoo) was estimated by inspection of the activities

226Ra deep in the cores and from activity measurements on several sections from each core (generally surface, middle, bottorn; every third section in core l3; Table 4-2).The

2roPb surface mixed layer (SML) depth was determined by eye from profiles (Figure 4-

3A). Analogous to the 'simple' model (Robbins, 1978), sedimentation velocities were calculated from the slope of the linear regression of the ln (excess 2rOPb activity (dpm ctn-';; in the deep layer on sediment depth (cm). These velocities were then used in the

t43 analytic expressions provided in Lavelle et al. (1985). Kbr and Kb2 values and the

2rOPb incident activity (Co, dpm cm-3) were calculated by least-squares fitting modelled

2rOPb activities to measured activities both in and below the mixed layer (Figure 4-34).

The sediment accumulation rate (r) was calculated from the sedimentation velocity (w,) by r: p,(1-<Þ)y,,, where p, is the density of sediment solids (here2.65 g

2roPb cm-3) and <Þ the average porosity below the mixed layer (Tabl e 4-2). Excess inventories 1X2rOPb*r¡ in each core were calculated by summing the activities Ci (dp* g-t) down to a depth where excess ''oPb was no longer detectable, according to the equation: inventory:XC¡m¡, where m; is the mass-depth increment (g "--') conesponding to the depth interval.

Table 4-2. Sediment core properties and sedimentation parameters t''Pb,uoo Core Water % SML "oRa 1surf. mid, btm or co Kb, depth (n) sand depth (cm) top to bottom) (dpm g-') (dpr g'') (dpm cm-3) (cm-2 a-r) 3 395 0.35,0.44,0.r5 0.90 11.2 5 4 153 63 )) )o )7 2.56 9.0 0.4

5 1t2 4 t0 0.89, 0.02, r.05 r.03 r3.8 1.5

6il9 ¿J 1.2,0.43,0.67 0.78 I 1.5 I

7 t06 13 2 nla,0.92, 1.6 1.3 r r 0.6 0.t I r50 05 2.2,1.9,2.0 1.84 10.9 0.4

934 262 n|a,0.57,0.33 0.45 4.3 I I0 200 l3 r2.3, r.5,0.58 5.40 n.4 2 lt 86 2t2 0.63, 0. r6, 0.20 0.33 8.5 I

12 il6 143 0.62,1.3,l.t 1.03 J.J 0.3 13 t45 82 2.0,2.3,2.0,1 .6, 1.2, 0.45, I .55 5.1 0.4 t4 244 03 s ?, u.:, i.1,3.5 9.8 5.1 0.5 I5 430 230 0.28,0.57,0.r4 0.90 3.5 0

HUD-4 60 2 1.2-1.6, n=10 1.3 7.0 5

FOGO-4 t20 2 2.3-2.6, n=6 2.4 8.7 I s40,s583 60,72

144 Table 4-2 (continued) Core ó," r I,,,CS t''uPb*, Pb Comments . -a t. (cm2 a-r) (gcm-a ) (dpm cm-2) (dpm cm-z) date 0.18 0.66 0.17 L53 73.2 r 880- (0. r 4-0.28) (0. r 2-0.2s) I 905 0.051 0.7s 0-034 0.48 t9.5 r 880 (0.04-0.06) (0.03-0.04) 0. 13 0.71 0. r 0 t.49 76.5

(0.r r-0.r5) (0.09-0. 1 2) 0.13 0.67 0.t2 3.57 49.9 r 880- (0.06-0. r 8) (0.06-0. I 6) 1928 2roPb; 0.05 t * 0.60 0.054* 0.72 16.8 low total profile (0.02-0.07) (0.02-0.07) may reflect mixing o r37Cs 0.17 0.70 0.13 3.71 71.6 r 880- deeper than (0.r5-0.r9) (0.12-0.25) 1892 predicted 9 0.t6 0.45 0.23 3.30 28.4 (0.14-0.22) (0.22-0.32) * l0 0.06* 0.74 0.04*+ 0.74 24.9 I 845- sediment textural change (0.04-0.08) (0.03-0.06) r 930 lt 0.15x* 0.45 0.22** 1.49 s0.4 textural change (0.r r-0.28) (0. r 5-0.30) r3tcs 12 0.15++ 0.60 0.I 6* + 0.15 r9.8 only in 1-2 cm; (0.t3-0.19) (0. r 3-0.20) textural change 13 0.073** 0.61 0.06* + 0.92 13.6 textural change (0.06-0.09) (0.06-0. r 0) t26Ra 14 0.052 0.77 0.032 0.53 r 0.0 I 835- high (0.04-0.06) (0.03-0.04) t9t 5 2roPb; 15 0.07* 0.6 t 0.07* 0.06 t2.0 low total profile (0.02-0.08) (0.02-0.08) may reflect mixing I-tuD-4 0.17 0.71 0.14 5.45 42.0 (0.12-0.r9) (0.09-0. r 7) FOGO-4 0.09* 0.71 0.07* l.3 r r 6.0 poor 'ttcs fit; profile (0.03-0. r 2) (0.03-0. r 0) nray refìect mixing s40,s583 0.043. 0.053 rates from d'Anglejan & Biksham (1988) t"'Pb SML= surface nrixed layer; ''"Pb.,,n,,=supported '"'Pb; C,, = incident activity; Kb¡ : upper Iayer nrixing râle (Kb2= 0.01 cnrr a'l for all cores except I 5 (-0 cnr2 a'r)); w. = sedimentation velocity; $^. = average porosity below the nrixed layer; r sediment 2"'Pb 206Pb/207Pb accunrulation rale; 12roPb.. = excess inventory; Pb date = apparent year ofentry ofcontaminant Pb, based on ratio. +should 'ttCs fit for first entry in I 954. Data for I-IUD-4 and FOGO-4 fronl Lockhart er al. ( I 998). be regarded as nraxinlunr values 2r"Pb (lorv total and profiles could be controlled sínrply by sedinrent ntixing processes); xx uncertain due to poor model fit and/or changes in sediment texture wilhin the core

t45 137 .3 Ca (dpm cm )

I o

ll¡

20 ttl

I 75 I Í t .,I ?t !Ð 3â Ð c; ( uÌ {J 0 I I ¡ ..t t

i I

1U J 1Ô

,ri !0 i I I t5 I '"i I

¡ I 1'r ?ci ?o I I I 4D I

I l 20

ù l 6.1 o (

i i

¡ '"j:

i 15 ;

FOGO-4

21OPb Figure 4-3. A) Profiles of the natural log of excess activity in the sediment cores. Points represent measured values while lines represent the model results. Arrow indicates the bottom of the surface mixed layer (SML). Supported 2r0Pb levels, SML depths, and other parameters are listed in Table 4-2.8) Porosity profÏles in the sediment cores. The shaded zones indicate textural changes.

146 t- a a a ta ti t ? a I t1 ¡ 30 2o ? t_: 5 6

0.4 0.5 0-s 0.7 0.E 0.9 0,5 0,6 0.r 0.4 o.5 o-ô o.7 0.8 0.9 0.4 0.5 0.6 0.7 0,8 0,9 û - ì1 0 i---.- u 0 I ir lç; t, |tì I I ia I I 5 ¿t 5 ltl1,. it' 5 'l 0i if rl I ìt t!t a t¡t i t0 i¡ to I il lìi I t it l!í -i la 10 Itì u1 i\ 15 I rt í\r I it {t ia lrl I i¡ I 20 i? m tft I il l5 ttl ol !l t il t{t I iô t\t ì ¡ i ¡ 7: i Ii i 't -- - r ' u! i Ezoo 30 Q 4 0.5 0.6 0.7 0.8 0 I oi 05 ó6 U 0.4 0.5 ú.4 ú.6 ó.6 0.7 ù.8 0.S E '-- o 0 Ë_0 T- .; o I J I I 2 a t 5 I I l j 4 a ¿ 10 J , 6 I t a /j I j t B I j I 10 I I a 12 ¡ 11 13 tl¿

0.3 0.4 0.5 $.0 Ð.7 0.8 o.4 0_5 0.õ o.f 0.0 0,9 0.1 0.5 0.{i o.7 (}.8 3.s an o-a) 06 Ol 08 O.9 !r Porosity ai ¿ !1:! .,ri i ¡¡ii 1 "ila .iit I Ìf 1c ii! J ,.ì ¡ il,i 15.

" ,,*";; o,'" r"r r'), )"u,'Å

147 Calculation of Organic Carbon Surface Fluxes and Burial and Oxidation Rates

Three different OC flux terms (i.e., accumulation on the seabed surface, burial below the reactive layer and in situ oxidation) were calculated from the organic carbon content in cores 5, 6, and 8, which all showed vertical profiles consistent with steady state supply of organic matter (Figure 4-4).We used a first-order kinetic model (cf.

Johannessen et al., 2003) to fit the down-core decrease ino/oOC measured in each of these cores, i.e., OC1n.,¡: OClm=o; * where m: mass depth (g a*-t), r: sediment "(-u'"), accumulation rate (g a-l), and k: oxidation rate constant the prof,rles showed "--t 1a-t;. little evidence of the effect of surface mixing, implying that OM degradation was rapid relative to mixing rates. Therefore, the expecte d %OC at the sediment/water interface was estimated from the OC intercept at mass depth: 0 g cm-2 of the linear regression of ln (g OC g sediment-'¡ on mass-depth. We multiplied this term (OCi,,.,=où by the sediment accumulation rate (Table 4-2) fo determine the flux of OC to the seabed surface (Table 4-

3). The OC burial rate for each core was calculated as the product of the average deep

2l0Pb %OC value, over the porlion of the core for which the model applies, and the core's sediment accumulation rate. The difference between the surface flux and the burial flux was taken as the core's OC oxidation rate. The carbon surface flux, burial rates and oxidation rates for marine and terrestrial OC were calculated for each core in an analogous manner after the fractions of marine and terrestrial carbon were estimated using a simple two end-member mixing model and ôr3C (see below).

148 --.- %oc 0.8 1.0 1.2 1.4 1.8 0.4 0.5 0.6 0.7 0.8 0.8 1.0 f¡'LJ \.r ó ^10 "P Þ oE qÞ o E15 qt Þ E- q o 6-i o a I I oi I o I 1

-25 -24 -23 -22 -21 -20 -25 -24 -23 -22 -21 -20 -25 -24 -23 -22 -21 -20 ---o- ô13c(%")

Figure 4-4. Downcore profiles of organic carbon content (%OC) and õr3C (%o) in cores a) 5, b) 6, and c) 8. The '/"OC profiles in these cores are interpreted as reflecting OC oxidation in the surface sediments.

Table 4-3. Organic carbon burial rates, surface fluxes and oxidation rates Burial Rate Surface Flux Oxidation %oc %oc \tz Rate Buried Oxidized

â,-l yrs x I0-3 g C x10-3gC x l0-3 g C cm"a'_t -t cm-a'-) _t cm-a'_t _l OC total Core 5 1.04 1.70 0.66 6t% 39% 0.036 t9 Core 6 0.7s 1.02 0.21 74% 26% 0.009 81 Core 8 1.24 1.76 0.51 t1% 29% 0.021 33 Avg. + Range 1.0 + 0.5 1.5 + 0.7 0.48 + 0.4 OC mariner Core 5 0.74 (0.22) r.2 (0.3s) 0.42 (0.13) 64% 36% 0.028 25 Core 6 0.s0 (0.r4) 0.72 (0.20) 0.22 (0.06) t0% 30% 0.0r3 55 Core 8 1.2 (0.4r) r.7 (0.ss) 0.45 (0.1s) t4% 26% 0.023 30 Avg. + Range 0.83 + 0.7 1.2* 1.0 0.36 + 0.2 OC terrigenous (by difference) 0.19+0.3 0.30+0.5 0.12*0.2 'Calculated from ô ''C values and simple two end-member mixing model. Bracketed values show uncertainty resulting from 0.3o/oo and 7o/oo variation in terrigenous and marine end-member values.

149 RESULTS AND DISCUSSION 2I0Pb-Derived Sedimentation Rates

The sedimentation rates, which ranged from 0.032 to 0.23 g cm-z a'' çTable 4-2¡, t'0Pb-derived encompass previously published sedimentation rates for two other cores collected in southeast Hudson Bay (0.043-0.053 g cm-z a-r, d'Anglejan and Biksham,

t'OPb 1988). Excess inuetrtories 1x2r0Pb*.:10.0 to 76.5 dpm c--t, Tabl" 4-2) corroborate these sedimentation rates, implying surface fluxes (product of X2l0Pb*, and the decay constant 0.03114 a-'¡ that agree (within30%) with the apparent surface fluxes, calculated as the product of sedimentation velocity (w,) and incident activity (C";Table 4-2).

2lOPb The distribution of and other tracers in marine sediments often reflect a combination of mixing and sediment burial (e.g., Anderson et al., 1988). In Hudson Bay,

ttoPb where glacigenic sediments comprise much of the seafloor (Josenhans et al., 1988), profiles may potentially reflect mixing between modern deposition and underlying glacigenic material. If mixing were not taken into account, such profiles would lead to overestimates of actual sedimentation rates (e.g., Anderson et al., 1988; de Haas et al.,

1997; Durieu de Madron et a1.,2000). As recommended by Smith (2001), we have

2loPb-derived l37Cs validated the sedimentation rates using and, for a number of cores,

r37Cs contaminant Pb (Table 4-2). The depth of penetration observed in each core (Figure

4-3A) was compared with the depth of penetration for '"Cs deposited in 1954 (date of first entry) as predicted by a numerical advective-diffusive model, which was constrained

2l0Pb-derived using the sedimentation parameters (Anderson et al., 1988; Johannessen et

r37Cs al., 2005). The observed penetration could not be reproduced solely through

2rOPb-derived sediment mixing; rather, the penetration data fit well with the

150 sedimentation and mixing rates combined. Two exceptions were cores 12 and 15, in which '"Cs was present in only the 7-2 cm section, suggesting possible loss of surface t3tcs material, and core 8, in which see*s to have penetrated more deeply than can be explained by the advective-diffusive model, possibly as a result of diffusion in pore water. This core (along with other southern Hudson Bay cores 6, 9 and HUD-4) has a

r37Cs large inventory (Table 4-2),perhaps because of river inputs, andl3TCs is readily mobilized when freshwater sediments are transferred to the marine environment

(Oughton et al., 1997).In the second validation process, the contaminant Pb profile as evidenced by206Pbl20tPb data (Gobeil et a1.,2001a) was compared with that predicted by

l37Cs. a numerical advective-diffusive model, in a similar manner to First dates of detectable entry ( I 845- 1 93 0; Table 4-2) were found to be in accord with results for nearby lakes where sediment mixing was not an issue (Outridge et al., 2002).

For a few cores (marked with asterisks in Table 4-2) additional sources of uncertainty may decrease the reliability of the derived parameters. One unceftainty is

226Ra 2lOPb variation in activities and, hence, in the supported activities, particularly in

226Ra cores 10 and 14. The activities in these cores are relatively high and variable (Table

2roPb 4-2) andthe total activities relatively low (<3 fold greater), which makes exact differentiation of supported and unsupported 2lOPb levels more critical. Limited available

226Ra data suggest the high levels are associated with high concentrations of solid-phase manganese (R2:0.72,n:7; Macdonald, unpublished), as observed elsewhere (Kadko et

t'6Ra a|.,1987). Lacking measurements for every core section, we used constant

2lOPb supported ''OPb le,rels based on average "'Ravalues and activities deep in the cores. Despite the problem of assigning a confident value for226Ra, we note that the

r51 derived sedimentation rates agree with the penetration of '"Cs in these cores (Table 4-2).

226Ra Among the remaining cores, variation in activities is low (averaging 30%) and

2rOPb unsystematic and probably an insignificant source of error in view of the high total activities(>4-foldhigherthansupporled2roPbincores 4,J,9,12,13 and 15,>g-fold higher in cores 3,5,6,8, i 1 and HUD-4).

2r0Pb The accuracy of the modeling also depends upon steady state in sedimentation and mixing. Visual inspection of the cores during sectioning suggested consistent sediment textures within most cores, which is supported by smoothly decreasing porosity profiles (Figure 4-38; note that porosity and o/osand content in near- surface layers of the cores (Table 4-2) are closely correlated; R':0.88, n:13). Four cores

(10, 17,72,73) were exceptions, with sediments becoming progressively more coarse and poorly-sorted in their deeper sections (beginning at 12-18 cm; see shaded portions of profiles in Figure 4-34,8). We conclude that parameters calculated from the profiles of these cores are less reliable than those calculated from the profiles of the other cores.

Estimate of the Sediment Sink

The general spatial distribution of the derived sedimentation rates (Figure 4-5) is one of higher rates on the shelves and lower rates in the central parl of the Bay, which agrees with regional sedimentation pattems interpreted from seismic data (Henderson,

1989; Josenhans et al., 1988). A regional sedimentation map (Josenhans et al. (1988) and

Geological Survey of Canada Open File Report #2215) shows postglacial sediments generally restricted to areas near river mouths, scattered pockets on the Bay's pseudo- shelves and localized depressions, while glacial till and glaciomarine sediments comprise most of the Bay's seafloor. Only four cores (3, 4, 8, and 14) were sited in areas of

t52 mapped continuous postglacial sediments (Figure 4-5). Core 5 was sited in an area of primarily bedrock but survey data collected from the Amundsen (http:ll www.omg.unb.ca) confirmed postglacial sediments at the core site. Four other cores (6,

7 ,9 and 15), and the previously-collected core HUD-4, were sited outside the survey boundaries,but Amundsen daTa indicated postglacial sediments at the site of core 6 and likely till at sites 7 and 15 (no data available for 9). Cores 10, 1 I , 72 and 13 were sited in mapped glacial till, as was the previously collected core FOGO-4 (Figure 4-5). We therefore exclude cores 7, 15, 10, 11,72,13, and FOGO-4 from the estimate of the average sedimentation rate for areas currently accumulating sediment in Hudson Bay, and

an average rate of 0.1 1 (+0.05) g u-' fro- the remaining eight cores. To calculate "*-t estimate the sediment sink, we apply the average sedimentation rate to the area of the seafloor presently capturing sediments according to the regional sedimentation map

(Figure 4-5; Geological Survey of Canada Open File Report #2215). We calculate (in a

GIS) the total capture areaat 125,000 km2 or roughly 15% of the total seafloor area. This area includes the mapped areas of continuous postglacial sediment (ca. 81000 km2, unit 5, continuous ponded mud, and 58, continuous postglacial sand, on original map) and an arbitrary 50% of the mapped area of discontinuous postglacial sediment (ca. 23600 km2 , unit 5/3, discontinuous mud over till and 5B/3, winnowed sandy mud/muddy sand over till on original map), and ISYo of the area outside the boundaries of the map and below 25 rn water depth (ca. 20300 km2, percent coverage of accumulation equivalent to inside the mapped area). We assume no accumulation in water depths less than 25 m, because this is generally regarded as the zone of wave action and ice scour (Henderson, 1989;

Hequette et al., 1999; Zevenhuizen et al., 1994). Although the surface area of

153 accumulation in nearshore areas is poorly constrained (Figure 4-5), our assumption that l5% of the unmapped area is accumulating sediments is probably conservative, considering that most of the cores collected outside the map boundaries, including two previous cores from southeast Hudson Bay (S583 and S40: Figure 4-5), indicate current accumulation in that area (d'Anglejan and Biksham, 1988). Postglacial sediment deposits have also been mapped locally near Manitounuk and Nastapoka Sounds along the east coast of the Bay, using high-resolution seismic data (Hill et al., 1999; Lavoie ef al.,

2008). Uncertainty about whether the percent coverage of accumulation in the area mapped as 'discontinuous accumulation" is 50% vs. 25o/o or 75o/o suggests about a 10o/o

(-12000 km2) uncertainty in the total estimated accumulation area. Applying an average sediment accumulation rate of 0.1 1 (+ 0.05) g cm2 a-l to the estimated accumulation area

(125,000 km2 + 12000 km2), yields a total sediment sink of 138 (+ 64) x 106 t a-r (Table

4-4).

Organic Carbon Burial and Oxidation in the Sediments

The organic carbon (OC) in surface sediments in Hudson Bay is generally <0.5yo-

1o/o,withthe exception of some of the deep basins, where the OC can exceed 1.5olo

(Leslie, 1963; Pelletier, 1986). Surface samples (0-1 cm) from the 13 sediment cores had

OC values between 0.40% and 1.55o/o (average 1 .0%).In several of the cores (e.g., 5, 6,

8), the %OC profiles showed highest values near the surface and downward decreases to roughly constant values at depth (Figure 4-4), reflecting OC oxidation in the surface sediments (Goni and Hedges,1995; Johannessen et al., 2003; Louchouarn et al., 1997).

Total OC burial rates in these three cores, calculated from their sedimentation rates

(Table 4-2) and the constant low %OC values deep in the cores (and over the range of the

154 Sediment Discontinuous Patches of Continuous accumulation Postglacial Postglacial Marine Sediment Marine Sediments Overlying Glaciomarine rate (g cm -2 a -1 ¡ Sediments, Bedrock or Till Ftacïion of terrestrial '-.' s::l;:d,Graciomarrn" Exposed Graciarrirr carbon il N /' l-ractron ol / marine carbon Exposed Bedrock

2l0Pb-derived Figure 4-5. sediment accumulation rates (black bars: g cm-2 a-l) and averâge proportions of marine and terrestrial carbon in the sediments derived from ôl3C data, shown as an overlay on a preliminary sediment classification map for the Bay, based on interpretation of Huntec D.T.S. and 3.5kHz seismic data (Josenhans et al., 1988; Geological Survey of Canada Open File Report#2215). Open bars show sedimentation rates for cores (HUD-4, FOGO-4) collected in t992193 (Lockhart et al., 1998) and previously published rates for two cores collected in 1985 (5583, S40; d'Anglejan and Biksham, 1988).

155 2r0Pb models), varied from 0.75 x l0-3 to 1.24 x i0-3 g C cm-2 a-r laverage 1.0 (+ 0.5) x

10-' g C cm-2 a-l: Table 4-3). The rates are more than an order of magnitude lower than

the rates reported for some highly productive temperate areas (e.g., Strait of Georgia,

Johannessen et a1., 2003) but within a factor of two of rates for less productive temperate

areas (e.g., mid-Atlantic continental slope, Alperin et al., 2002). The estimated burial

rates are also similar to those in the Chukchi Sea (0.67-1.0 x l0-3 g C cm-2 a-I, Naidu et

aL,2004). Applying the average burial rate to the area of postglacial sediment

accumulation in Hudson Bay (125,000 + 12000 km2¡ yields a total OC burial flux of 1.27

(+ 0.6) x 106 t C a-r. If the OC were l}}%marine in origin, this capture rate would

represent about 8% of the Bay's annual new primary production.

The surface fluxes of OC estimated from the core profiles (Figure 4-4) varied

from 1.02 x 10-3 to 1.76x 10-3 g C cm-2 a-tlaverage 1.5 (+ 0.7) x 10-'g C cm-2 a'¡, which

suggests OC oxidation rates of 0.27 x 10-3 - 0.66 x 10-3 g C cm-2 a-' laverage 0.48 (+ 0.4)

x 10-3 g C cm-2 a-r: Table 4-3). The rate constants (0.009-0.036 a-t;ttn: 19-81 years) are

similar to values calculated for bulk organic matter and lignin degradation in the St.

Lawrence Estuary (Louchouarn et al., 1997).

To estimate the proporlions of marine and terrestrial carbon (Fn,u,. and F1".,,

respectively) contributing to the surface flux and the oxidation and burial rates, we used

the ôr3C composition of the organic carbon in the cores (Figure 4-4) and a simple two-

compartment mixing model: ôl3C : (F",o,Xðl3cn,o,) -t (1-F'u.X õl3C,"rr)), where ôt'Cn,u,

and ô13C,... refer to the marine terrestrial end-member compositions, respectively. For the terrestrial end-member, we take the average of observed ôl3C values in river suspended

sediment, giving equal weight to the Nelson River and to all other rivers combined

156 (-28.2o/oo (+1.5%o); Table 4-1), which roughly reflects their contributions to the Bay's

total yearly runoff (Dery et al., 2005). Similar terrestrial end-member values have been

used to model terrestrial and marine organic matter distributions in the North Bering-

Chukchi Seas (-27o/oo; Naidu ef a1.,2000), the central Arctic Ocean and Yermak Plateau

(-27.1%o;schubert and calvert, 2001) and the Kara Sea (-2[%o;Gaye et a1.,2007).

Generally high CIN ratios (10 - 24) and low ôr3C values (-29.9%oto -26.3%o) in 17 SpM

samples collected from nine Hudson Bay rivers (Table 4-I) arc consistent with tenestrial

vascular C3 plant sources for these materials (Goni et al., 1997;2lÌS;Ruttenberg and

Goni, 1997). The C/lr{ and ô13C data (Table 4-1) indicate negligible contributions of

(glacio) marine carbon to the particulate carbon load discharged by rivers, which is

surprising in view of Tynell Sea (glaciomarine) deposits providing a major source of

inorganic sediment to the rivers of James Bay and southern Hudson Bay (Adshead, l9B3;

d'Anglejan, 1980; 1982; Dredge andNixon, r99z;Lavoie etal.,200z), and organic matter generally cycling slowly when associated with mineral surfaces (Hedges and

Oades, 1997) or sequestered in peatlands or permafrost (Benne r et a1.,2004; Goni et al.,

2005; Guo et a1.,2004). Nevertheless, similar or even more highly depleted ôr3C values observed in the Great Whale River in southeast Hudson Bay in July-August (less than -

30%o, Retamal et al., 2001) and in natural lakes and reservoirs near eastern James Bay

(less than -29%o, Montgomery et al., 2000) provide further evidence of negligible contributions of (old) marine POC. The small differences in ôl3C values among rivers

(-28-2%o x l.5%oo discharge-weighted average; Table 4-l) may reflect some influence of freshwater phytoplankton or, more likely, variation in the composition (size, age, degradation state) of the terrigenous carbon pool (Guo and Macdo nald,, 2006).

157 Table 4-4. Sediment and organic carbon budget 1x106 t a-r¡ Sediment POC DOC

Total Marine Terrigenous

Inputs

Marine PP 24.2 16.1 l6.l 4.0 +t7 *7 +.7 *3 River input I 1.6 0.46 0.46 3.6 +6 * 0.33 + 0.33 T2 Coastal erosion 8.0 0.072 0.012 +4 + 0.04 + 0.04 Atmospheric input 0.50 0.12 0.12 * 0.01 + 0.08 + 0.08 Combined Inputs 44.3 16.8 16.1 0.6s 7.6 +18 t/ +7 + 0.3 =-) Sinks & losses

Sediment burial 138 1.27 1.03 0.23 +64 * 0.6 + 0.5 + 0.4 Loss by oxidation in 0.60 0.60 0.45 0.15 sediments + 0.5 + 0.5 + 0.3 + 0.2 Loss by ice export 0.00 0.000 0.000 0.000 + 0.35 + 0.003 + 0.001 + 0.002 Loss by advection -0 -0 -0 -0 (s e¡

Apparent oxidation in 21.8 16.0 l5.l 0.85 l< 9) water columnl Combined Sinks t60 17.8 16.6 t.2 s9

Apparent by 116 t.04 0.46 0.58

Assumed 90%o oxidation/dissolution of algal mass for sediment budget (Gaye et a1.,2007; Macdonald et al., 1998). For POC budgets, values represent the unbalanced terms (combined inputs plus POC associated with apparent resuspension load minus combined sinks). 2Sediment term is by difference (sinks minus inputs) to balance sediment budget. POC supply is derived from sediment supply, assuming OC content of 0.\-IYo, of which 44%o is marine and 56%o is terestrial.

To estimate the composition of the marine endmember, we make use of lignin measurements from sediment samples in the upper portions of i 1 cores (Chapter 3) and the linear relationship that exists between lignin yields (À8), defined as the sum of syringyl, vanillyl and cinnamyl phenols (Hedges and Mann, l97g),and ôr3C (f:0.79, p<0.0001, excluding two outliers: Figure 4-6). Because lignin is a major component of

158 terrestrial vascular plants but essentially absent from marine organisms (Goni and

Hedges, 1995; Hedges and Mann, lgTg),the strength of the lignin - ôl3C relationship provides confidence that the ôl3C values are generally applicable to tracing the distribution of terrestrial OC (OCt.,r) in Hudson Bay, at least among modern sediments

(cf. Lavoie ef a1.,2008).

_20 CORE y = -6.36x-20.8 ø4 -21 rL-1.79, p=0.001 X6 +7 ¿B _22 v9 -,¡ l0 -23 r" 11 o -¿o w12 -24 ê13 *14 * 15 -¿c

-26 0 0 0.5 ¡.u 1.5 Lignin Yield (nB) (mg/100 mg OC)

Figure 4-6. Retationship between ô13C values (%o) and sedimentary lignin yields (48), defïned as the sum of vanillyl, syringyl and cinnamyl phenols (Hedges and Mann, 1979), in units of mg/100 mg OC.

The ôl3C value of the marine endmember influencing the sediments may be estimated from the y-intercept (for x:0) of the linear regression (õr3C : -6.36(r\8) -

20.81). The selected value (-20.8o/oo) is similar to the marine endmembers chosen for the

North Bering/ Chukchi Sea (-2lo/oo, Naidu et al., 2000) and the central Arctic Ocean (-

21.3o/oo to -I9.0o/oo, Schubert and Calvert, 2001). Uncertainty in the end-member value of about lo/oo is suggested by the variation in ôl3C at low values of Â8 (Figure 4-6).

The proportions of marine carbon in surface sections of all the cores estimated using these endmembers vary from >80% in central Hudson Bay to 60%-80% in the

159 southeast shelf, and a minimum of 40o/o-50Yo along the southwest inner shelf (Figure 4-

5). In cores 5, 6 and 8, downcore profiles of ôr3C (Figure 4-4)yieldproportions of marine carbon of 78%o-84Yo,56yo-72yo, and72o/o-850lo, respectively. Downcore profiles of yooc^u, generally resembled profiles of (total) o/oOC, with steep decreases near the surface of several cores (not shown). Applying the same modeling approach to these profiles as above yields an average marine carbon surface flux of 1.2 x l0-' g C c^-' u-t

(range + 0.4 gC cm-2 a-r), burial rate of 0.83 (* 0.7) x 10-' g C cm-2 a-r, and oxidation rate of 0.36 (+ 0.2) x 10-3 g C cm-2 a-' iTable 4-3). Applying the average burial rate to the area of postglacial sediment accumulation in Hudson Bay yields a total burial sink for marine

OC of -1.03 + 0.5 x 106 t C a-r and an additional loss to oxidation in the sediments of

0.45 (+ 0.3) x 106 t C ar lTable 4-4).Theterrestrial carbon sink and oxidation loss are estimated (by difference from total OC) at0.23 (+ 0.4) x 106 t C a-r and 0.15 (+ 0.2) x 106 t C a-r (Table 4-4).Thepropagated effors highlight large uncertainties in estimates of the terrestrial carbon sink and oxidative losses.

Other Losses and Sinks

Ice Export

Ice rafting is believed to be an important mechanism for moving sediment locally within Hudson Bay (i.e., on scales of 10s to 100s of meters, Hequette et al., 1999) and to a lesser extent regionally (e.g., between where it forms in the north and melts in the south and also inshore to offshore, Henderson, 1989; Pelletier, 1986). However, we expect minimal impact of this mechanism on the sediment and OC budget for the Bay because most of the ice forms and melts within the Bay (Prinsenberg, 1988). The only area of significant ice exchange is in the Southampton, Coats and Mansel Islands area, where

160 there appears to be significant melt of imported ice (Tan and Strain, 1996), including

sediment-laden ice imported from Foxe Basin (Leslie, 1963; Pelletier, 1986; Prinsenberg,

1986). Nevertheless, net export of sea ice has been estimated at about 35 km3 a-' llVturry

and Barber , 197 4) cited in (Prinsenberg, 1984)). Here, we take as our best estimate a

negligible net exchange but use this export volume and average 'background'

concentrations of ice-borne sediment of about 10 mg L-l as determined elsewhere in the

Arctic (Eicken, 2004), to calculate an import or export flux of ice-rafted sediment of

+0.35 x 106 t a-r. Assuming an organic carbon content of 0.8-1% (like shoreline

materials), the carbon input or loss by ice-rafting is about +0.003 x 106 t C a-r or +0.001

6 and +0.002 x 10 t a-l of marine and terrestrial carbon, respectively, assuming the same

carbon composition (44% marinel56%o tenestrial) as our most nearshore core (core 9).

Advection

Another potential sink for sediment and carbon is advection into Foxe Basin

and/or Hudson Strait. Estimating advective losses requires measurements of water

volume transports as well as particle/carbon loads. Net freshwater transpoft rates out of

Hudson Bay have been calculated at about 0.03 Sv (l Sv: 106 m3 s-r) using freshwater

budgets (Prinsenberg, 1986). However, total imports and exports are not well known and

the water exchanges are complex (Saucier et a1.,7994;200fl. The total transport in and

out of Hudson Bay may be as high as 0.55 Sv (Ingram and Prinsenberg, 1998) with -0.25

Sv below 100 m, implying a replacement time of four years (Saucier et al., 2004).

Although we have no direct information about the parliculate or carbon load of the water being exchanged, a previous study of sedimentary lignin distribution concluded that only a minor amount of terrestrial particulate organic material escaped from the Bay into west

161 Hudson Strait (Chapter 3). Similar inshore and offshore SPM concentrations (0.71 (+

0.22) mgl-r and 0.72 (+ 0.48) mg L-r, respectively) and POC concentrations (0.074 (+

0.024) mg L-r and 0.053 (+ 0.021) mg L-r, respectively) (Anderson and Roff,, 1980) also suggest minimal particulate transport out of the Bay with the coastal current outflow.

Likewise, POC concentrations were consistently 0.2-0.3 mg L-l along the coastal current transport pathway from James Bay northward to the channels to Hudson Strait (Harvey et al.,1997). These greater concentrations may reflect localized distribution of POC in the surface coastal current (Saucier eta1.,2004), but the absence ofa north-south gradient seems inconsistent with northward transport. Clearly, any net exchange would likely be an export. However, based on the available evidence, we assume negligible net exchange for suspended sediment and POC between Hudson Bay and Foxe Basin and Hudson

Strait.

Loss of DOC through advection is difficult to estimate because there are few published DOC measurements for Hudson Bay waters. Coastal waters in western Hudson

Bay had DOC concentrations less than 10 mg L-' (.50 uM, Chapter 2) and surface waters offshore from the Great Whale River estuary in southeast Hudson Bay contained2.47 (+

0.13) mg t-' 1n:S; (Retamal et a\.,2007). Offshore from the Nelson River estuary in southwest Hudson Bay, surface water DOC concentrations decreased from >7 mg L-l to as low as 0.48 L-r lBaker eta1.,1993). Thus, the range of values observed is similar to that in the Arctic Ocean, where coastal watels near rivers contain 5-10 mg L-' (Opsahl et a1.,7999), and surface waters away from major rivers contain 0.4-1.3 mg L-r (Wheeler et aL.,1997). Recent data show that chromophoric (coloured) dissolved organic matter

(CDOM) roughly doubles in concentration between Foxe Channel and northern Hudson

r62 Strait and the Bay itself, including outflowing waters in the northeast (Granskog et al.,

2007). The differences in CDOM were present both in the surface mixed layer and deep

waters; however, the (summer) measurements likely reflect maximum values. Likewise,

dissolved nitrate and phosphate concentrations (J.-E. Tremblay, pers. comm.) and

Redfield stoichiometry (i06C:l6N:lP) imply carbon concentrations of 0.7 and 1.8 mg L-r

(respectively) in the deep waters of the Bay in the nofihwest (i.e., in the inflow) and 1.1

and2.4 mg L-l in the east near the outflow to Hudson Strait, which indicate an apparent

1.4- to 1.5-fold increase as waters circulate within the Bay, assuming OC remains a

constant proportion of the total carbon. This approach neglects the possibility of

oxidation and conversion of OC to dissolved inorganic carbon (DIC). Assuming 0.55 Sv

exchange, a DOC concentration of -1.0 -g L-r in the deep inflowing waters (similar to

waters in the Arctic Ocean) and a concentration of -1.5 mg L-r in the outflowing waters

(consistent with the increase implied by nutrients and CDOM), we estimate a likely maximum DOC export of -9 x 106 t C a-r.

Inputs of Sediment and Organic Carbon

MarÌne Primary Producríon

Hudson Bay has a relatively low standing crop of phytoplankton at the lower end of what has been observed in other Arctic seas (cf. data fiom various Arctic areas compiled in Subba Rao and Platt, 1984). During a bay-wide survey in August-September

7975, surface chlorophyll concentrations averaged only 0.09 rng chl a m-3 offshore and

0.28 mg chl a m-3 inshore (Anderson and Roff, 1980). Similar chlorophyll contents (0.25-

1.25 mgchl a rn-3) were observed during a September 1993 transect up the eastern coast of Hudson Bay (Harvey et al., 1997). Chlorophyll levels between 0.2 and2 mg chl a m-3

t63 in coastal areas including Chesterfield Inlet (Roff et al., 1980), the Churchill River

estuary (Baker et al., 1994) and Manitounuk Sound (Legendre et al., 1981) are consistent

with the data from the large-scale surveys. Springtime chlorophyll measurements in

inshore regions of Hudson Bay indicate a relatively well-developed ice algae community

(Freeman etaL.,1982; Gosselin et al., 1985; Michel etaL.,1993; Welch et al., 1991).

Chlorophyll in the bottom of the ice was as high as 8 mg chl a m-2 in Chesterfield Inlet

(Welch et al., 1991) and 6 mg chl a m-' near Manitounuk Sound (Michel et al., 1993).

Primary production data for Hudson Bay are scarce compared to the biomass data but values of about 2.5-3 mg C m-3 hr-l have been observed in southeastern Hudson Bay

surface waters (Grainger, 1982;Legendre and Simard, 1979). A production estimate of

70-100 g C m-2 ar was derived from these data (Pett and Roff, lgS2) but likely

overestimates the bay-wide average. Roff and Legendre (1986) estimated that inshore productive areas may reach 70 g C m-' a-' but proposed an overall average of 35 g C m-2 a-1, excluding the main spring diatom bloom, which has yet to be assessed directly, and ice algal production, which these authors suggest may contribute l0 g C m-2 a-'. Hudson

Bay probably has a relatively long growing period for ice algae because of favourable spring light conditions (Cota et al., 1991). Considering the evidence for relatively high ice algal production in Hudson Bay, Sakshaug (2004) estimated a total primary production rate of 50-70 g C m-t a-r, which is than g a-' slightly higher the 30 C ^-' typical for seasonally ice-free Arctic areas (Subba Rao and Platt, 1984).

New production, which conesponds to the rate at which POC can be exporled downward from the euphotic zone, was estimated at roughly 50% of the total production values (25-35 g C m-2 a-') by Sakshaug (2004) and a similar value Qa gC--t u-')

t64 derived from carbonate data from the northern part of the Bay (Jones and Anderson,

1994). These estimates generally agree with the estimates for other Arctic shelves (e.g.,

7-17 g C m-2 a-t on the Beaufort Sea shelf, 25-50 g C m-2 a-l in Baffin Bay, Sakshaug,

2004). Here, we assume an average new production of 24 (+ 10) g C m-' a-' over an arca

of 841,000 km2 and assign 80% (+ l0%) of it to the POC budget for a total of 16.1 (+ 7)

x 106 t C a-r lTabl e 4-4). The remaining2)% (4.0 (+ 3) x 106 t C a-r) we assign as an

input of dissolved organic carbon (DOC) following the estimates of Gosselin et al.

(1997). Although many OC budgets neglect DOC (Gebhardt et a1.,2005; Macdonald et

al., 1998), we include it here despite the unceftainty in the values because it may

represent a relatively large fraction (up to 65%o) of the marine carbon produced in Arctic

systems (Gosselin et al., 1997) and hence suppoft alarge and active microbial loop (Rich

et al., 1997).In Hudson Bay, DOC comprises an estimatedg0% of the total organic

carbon discharged by some rivers (Great Whale River, Hudon et al., 1996), is present at

elevated concentrations in surface waters, where the chromophoric component may

influence various thermo- and photoprocesses (Granskog et aL,2007) and appears to

contain a labile component, which supports a heterotrophic food web (Else et al., 2008a;

Else et al., 2008b).

We also include in the budget an estimate of the total particulate input from primary production (algal biomass), because diatoms represent an important component of the phytoplankton community in the Bay (>50% in the nofth and west (Harvey et al.,

1997),20%-60% in the east (Bursa, 1961)) and their remains appear to be preserved in the sediments in at least some areas of the Bay (i.e., sufficient for paleoceanographic studies in Nastapoka River estuary, eastern Hudson Bay, Lavoie et al., 2002).

165 Microscopic examinations revealed abundant diatom remains in the upper sections of

some of our cores (personal communication, Dr. Dave Scott, Dalhousie University,

2008). This is consistent with findings of near-constant biogenic silica fluxes with depth

in a short-term sediment trap study (Lapoussière eta1.,2005). Assuming a contribution of

50% (+25o/o) ftom diatoms to total PP, and a ratio of total biomass to POC of 3:1

(Brzezniski, 1985), we estimate a total parliculate input of 24 .2 (+ 17) x 106 t a-r.

River Input

Regionally, rivers discharging to Hudson Bay from the province of Manitoba

(Figure 4-1) and rivers discharging indirectly to Hudson Bay through the east and west

sides of James Bay account for the majority (70%) of the total yearly runoff into the Bay

(Table 4-5). The Nunavut rivers account for an additional 12o/o, and nofihern Ontario and

Quebec rivers (excluding those flowing into James Bay) 8% and 70Yo, respectively

(Prinsenberg, 1977; Prinsenberg, I 980).

Although freshwater discharges are relatively well known, with observations for

about 35 rivers (80% of the total discharge) compiled in Environment Canada's

Hydrometric Database (HYDAT (Environment Canada,2004); for a summary see Dery

et al., 2005), there is relatively little information about the nature or quantity of parliculate material carried by these rivers (Environment Canada,2004; Hudon et al.,

1996; Stichling, 1974). Maps interpreted from data collected during 196l-1970

(Stichling, 1974) show suspended sediment concentrations of <50 mg L-r in most rivers

draining into the Bay from Nunavut, Manitoba, northern Ontario and Quebec and 50-200 mg L-r in a few exceptional rivers in Manitoba (Churchill River) and Quebec (Great

Whale River). Estimated concentrations were in the 50-200 mg L-r range or even the

166 200-400 mg L-r range for most of the rivers flowing into James Bay. Sediment yields

were estimafed at2-10 tkm-2 a-l for Manitoba rivers draining the Prairies and 20-100 t

kr¡r2 a-r for rivers draining into the eastern side of James Bay (Stichling, lgl4).

More recent suspended sediment data are sparse, but summertime (April-

November) measurements have been made on rivers in Nunavut, Manitoba, and Ontario,

and more numerous measurements on several rivers flowing into the western side of

James Bay (Baker, 1989; Baker et al., 1993; Environment Canada,2004; Jeffries et al.,

1994). Much of the available data (except for that from the Nelson River) is compiled in

Environment Canada's Hydrometric Database (HYDAT; Environment Canada,2004).

The average summefiime suspended sediment concentration is greatest for western James

Bay rivers (51 mg L-r: Table 4-5), followed by Manitoba rivers (37 mgl--r), and lowest

for Nunavut rivers (2.9 mgI--t¡. The standard deviations of the measurements represent

75%oto 125% of the mean values (Table 4-5). Although these averages are generally at the low end of Stichling's (1914) estimates, we consider the western James Bay values to be reliable, because they reflect a relatively large number of samples from stations that represent about 40%o of the drainage area. As annual averages, these summertime values are probably much too high; seasonal studies have shown that SPM concentrations are low for most of the hydrological cycle except for a brief spike during peak discharge in spring (d'Anglejan, 1980; d'Anglejan, 1982; Hudon et al, 7996). Without data reflecting the wide seasonal variations in flow, the concentrations showed at most weak linear associations with discharge çr2 < 0.2¡, which precludes using discharge-concentration relationships derived from these data to estimate annual sediment flux.

167 Table 4-5. Suspended sediment concentrations measured in Hudson Bay drainage areas and estimated regional sediment inputs to the Bay

Region Area Dis- Runoff Ave. Susp. Sed. Conc. Est. Est. Est. charge (Apr-Nov) Winter Sed. Sed. (mg L-') Conc. Flux Yield x 103 km3 mm a-t Mean SD mg L-r x 106 t a-r t km-2 _t km2 a-l (range) A' 2.9 t0 NU 484 81 r80 (0.3-7.0) 2.2 0.2 + 0.12 0.4 JI 55 MB \51 5 180 115 (8.0- r 97) 36.9 5.0 + 4.62 3.2 t1 43 ON 198 56 280 (2.0-62) 14.1 0.8 + 0.62 3.9 West 327 James 5l Bay 372 129 341 (r.0-7r 1) 64.0 10 4.7 + 5.62 2.8 East James Bay 310 r89 610 3.9 +3.23 12.6 QC 161 12 441 0.5 + 0.14 2.9

Total 15.0 + g5 Using 60% of James Bay 11.6 + 6

rivers discharging indirectly to Hudson Bay through James Bay. Most Suspended Sediment Concentrations are from Environment Canada (2004), with additional data for Manitoba (MB) from Baker ( 1989), Baker et al. ( I 994), and G. McCullough (pers. comm.); and for Nunavut (NU) from Jefferies et al. ( 1994) 2Calculated as: (assumed winter concentration x 35% of discharge) + (summer average concentration x 65%o of discharge); propagated error represents standard deviation ofobserved concentrations and22Yo interannual variability in discharge 'No data available from Environment Canada. Estimated lluxes f'or four rivers from East James Bay (d'Anglejan, 1980; 1982) pro-rated to total runofffor East James Bay region; propagated error represents 787o variation among flux estimates based on the four rivers and 22o/oinferannual variability in discharge 4Estimated Great Whale River fluxes (Hudon et al., 1996) pro-rated to total runofTfor Quebec region; propagated error represents 2lolo interannual variability in sediment f'luxes and 22%ovariablliTy in discharge sError propagated from estimated fluxes for each drainage area

Here we estimate fluxes from regional runoff rates and recent suspended sediment

data (Table 4-5) for western James Bay, Ontario, Manitoba and Nunavut. We calculate

summeftime and wintertime fluxes separately, because some of the summertime (April-

Nov) measurements probably represent peak flow (and hence sediment discharge), and

by incorporating assumed (low) winter values, we may avoid overestimating the annual

loads (cf. d'Anglejan and Biksham, 1988; Hudon et al., 1996). To estimate summertime

fluxes, we multiply the average summefiime concentrations by 65% of the annual runoff

168 rate for each area, which is roughly the percentage of runoff discharged during the

summer months (Prinsenberg,l9TT; Prinsenberg, 1980). To estimate wintertime fluxes,

we take the concentrations observed during summertime periods of low flow (Table 4-5)

and multiply them by the remaining 35%o of the annual runoff rates. The eror associated

with these estimates (Table 4-5) reflects propagation of the error in observed average

suspended sediment concentrations (i.e., standard deviations) and 22Yo interannual

variability in runoff rates (Dery et al., 2005). Lending confidence to our estimates is good

agreement between the sediment yield for the Nunavut rivers (0.4 t km-2 a-r: Table 4-5)

and independent estimates for two rivers in the Nunavut region that drain into the Arctic

Archipelago (0.2-0.6 t km-2 a-', AMAP, 1998). Likewise, estimated sediment yields for

other Hudson Bay drainage areas (3-12.8 t km-2 a-'¡ compare well with those for Russian rivers including the Yenisey,Lena, Ob', and Severnaya Dvina (1.9-12 t km-2 a-', Holmes

et al., 2002).

For eastern James Bay rivers, which are not included in HYDAT, we estimate the regional sediment flux using published estimates of annual suspended sediment fluxes for four relatively large rivers (d'Anglejan, 1980; d'Anglejan, 1982). Scaling the combined fluxes of the rivers (2.08 x 106 t a-r) in proportion to their runoff (collectively about 53o/o of the total eastern James Bay outflow), we calculate a total suspended sediment flux for eastern James Bay of 3.9 x 106 t a-r lTable 4-5). The error associated with this value (+3.2 x 106 t a-l¡ represents the variation (75%) among fluxes scaled-up individually fi'om each river and 22o/o interannual variability in discharge (Dery et al., 2005). The estimated flux implies a sediment yield for eastern James Bay (12.6 t km-2 a-r¡ which is very similar to the one derived from the HYDAT data for western James Bay (12.8 t km-2 a-r).

t69 For Quebec rivers discharging directly to Hudson Bay, we base our estimate of

sediment flux on the results of a quantitative year-round assessment (sampling every i -3

weeks) of the Great Whale River, which contributes about 29Yo of thetotal discharge

from the Quebec drainage area (Hudon et al., 1996). This estimate probably provides an

upper limit because the Great Whale River drains a tundra,/boreal forest transitional area

with widespread glaciomarine deposits and glacial till cover, whereas the more northerly

Quebec rivers drain relatively barren Canadian Shield (Stewart and Lockhart,2005).

Hudon et al. (1996) found that SPM in the Great Whale River was as high as 40 mg L-l

for a brief period around ice breakup, fell to 2-10 mg L-l for most of the summer period

and fell further (<2 mgl-t) for the winter. Using a polynomial equation to model these

changes, Hudon et al. (1996) calculated ayearly particulate load (inorganic * organic) of

115,000-156,000 t a-1. Pro-rated to total runoff for the Quebec region, the total sediment

load is about 0.5 (+ 0.1) x 106 t a-r.

combining the above estimates gives a total sediment load of 15.0 (+ g) x 106 t

a-1, with rivers that flow into James Bay representing about 58% of the total, which is

comparable to the 63Yo estimate by Kranck and Ruffman (1982). It remains unknown

how much sediment is trapped locally within James Bay. James Bay (and southwest

Hudson Bay) rivers are deeply incised into old marine (Tynell Sea) deposits that now

surround the Bay; hence these materials comprise much of the river sediment flux

(Adshead, 1983; d'Anglejan, 1980; Dredge and Nixon, 1992). According to the available

data,the coarse-grained portion of the sediment load is deposited nearshore, whereas some of the silts and clays, which make up about 50%-60% of the total sediment load

(d'Anglejan 1982), are flushed seaward and distributed across the seafloor of James Bay

t70 and also transported northward into Hudson Bay (d'Anglejan, 1980; Kranck and

Ruffman, 1982). Concluding that the silt- and clay-sized fraction (maximum 60%) represents an upper limit to the sediment load of James Bay rivers that actually reaches

Hudson Bay, we reduce our estimate of the river sediment input from James Bay rivers to

Hudson Bay to 60%o of thetotal James Bay inputs. The estimated total riverine sediment input to Hudson Bay is thus 11.6 (+ 6) x i06 t a-t (Table 4-4;.

The OC content of SPM from northern rivers generally va¡ies from 0.5% to 5o/o, with lowest proportions characterizingthe material discharged during freshet. On the

Mackenzie River, spring/early summer values avetage 1.4% (+ 0.2%; n:10) and winter values 4.8% e 13%; n:18) (Goni et al., 2000; Yunker et al., 1993). Similarly, a range of

0.4%-2.7% OC was obtained for spring/suTnmer suspended sediment samples from 12

Russian rivers (Lobbes et al., 2000). Measurements of OC content for Hudson Bay rivers are generally within the same range. The OC content of SPM samples collected from the

Nelson and Hayes Rivers in July and October 2005 averaged 2.2% (+ L4% OC; n:7) and

3.8% (+ 1.I% OC; n:7), respectively (G. McCullough, pers. conìm.). Earlier samples from the Nelson River showed values as low 0.4%o in spring and as high as 6.4% (L I.3%

OC; n:l8) in summerlfall (Baker, 1989; Baker et a1., 1993). Hudon ef al. (1996) estimated 4-I5% OC content in the suspended solids load from the Great Whale River, with lowest proportions during spring freshet (May-June). Data from James Bay rivers are relatively scarce, but the Tyrrell Sea deposits that provide the major source of sediments to many of those rivers (Adshead, 1983; d'Anglejan, 1980; Dredge and Nixon,

1992) contain only about 0.8% OC (d'Anglejan, 7982; Hillaire-Marcel and Fairbridge, lg78). Here we use 4o/o (+ 2%) as an aveïage OC content, which agrees with the overall

t7l average (3.9% OC) estimated for the Ob' River (Gebhardt et al., 2005). Combining 4%

(+ 2%) OC with a sediment load of i I .6 (+ 6) x 106 t a-r, we arrive at a particulate carbon loading from rivers of 0.46 (+ 0.33) x 106 t C a ' (Tabl e 4-4).

Dissolved organic carbon (DOC) represents a second important component of the river material transported to Hudson Bay. Among nine rivers sampled in September-

October (2005) (Great Whale, Little Whale, de Povungnituk,Innuksuac, Nastapoca,

Winisk, Hayes, Nelson, and Churchill River; Table 4-l), the Churchill River had the highest (14.1 mg L-r) and the Great Whale and de Povungnituk Rivers the lowest (2.3 mg

L-r¡ OOC concentrations (Granskog et a1.,2007). Earlier data from the Churchill River showed DOC concentrations of 9 mg L-l in late winter and about 22 mgl-r during freshet

(Chapter 2). DOC concentrations measured in the Nelson River over the summer months

(July to October, n:23) averaged 7 .9 (+ 1 .6) mg L-r. Annual averages estimated for the

Churchill and Nelson Rivers (i5.1 (+ 19.7) mgl-r and 20.5 (+ 15.5) mg L-r, respectively) fall within a similar range (Kirk and St. Louis, 2009). On the Great Whale River, DOC concentrations varied between 3 and 7 mgL-t annually, with the higliest concentrations occurring during spring freshet (Hudon et al., 1996) and concentrations of about 3 mg L-r during the summer (Retamal e|. a1.,2007).

To estimate the average annual DOC input from rivers to Hudson Bay, we take advantage of a relationship between area-specific DOC transports and river runoff.

Hudon et al. ( I 996) found that the area-specific DOC transporls of rivers in the Hudson

Bay and Ungava Bay drainage basins could be reasonably estimated from runoff

1R2:O.SS, n:10) despite differences in drainage basin properties such as vegetation

(Hudon et aL,1996). 'We observe a sirnilarly strong relationship here using the Hudson

172 Bay river data provided by Hudon et al. (1996) (four rivers, 1 sampled in duplicate) and

the additional dxa from Granskog et al. (2007) (8 rivers: Table 4-1). The area-specifîc

DOC transport (t k--'¿') vs. runoff (mm a-') relationship is described by the equation:

transporFO.OOl9 x runoff + 0]0 (12:0.+5, n:l3, p:0.03, not shown).

Using this relationship and the total runoff rates for the six Hudson Bay drainage areas

(Table 4-5), the area-specific DOC transports for the six drainage areas vary from 0.9 t C

km-2 a-r (Manitoba) to 1.9 t C km-2 a-r (east James Bay), implying total transports of 0.25

to 1.45 x i06 t c a-r and an overall total Doc supply of 3.6 x 106 t c a-r. From the

confidence limits on the regression slope and intercept, we estimate uncertainty on the

total supply in the order of t 2 x 706 t C a-t .

Coastal Erosion

Subaerial coastal erosion is an important source of sediment to many of the Arctic

Ocean's shelf seas, with coastal retreat rates of about 3 m al having been observed in the

Kara, Laptev, East Siberian, Chukchi, and BeaufoÍ Seas (Grigoriev et a\.,2004).

However, the impact of this process in Hudson Bay, where relative sea level (RSL) is

declining, is presumably much less than in these other areas where RSL is increasing

(Beaulieu and Allard, 2003). The present rate of uplift of the southern and eastern

coastlines of the Bay is about i cm a-r (Hillaire-Marcel and Fairbridge, 1978; Sella et al.,

2007), and the shorelines are generally characterized by progradation (up to 15 m a-r,

d'Anglejan, 1980). Thermal erosion of permafrost along coastal cliffs is the only major driver for subaerial coastal erosion, and thus observed average annual erosion rates, even in susceptible areas, are relatively low (0.6 m a-t, Zevenhuizen et al., 1gg4). Thermal erosion along a 15 km stretch of coastline in eastern Hudson Bay, where the coastal cliffs

173 (ca. 1 m high) contained 60%o fine-grained soils reportedly generated a sediment supply

rate of about 0.001 x 106 t km-r al (Zevenhuizen et al., lgg4). This rate of sediment

supply is comparable to what occurs along relatively unsusceptible, rocky and non-icy

coasts in the other shelf seas, for example -0.0015 x 106 km-r ar in the Laptev and East

Siberian Seas (Grigoriev et a1.,2004). Assuming a similar supply rate along the -8000

km of shoreline containing at least some stretches of unconsolidated coastal cliffs

(Martini, 1986; Stewart and Lockhart, 2005), we estimate a total sediment supply to

Hudson Bay from coastal erosion of 8.0 (+ 4) x 106 t a-r (Table 4-4). Assuming an

organic carbon composition of 0.8-1% (all terrestrial), yields a total subaerial coastal

erosion input of 0.072 (+ 0.04) x 106 t C a-r. Based on our calculations, it appears that the

input of sediment by coastal erosion is comparable to the inputs by rivers, while the input

of terrigenous OC by coastal erosion is very low (Table 4-4).

Atmospheric Input

Dry deposition fluxes of sediment to the sea surface in the Russian and central

Arctic have been estimated at 570 mg m-2 a-r and fluxes of eolian material to the drifting

ice surface estimated independently at 624 mg m-2 al, yielding an average flux estimate

of 597 (+27) mg m-2 a-r (Rachold et al., 2004). Applying an average organic carbon

content of about 23% (* 15%) implies an OC flux of 137 mgC m-2 a-r (Rachold et al.,

2004). These values are more than an order of magnitude higher than traditionally accepted values (compare fo ca. 77 mg C m-' a-t estimated for the Beaufort Shelf,

Macdonald et al., 1998) but they are based on data collected over several years of aerosol research. Eolian inputs are not believed to be significant sediment sources to Hudson Bay

(Henderson, 1989) but as an upper limit for the budget, we pro-rate the Arctic Ocean

174 estimates to the surface area of Hudson Bay (841000 km2¡ to obtain 0.50 (+ 0.02) x 106 t

ar and 0.12 (+ 0.08) x 106 t C a-' for sediment POC, respectively (Table 4-4).

Mass Balance and Apparent Impact of Resuspension and Lateral Transport

Our calculations of the contemporary supply of sediment from all sources yield an

input of ab out 44 (+ 18) x 106 t a-'. Thit supply represents only one-third the estimated

sediment sink (138 (+ 64) x 106 t a '). This imbalance is beyond the margin of error for

the various inputs. Furthermore, inputs have been independently estimated and are unlikely to be systematically in enor. On the other hand, our figure for sediment burial in the Bay is likely to be an underestimate because small areas of accumulation have not been effectively mapped (Josenhans et al., 1988), and more widespread (>15%) accumulation likely occurs on the inner shelf outside the mapped area used here. If we consider other sinks and losses of particulate matter such as dissolution or oxidation of algal biomass (including siliceous remains) (ca.90o/o, de Haas et al., 2002; Gaye et al.,

2007; Macdonald et al., 1998), our apparent total sink increases to 160 x 106 t a-r (Table

4-4), which is even further out of balance with the calculated sediment input.

To address the imbalance between sediment supply and sedirnent sink, we propose that resuspension and lateral transport of sediments from shallow-water deposits and winnowing from topographic highs, a likely consequence of a long-term decline in

RSL, accounts for the required -1 l6 x 106 t a-r to balance the budget (Table 4-4).

Redistribution of fine particles from coastal (largely glacigenic) deposits in response to changes in RSL has been proposed previously as the major source of sediments in

Manitounuk Sound and Nastapoka Sound in eastern Hudson Bay (Lavoie et al., 2002;

2008;Zevenhuizen et al., 1994). Henderson (1989) proposed that postglacial sediments in

175 central Hudson Bay also largely result from the erosion and remobilizationof nearshore

glacigenic deposits. Although the proposed resuspension term is uncommonly large

compared to shelf seas in the Arctic (e.g., Stein and Macdonald,2004a), or subarctic and

temperate regions (e.g., Dunieu de Madron et al., 2000; Johannessen et al., 2003), the

continuous falling RSL that has characteúzed Hudson Bay, especiatly the southern

portion (rates of almost 1.2 cma-l;, is exceptional (Peltier, 199S). In northern parts of the

Baltic Sea, where rates of rebound approach those in Hudson Bay (as high as 0.9 cm a-l),

resuspension of shallow water sediments and subsequent redeposition in deeper

sedimentary basins is proposed to account for about 80% of the accumulating sediments

((Jonsson, 1992) as cited in (Hakanson et al., 2004) and see also (Jonsson and Carman,

1994; Jonsson et al., 1990; Leivuori and Niemisto, 1995). As old bottom areas rise after

being released from the weight of glacial ice, former areas of active sedimentation reach

the critical depth above which waves can exert a direct influence on, and resuspend, the

sediments (Håkanson et al., 2004; Ignatius et al., 1981). It is equally plausible that

isostatic rebound in Hudson Bay has brought new sea floor areas (especially in the south)

up into the zone where they are subject to erosion from waves, ice scouring, storm surges,

and tides (Hequette et al., r999;Lavoie et a1.,2002; zevenhuizen et al., lgg4),and by renewing the supply of f,tnes, supports resuspension at the millennial scale.

Fine material, which accumulated when the area was inundated by the postglacial

Tyrrell Sea, comprises ca. 60%o of coastal deposits along the low-lying southern Hudson

Baylnorthern James Bay coastlines (Zevenhuizen et al., 1994). Assuming active resuspension and lateral transport along these coastlines (46,900 krn2), where the present rate of isostatic rebound is roughly i cm a-t lHillaire-Marcel and Fairbri d,ge,1978),

176 implies a continuing long-term supply of material of roughly 800 x 106 t a-r (using a

sediment density of I.7 g cm-3). Thus, offshore transport of only about 75o/o of material

made available in this maïìner could supply the 1 16 x 106 t a-r of sediment needed to

balance the sediment budget (Table 4-4). Sub-tidal sediment erosion rates of about 50 cm

per hundred years (equivalent to 400 x 106 t a-' fto* a 46,900 km2 coastal area) have

been observed in a relatively sheltered region of eastern Hudson Bay, apparently

resulting from a combination of ice-scour and weak near-bottom currents (Hequette et al.,

reee).

The proposed resuspension sediment supply (ca. 116 x 106 t a-r: Figure 4-7),

assuming an OC content of 0.8-1% OC, of which44Yo is of marine origin and 56Yo

terrigenous (like our most nearshore core), would also deliver about 0.46 x 106 t C a-l

marine and 0.58 x 106 t C a-l terrestrial carbon. This amount of marine carbon is

insignificant compared to primary production (16.1 (+ 7) x 106 t C a-'); however, this

ancient marine carbon would be highly degraded and recalcitrant, which might help to

explain the dominantly 'marine' character of the Hudson Bay sediments (-80%)

compared with other Arctic shelf seas (e.g., Kara and Beaufort Seas), where terrigenous

OC predominates (Gebhardt et al., 2005; Macdonald et a1.,1998). The input of

tenigenous carbon associated with the apparent redistributed sediment load would

represent almost half of the total input, comparable to that from rivers (0.46 (+ 0.33) x

106 t C a-r: Table 4-4) andmuch greater than that from coastal erosion (0.072 (+ 0.04) x

106 t C a-'¡ or atmospheric inputs (0.12 (+ 0.1) x 106 t C a-r).

The unbalanced terms in the marine and terrigenous POC budgets (15.1 and 0.85 x 106 t C a-r, respectively: Table 4-4) arehere tentatively attributed to oxidation or

177 leaching (conversion to DOC) in the water column (Figure 4-7). This implies oxidation

(or leaching) of about 650/o of the POC¡"., that enters the system, burial of <20Yo and oxidation in surface sediments of 95%) of marine carbon recycling have been estimated for the Beaufort Sea (Goni et al., 2005; Macdonald et al., 1998), productive areas of the arctic Alaskan shelf (Naidu et al., 2004), and the central Kara Sea (Gebhardt et a1.,2005). The apparent oxidation of terrestrial OM in

Hudson Bay seems large considering that Arctic river OM is largely refractory as a result of extensive degradation on land (Dittmar and Kattner, 2003; Hedges et al., 1994; Lobbes et al., 2000), as is recycled OM associated with resuspended sediments (Hedges and Keil, i995). However, a roughly 65% loss seems reasonable in the context of an observed 80%

- 90% loss of modem terrigenous OM and 650/o loss of ancient terrigenous OM on the

Beatrfort Shelf (Goni et al., 2005). Degradation of old OM in Hudson Bay may be enhanced by repeated cycles of resuspension and deposition during transport to ofßhore sinks (Blair et a1.,2004; Hulthe et al., i998).

178 Atmospheric input River input 0.50 11 6 Erosion PP o 21.8 -0 116 lce export Resuspension () ,tr, 138 Leach/oxid

0.60 (A)Sediment Leachioxid

Resuspension

(B) Marine OC POC ------> DOC'.--t-t

Atmospheric input River inout u.4b Erosion 0.072

Resuspension

(C)Terrigenous OC POC ------> DOC,-t-z

Figure 4-7. Budgets (106 t a-r) for sediment (A), marine organic carbon (OC) (B) and terrigenous OC (C). Straight arro\,vs represent particulate organic carbon (POC) and wavy arrows dissolved organic carbon (DOC). Arrows into the water column represent inputs from autochthonous primary production (PP), river input, coastal erosion, resuspension and lateral transport of coastal deposits, and atmospheric inputs. Arrorvs out of the lvater into the sediments represent sediment burial and arrolvs out of the water column to the right represent export by ice and advection or oxidation of DOC. Circular arro\,vs represent internal losses (oxidation or leaching). The resuspension term in the sediment budget was derived by difference to balance the sediment budget and the unbalanced terms in the OC budgets assigned to water column leaching/oxidation.

179 Progressive degradation of sedimentary lignin from the margin to the interior in

Hudson Bay (Chapter 3) supports signif,rcant POC¡.,, degradation during transport to offshore sinks. Further evidence that OC1.,, becomes recycled in nearshore waters is provided by the hnding that these waters are supersaturated with dissolved inorganic carbon (DIC) with respect to the atmosphere, in contrast to undersaturated waters offshore (Else et al., 2008a). DOC data for Hudson Bay waters are needed to properly constrain the conversion of POC to DOC. AlaC measurements would provide an independent assessment of the age of the POC cycled in this region and help constrain the magnitude of glacial-age POC resuspension and redistribution within Hudson Bay.

The budget presented here should be viewed as preliminary, with additional data needed to constrain the budget terms more closely and to confirm the apparent importance of resuspension and recycling of old (glacigenic) OM. Collecting additional sediment cores in close collaboration with seismic profiling would make the sediment sink estimate more robust and confirm the resuspension term. The resuspension process could be witnessed directly by examining sectional data fol optical properties

(%Transmission, see for example Carmack and Macdonald,2002; Johannessen et al.,

2006) and fluxes estimated by deploying sediment traps (e.g., Forest et al., 2007; Forest et a1., 2008; O'Brien et a1.,2006). Exchanges of water, sediment, POC and DOC between

Hudson Bay and Hudson Strait and Foxe Basin are also poorly known, and the export of

DOC is, in particular, an important gap because it likely represents the major OC export from Hudson Bay. Particle flux data collected from a larger number of rivers and with greater frequency, and additional data from which marine primary production could be assessed, will help constrain the Bay's sediment and OC inputs.

180 The budget presented here implies important differences in sedimentary regime

and OC cycling between Hudson Bay and marginal seas of the Arctic Ocean due to a

legacy ofdifferent post-glacial histories. These differences need to be clearly understood

before using Hudson Bay as a sentinel for change in the Arctic's marine environment.

Specif,rcally, in Hudson Bay, the apparent dominance of resuspension and lateral transport of coastal materials in the sediment and terrestrial POC budget implies a susceptibility to the increased storminess and wave base erosion that might accompany reduced sea ice cover (Gagnon and Gough,2005; Wu et a1.,2}}5).Increased sediment and organic carbon burial in the Bay may result, but other factors, such as the carbon loading of particles (e.g., through increased primary production or river delivery) as well as resuspension-driven burial might be necessary for OC burial to increase. Our approach of setting the resuspension flux to balance the sediment budget in Hudson Bay may be appropriate on the time scales of the sediment cores (decades to hundreds of years) but over rnillennial timescales, isostatic rebound represents a continued perturbation to the system with probable consequences for sediment supply (e.g., Ballantyne, 2002). The sedimentary and OC regimes in the Bay may also lag isostatic rebound over the long term, such that modern sedimentation rates reflect sediment supplied to the margins of the system in the past, when regional emergence rates were very high (e.g.,4.5 cm a' about 3500 BP (Gonthier et al., 1993)). If sedimentation and OC burial in the Hudson

Bay system are already in transition due to the history of isostatic rebound, it will be more diff,rcult to predict and measure the additional consequences of river diversions and climate-related changes in sea ice or permafrost or to make inferences fiom this 'sentinel' system to regions elsewhere in the Arctic.

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195 Chapter 5: Elemental and Stable Isotopic Constraints on River

Influence and Patterns of Nitrogen Cycling and Biological Productivity

in Hudson Bay

ABSTRACT

Elemental (carbon and nitrogen) ratios and stable carbon and nitrogen isotope ratios (ðl3C and ô'tN) are examined in sediments and suspended particulate matter from

Hudson Bay to study the influence of river inputs and autochthonous production on organic matter distribution. River-derived particulate organic matter (POM) compositions suggest that terrigenous OM sources to Hudson Bay are heterogeneous, nitrogen-poor and isotopically depleted, consistent with expectations for terrestrial C3 vascular plant sources, and distinct from marine OM sources. Both ôl3C and C/ÌlI source signatures seem to be transmitted to sediments with little or no modification, therefore making good tracers for terrigenous OM in Hudson Bay. They suggest a general pattern of progressively larger contlibutions from marine sources with distance from shore and secondarily fi'om south to nofih, which broadly corresponds to the distribution of river inputs and opposes inshore-offshore gradients in surface phytoplankton biomass.

Processes other than mixing of marine and tenigenous OM influence sedimentary ôl5N values, including variability in the ð'5N of phytoplankton in the Bay's surface waters due to differences in relative nitrate utilization, and post-production processes, which bring

lsN-enrichment about an apparently constant between surface waters and underlying sediments. Variability in the ô'5N of phytoplankton in the Bay's surface waters, in contrast, seefiìs to be organized spatially with a pattern that suggests an inshore-offshore

t96 difference in surface water nitrogen conditions (open- vs closed-system) and hence the

ðl5N value of phytoplankton. The ðl5N patterns, supported by a simple nitrate box-model budget, suggest that in inshore regions of Hudson Bay, upwelling of deep, nutrient-rich waters replenishes surface nitrate, resulting in 'open system' conditions which tend to maintain nitrate ôl5N at low and constant values, and these values are reflected in the sinking detritus. River inflow, which is constrained to inshore regions of Hudson Bay, appears to be a relatively minor source of nitrate compared to upwelling of deep waters, according to the budget. However, the large river inflow may contribute indirectly to enhanced inshore nutrient supply by supporting, in part, the large-scale estuarine circulation in the Bay and consequently entrainment and upwelling of deep water. In contrast to previous proposals that Hudson Bay is oligotrophic because it receives too much fresh water (Dunbar, 1993), our data together with the box model, support most of the primary production being organized around the margin of the Bay, where river flow is constrained.

INTRODUCTION

Coastal oceans are particularly active sites for organic carbon cycling because of larger autochthonous production of organic matter (primary production) relative to the open ocean and additional allochthonous inputs from the surrounding watershed (Liu et a1.,2000; Smith and Hollibaugh, 1993). An estimated 80%-95% of the organic matter preserved in the oceans (and thereby removed from cycling) is buried in coastal ocean sediments (Chen et al., 2003; Hedges and Keil, 1995) but the intensity of physical and biogeochemical processes controlling the organic carbon cycle in coastal seas also provides potential for rapid organic matter turnover (Eglinton and Repeta,2004; Liu et

r9l a1.,2000; Mackenzie et al., 2004). The coastal ocean has been profoundly impacted by

human activities (e.g., damming, eutrophication, over-fishing) and is now expected to be the most sensitive part of the marine environment to global climate change (ACIA, 2005;

Chen et a1.,2003; Macdonald and Anderson, 2008; Mackenzie eT a1.,2004).

Hudson Bay is a large (830 x 103 km2), estuarine, shelf-like sea located at the

southern margin of the Arctic. Hudson Bay and its watershed are rapidly changing as paft

of global climate change (ACIA, 2005; Cohen et al., 1994; Dery et a1., 2005; Stirling et

al., 1999; Tynan and DeMaster, 1997) and it is proposed that the consequences of climate

change may occur sooner and more strongly in Hudson Bay than in the Arctic Ocean,

owing to its more southerly latitude and close association with terrestrial systems (ACIA,

2005; Gagnon and Gough,2005; Gough and Wolfe, 2001; Westmacott and Burn, 1991).

Because few oceanographic surveys have been conducted in the region and consequently

data are extremely sparse, it remains one of the last Arctic areas for which important

components of its modern organic carbon cycle are poorly known (e.g., Saucier et al.,

2004; Stein and Macdonald ,2004a). Hudson Bay receives a very large influx of river

1;, runoff (about 713 km3 a equivalent to about 0.85 m of freshwater distributed over the

entire surface of the Bay (Prinsenberg, 1977;Prinsenberg, 1984). River inflow is, in fact,

largely constrained to nearshore waters in Hudson Bay (Granskog et a1.,2007;

Prinsenberg, TgS6a; Saucier et al., 2004), contributing >3 m of freshwater to the surface

mixed layer along most of the southern coast (Granskog et a1.,2007). River inputs

influence organic carbon cycling in continental shelf sea systems in numerous ways, both

directly by delivering sediments and particulate and dissolved organic materials, and

indirectly by enhancing or reducing marine primary productivity through nutrient inputs,

198 effects on the stability of the water column, estuarine circulation, entrainment and

upwelling (Mann andLazier, 1991). Recent information suggests organic matter

delivered by rivers is widely distributed throughout Hudson Bay (e.g., Granskog et al.,

2007; Chapter 3), while earlier findings suggest river inputs also play an important role in

the Bay's biological productivity, influencing surface phytoplankton biomass (Anderson

and Roff, 1980a) and community structure (Harvey et al., 1997) but perhaps limiting the

Bay's overall productivity (Dunbar, 1993).It remains unclear how allochthonous and

autochthonous inputs compare in overall importance for organic matter preservation and whether the effects of river inputs are direct (e.g., nutrient contributions) or indirect (e.g., through effects on stratification, circulation, light climate or other factors) (Anderson and

Roff, 1980a; 1980b; Granskog et al., 2007; Harvey et al., 1997). A better understanding of the relationship between river inputs, nutrient cycling, and biological productivity in the Bay is critical to predicting the response of the Hudson Bay system to projected climate change (Gagnon and Gough,2005; Johannessen et al., 2004; Lawrence and

Slater, 2005), including major changes in the quantity and timing of river runoff, which are already underway (Dery et al., 2005). Understanding the current nutrient cycle in

Hudson Bay will provide a better basis for drawing connections between productivity, natural climate variability, changes related to freshwater diversions and dams

(Prinsenberg, 1983; Prinsenberg,1994), and 'downstream' consequences such as productivity on the Labrador Shelf (Sutcliffe et al., 1983).

Sediments represent a repository for organic matter produced within a marine system as well as that entering the system from the surrounding watershed. Information about these sources, including the environmental condìtions under which the marine

199 organic matter was produced, can be obtained from various organic geochemical proxies.

For instance, low carbon/nitrogen (C/l.J) ratios and high (or enriched) ð'3C values in marine sediments distinguish organic matter (OM) derived from marine sources

(phytoplankton or ice algae) from contributions of terrigenous OM, which is characteristically nitrogen-poor and isotopically depleted (Fernandes and Sicre, 2000;

Hedges et al., 1988; Meyers,1994; Ruttenberg and Goni, 1997). Use of these proxies together helps confirm that differences are indeed related to source and not post- production processes (e.g., microbial degradation, zooplankton scavenging). Stable nitrogen isotope ratios (ôttN) of sedimentary OM have also been used to estimate relative contributions of marine and terrigenous OM (e.g., Naidu et al., 2000) but ðrsN may, under other circumstances, be used as a proxy for surface water nutrient conditions under which the marine OM was produced (Altabet, 2004; Altabet and Francois,1994; Francois et al., 1992). Because of isotopic fractionation during nitrate uptake for photosynthesis, a large or replenished pool of nitrate leads to lower õrsN values in phytoplankton, compared to conditions under which nitrate is severely drawn down (Altabet, 1988; Saino and Hattori,1987). The variation generated at the surface is transmitted to the sediments when the planktonic OM sinks (Altabet and Francois, 1994;Farrell et al., 1995; Voss et aL,1996).

Here, we use CÀtr, ôl3C and õlsN data for 13 sediment cores and samples of suspended particulate organic matter (POM) to characterize the rnarine and terrigenous sources of OM to Hudson Bay and delineate the importance of these sources to sedimentary OM. We then estimate ôrsN values of the marine-delived portion of

200 sedimentary OM to assess nitrate conditions within the Bay and construct a simple budget to test the plausibility of the inferred nitrate scheme.

STUDY AREA

Hudson Bay is alarge shallow basin, with generally gentle bathymetry, a mean depth of only about 125 m, and maximum depths (200-250 m) lying toward the centre of the Bay (Leslie, 1963). Waters flowing out of the Arctic Archipelago through Foxe Basin enter the Bay at its northwestem end (around Southampton Island), while two-way exchanges with Hudson Strait occur at its northeastern end. There is a basic cyclonic circulation in the Bay, including a strong surface coastal current in eastern Hudson Bay, which intensifies from southwestern Hudson Bay toward James Bay, then north along the eastem shore of Hudson Bay, and finally east through southern Hudson Strait (Ingram and Prinsenberg, 1998; Saucier et al., 2004; Wang et al., 1994). Tides are semidiurnal and have a range of more than 4 m along the west coast and less than half of this along the east coast (Prinsenberg and Freeman, 1986). A nearly complete sea-ice cover forms during most yeals, with ice formation beginning in late October and coverage being virtually complete by December. Maximum ice thickness and extent tend to be reached by April, melt progresses through May, June and July, and open water conditions typically return by the first week of August (Markham, 1986). Melt of the seasonal sea ice cover adds an estimated additional 140-160 cm of freshwater to the surface of the Bay between May and July (Prinsenberg, 1984; Prinsenberg, 1988).

The large seasonal freshwater inputs to the surface waters of Hudson Bay result in strong verlical and horizontal density gradients. During the summer rnonths, the upper portion of the water column is generally strongly stratified, with the depth of the mixed

20r layer varying from about 5-10 m in the southern part of the Bay to 20-25 m in the northern part of the Bay (Granskog et a1.,2007; Prinsenberg, i986b). Salinity in the mixed layer varies from about <24-28 in the coastal waters of the Bay and averages about

30 offshore and >31.5 outside the mouth of the Bay in Foxe Channel and west Hudson

Strait (Granskog et a1.,2007; Harvey et al., 1991). The surface mixed layer thickens and becomes more saline in the fall because of tides, stronger winds, and buoyancy loss and thickens further in the winter due to brine rejection associated with sea-ice formation, ultimately reaching between approximately 50 to 60 m in southern and eastern Hudson

Bay and over 100 m in the northern and western parts of the Bay (Saucier et al., 2004).

The standing crop of phytoplankton in Hudson Bay has been surveyed on only a couple of occasions but appears to be at the lower end of what has been observed in other arctic seas (Subba Rao and Platt, 1984). During a bay-wide survey in August-September

1975, surface chlorophyll concentrations averaged only 0.09 mg chl a m-3 offshore and

0.28 mg chl a m-3 inshore (Anderson and Roff, 1980a). A subsurface chlorophyll a maximum, which was widespread in the offshore, contained 1.8 to 63-times the surface chlorophyll values (0.3-10.75 mg chl a m-\ (Anderson and Roff, 1980b). Many of the same patterns and similar chlorophyll contents (0.25-1.25 mgchl a m-3) were observed during a (September) 1993 transect up the eastern coast of Hudson Bay (Harvey et al.,

1997). Springtime chlorophyll measurements in inshore regions of Hudson Bay indicate a relatively well-developed ice algae community (Freeman et al.,7982; Gosselin et al.,

1985; Michel et al., 1993; Welch et al., l99l). Primary production data for Hudson Bay are scarce compared to the biomass data. Estimates vary from 70-100 g C m-2 al, based on inslrore data (Pett and Roff, lgï2),to 35 g C m-2 a-' for an overall avefage, excluding

202 the main spring diatom bloom, which has yet to be assessed directly, and ice algal production, which may contribute 10 g C m-2 a-' 1Roff and Legendre, 1986). New production was estimated at roughly 25-35 g C m-t a-l by comparison to other arctic areas

(Sakshaug, 2004) and a similar value Qa g C ^'' u-t) derived from carbonate data from the northern part of the Bay (Jones and Anderson,7994).

METHODS

Samples include suspended parliculate organic matter (POM) collected from nine rivers discharging to Hudson Bay (labelled in Figure 5-1, Table 5-1), POM collected from surface waters (0-5 m) at 19 stations in Hudson Strait and Hudson Bay (open squares in Figure 5-1, Table 5-2), POM collected from the subsurface (10-300 m depth) at five of these 19 stations (labelled I-2, Figure 5-1), and 13 sediment box-cores (fîlled circles in Figure 5-1, Table 5-3). The majority of samples were collected during the

ArcticNet (http ://www. arcticnet-u laval. ca) 05 02 expedition, onboard the Canadian research icebreaker CCGS Amundsen, between 15 September and 26 October 2005. The water column was profiled at each sampling station using the ship's rosette, which was equipped with an SBE-9l1plus (Sea-Bird Electronics, Inc.) conductivity-temperature- depth (CTD) sensor and an auxiliary chlorophyll fluorometer (Seapoint Sensors, Inc.). At select stations (denoted with X in Figure 5-1), salinity and nutrient samples were collected from the rosette bottles. Salinity was measured onboard with a Guildline

Autosal 8400 salinometer (precision generally better than 0.002, standardized against

IAPSO Standard Sea'Water) and concentrations of NO¡- + NOz- determined on fresh samples using an Autoanalyzer 3 (Bran+Luebbe) with colorimetric methods adapted from

Grasshoff (1999). Additional nutrients and POM samples were collected from major

203 rivers accessed by zodiac or helicopter deployed from the ship. In a few cases, we also

collected river POM samples from the shore or the ice surface during other periods of the

year.

Figure 5-1. Map of study area and location of samples, including dissolved nutrients (X), particulate organic matter (POM) from surface waters (open squares), POM from subsurface waters (open squares labelled I-l),, and sediment boxcores (closed circles).

River SPM

Suspended particulate matter (SPM) samples were collected from nine rivers

(Table 5-l), which represent collectively about 34Yo of Htdson Bay's annual river inflow

(including James Bay) or -64% of the inflow entering Hudson Bay directly (Dery et al.,

2005). In addition to samples collected during September-October 2005, samples were collected from the Nelson and Hayes Rivers in July 2005, the Gleat Whale, Little Whale, and Nastapoca Rivers in June 2006, and the Churchill River during April-May 2005.

Samples were obtained by pumping water from 0.5-1 m depth and collecting the

204 particulate phase on stacked pre-combusted GF/D and GFÆ glass fibre filters (Whatman,

142 or 293 mm diameter; nominal pore size 2J and 0.7 pm, respectively), mounted on

stainless steel filter holders (Axys Technologies, Sydney, BC). We used a submersible

pump (Shurflo 9300 Series, SHURflo,LLC, Cypress, CA) connected to the filter

apparatus by Teflon-lined stainless steel hose.

Table 5-1. Characteristics of suspended particulate matter (SPM) and dissolved nutrient concentrations in Hudson Bay rivers River Prov. Discharge' Month SPM Nitrate' km3 a-' ô''C (%o) ô''N (%o) Churchilì MB 20.51 3.6 Oct. 9.3 -28.61 3.3 s 2.02 May r 0.5 -27.90 2.s8 April 10.8 -27.95 2.91 April 10.7 -29.82 4.54 April r 0.8 -29.72 4.88 Nelson MB 94.24 16.3 Oct. 16.3 -21.65 5.3 r 6.31 JulY 24.1 -28. I 5 4.05 Hayes MB 18.62 3.2 Oct. 21.7 -28.81 5.51 1.52 July 19.5 -29.42 2.18 Winisk oN 14.69 2.5 Oct. 13.6 -28.92 4.36 1.46 Great Whale QC 19.17 3.4 Oct. 7.6 -26.10 4.63 0.86 June 9.9 -28.98 I .13 Little Whale QC 3.74 0.6 Sept. 12.7 -27.60 4.6s 1.45 June 9.3 -29.30 r .01 Nastapoca QC 1.86 1.4 June 8.6 -29.80 0.61 0.35 Sept. r 0.4 2.12 0.45 Innuksuac QC 3.2s 0.6 Sept. 9.4 -26.28 2.31 0.29 de Povungnituk QC I L63 2.0 Sept. 9.9 -28.98 I .13 2.02 D is c h a r ge l4re i ghted Av e r a ge3 r 6.1 -28.2 3.9 J. / x 2.7 + 0.3 + 0.6 + 2.1 Discharge, runoff and area from Dery et al. (2005) 2Nitrate + nitrite; mean of 1-5 replicates 'weighted average

Filter samples were stored frozen (-20'C) in the f,reld and subsequently at the

Freshwater Institute (FWI), Fisheries & Oceans Canada. Just prior to analysis, they were thawed at room temperature and several sub-samples (-21 mm diameter) were punched

out of the larger filter with a stainless steel punch. The replicates were then oven-dried at

60"C, packaged individually in Petri dishes, and shipped to the University of British Columbia (UBC). At UBC, organic carbon (OC) was measured on acid-decarbonated filter samples using a Carlo Erba NA- 1500 Elemental Analyzer (Verardo et al., 1990) and its stable isotope composition (reported as ôr3C relative to PDB) determined by an in-line isotope ratio mass spectrometer. Total nitrogen (TN) and its stable isotope composition

(reported as ôlsN relative to atmospheric N) was determined on separate untreated subsamples.

Marine POM

Samples of POM from the surface waters of Hudson Bay and Hudson Strait

(Figure 5-1; Table 5-2) were obtained during September-October 2005 by deploying submersible pumps from the deck of the Amundsen. The submersible pumps (Shurflo

9300 Series, SHURflo, LLC, Cypress, CA and Franklin Electric stainless steel submersible water-well pump motor with Micropump pump head) were connected to a f,rlter apparatus on deck by Teflon-lined stainless steel hose. Seawater was drawn from 0-

5 m depth and the particulate phase collected on stacked pre-combusted GF/D and GF/F glass fibre filters (Whatman, 742 or 293 mm diameter; nominal pore size 2.7 and 0.7 ¡tm, respectively), mounted on stainless steel filter holders (Axys Technologies, Sydney, BC).

Samples of subsurface POM were collected from five surface POM sampling stations in the vicinity of the sites of sediment cores 8, 10, 1 1, 74 and 15. The samples were obtained from depths of 10-300 m by drawing water frorn the ship's rosette, which was equipped with l2-L Niskin bottles as well as a SBE-91lplus (Sea-Bird Electronics,

Inc.) conductivity-temperature-depth (CTD) sensor (Seapoint Sensors, Inc.). The water was transfened from the Niskin bottles into stainless steel cans (rinsed three times before use). From these containers, a peristaltic pump was used to collect the particulate phase

206 on pre-combusted GF/F glass fibre filters (Whatman, 142 mm diameter; nominal pore size 0.7 pm), mounted on stainless steel f,ilter holders (Axys Technologies, Sydney, BC).

The marine POM filter samples were treated similarly to the river SPM samples and stored frozen (-20'C) until just prior to analysis, when they were sub-sampled with a stainless steel punch (-21mm diameter). Analysis was conducted at UBC using the same methods described above.

Table 5-2. Properties of particulate organic matter (POM) from Hudson Bay surface waters Latitude Longitude Station Surface Molar Estimated õ''C ô,,N ("N) ('w) Depth (m) satinity c/N' oc/oN2 (/*) (%,) 6t 16.38 64 49.03 262 32.9 6.5 1.2 -2s52 4.69 60 s0.91 64 42.41 381 32.2 6.1 6.1 -21.s0 4.56 6231.47 70 52.08 344 32.1 5.9 6.6 -24.54 5.22 6216.41 71 s8.01 338 32.1 6.1 1.J -22.s2 s. r 9 6419.69 78 05.02 115 31.8 5.9 6.8 -24.66 4.65 64 0t.74 79 12.88 311 31.5 5.8 6.8 -25.00 s.97 6245.46 8I 03.44 198 30.0 6.8 7.4 -25.53 5.31 62 08.07 78 42.87 156 27.9 6.2 7.2 -25.92 4.47 60 09.44 79 01 .46 140 21.1 5.3 6.4 -26.3s 4.22 s5 17.08 77 s3.93 92 20.5 6.2 6.7 -23.35 2.46 54 41.10 79 58.84 61 23.9 6.8 1.6 -23.36 4.09 56 45.07 80 49.75 t78 27.5 5.4 6.4 -25.03 3.54 5823.89 83 17.49 r8r 28.s 6.4 7.1 -26.s1 3.18 57 34.41 91 31.18 60 29.4 l.l 8.8 -23.46 s.83 s9 00.68 8136.68 192 29.6 6.4 1.2 -2s.89 4.15 58 46.90 9t 31.25 80 29.0 6.5 6.9 -24.99 s.17 60 26.19 89 22.29 141 29.8 6.3 1.1 -2s.82 4.24 5123.64 91 56.49 JJ 29.4 17.g3 n.g3 -24.94 5.30 s7 23.64 91 s6.49 JJ 29.4 ruJ3 14.13 -24.91 4.71 A'¡,erage t SD 6.2 7.1 -24.70 4.s6 i J-t + 0.6 + 1.31 + 0.84 Ratio of OC to TN 2Estimated by subtracting positive TN intercept at OC:0 from TN values 3Outliers excluded from OC vs. TN relationship and from summary statistics. (Sarnples collected fronl Nelson/Hayes River estuaries resernble river POM in their propefties.)

201 Sediment Cores

Sediment coring sites (stations 3 to 15: Figure 5-1) were selected along the

Amundsen cruise track from bathymetric and sub-bottom data gathered by an EM300 &

3.5kHz transducer array (http://www.omg.unb.ca). Their sampling and analysis has been described previously (Chapter 3). Briefly, the cores were sectioned aboard the ship, generally into I cm intervals for the top 10 cm,2 cm intervals for the next l0 cm, and 5 cm intervals for the remainder of the core. Sediment from the outer 5 cm of the box was discarded. Each section (sample) was homogenized, subsampled for elemental and isotopic analyses, and stored frozen (-20"C). At the end of the cruise, samples were shipped to FWI, freeze-dried and ground with a mortar and pestle, and from there distributed to the Environmental Radiochemistry Laboratory, Soil Science Department, of the University of Manitoba (U of M) for radioisotope analyses and to UBC for

2lOPb elemental and isotopic analyses. Sedimentation rates derived from profiles and

l37Cs verified against and, in some of the cores, contaminant Pb, varied from 0.032-0.23 g cm-2 a-' lTable 5-3; Chapter 3,4). At UBC, total C and N measurements were performed using a Carlo-Erba elemental analyzer. Total inorganic C was determined by

CO2 coulorneter, and then organic C obtained by difference. Stable carbon (ô'3C) and nitrogen (ô''N) isotope composition was determined by an in-line isotope ratio mass spectrometer, following pre-treatment of the ðl3C subsample with l)%Hclto remove inorganic carbon. Replicate samples (n:16) indicated a precision of +0.3o/o and +1 .60/o for

OC and TN, respectively, and +0.l4%o and,L0.17o/oo for ôr3C and õr5N, respectively.

208 RESULTS AND DISCUSSION

Characterization of terrigenous and marine sources using river and Hudson Bay

surface water POM

River SPM was charucterized by relatively high but variable C/l.J ratios and very low and consistent stable isotopic values (Table 5-1). Weighted according to the proportional discharge of the sampled rivers, CÆlI ratios averaged 16.1 and ôl3C values

-28.2%o (Table 5-1). Variation in CAI ratios among the rivers (e.g., values >16.3 in the

Nelson and Hayes Rivers vs. <13.6 in the others) probably reflects differences in the proportions of nitrogen-poor coarse plant debris (typical CIl.{ ratios of 10-200, Ruttenberg and Goni, 1997) and soil organic matter (typical CA{ ratios of 8-11, Onstad et al., 2000).

These different forms of terrigenous OM as well as differences in age or degradation state of the POM (cf. Guo and Macdonald, 2006) probably also account for the small variations in ðr3C among the river SPM samples (-29.80o/ooto -26.28o/oo; Table 5-1). The river SPM ôl3C values are at the low end of observed values in SPM from other Arctic rivers (-27 .7o/oo to -24.9o/oo, Lobbes et al., 2000; Naidu et al., 2000), which supports a minor contribution fi'om freshwater phytoplankton, as previously suggested

(Montgomery et al., 2000; Retamal et al., 2007). However, freshwater algae can be no more than minor sources, considering the C/lr{ ratios greatly exceed those typical of algal material (6.6, Redfield et aL,1963). The average ðrsN value in river SPM (3.9%o) is slightly higher than values reported for other Arctic rivers (2.640/oo in the Chena River,

Alaska (Guo et a1.,2003),0.31-1 .6%o in the Yukon River, Alaska (Guo and Macdonald,

2006), 7.4-2.8o/oo in the Mackenzie River (Naidu et al., 2000) but values during May-June

(0.670/oo-2.58%o) when greatest particulate discharge occllrs (e.g., Hudon et al., 1996), are

209 comparable. Altogether, the river SPM compositions suggest that terrigenous OM sources to Hudson Bay are heterogeneous, N-poor and isotopically depleted, consistent with expectations for terrestrial C3 vascular plant sources (e.g., Hedges and Oades, 1997;

Onstad et a1.,2000; Ruttenberg and Goni, 1997).

The õr3C of POM from Hudson Bay surface waters varied from -26.570/oo to

-21.50o/oo (Table 5-2) and was significantly higher than values in river SPM, as expected for marine phytoplankton (Rau et aL,1982). Nevefiheless, the average ôr3C value of

-24.7o/oo in surface water POM in Hudson Bay is lower than the average of -20.9o/oo

(n:14) in the nofih Bering, Chukchi, and Beaufort Seas (Rau et a1.,1982) and -22.3o/oo

(n:38) in the North Water polynya (Hobson et al., 2002), which implies the values in

Hudson Bay are influenced by tenigenous OM contributions. Thus, the maximum surface

POM ôr3C value (-27.50o/oo) might be closer to the value in Hudson Bay marine phytoplankton. Terrigenous OM influence on Hudson Bay surface POM is also suggested by the relationship between total nitrogen content (TN) and organic carbon content (OC) in the surface POM (not shown). A positive TN intercept, equivalent to about 10%-20% of TN, at0o/o OC indicates the presence of inorganic nitrogen, likely ammonium associated with clays (Muller, 1977; Schubert and Calverl, 2001). By subtracting the intercept value from TN, we revise our estimates of the CÆli ratios in the surface POM to about 6.4-8.8 (OC/ON; Table 5-2). The low end of this range is comparable to the average CÆ.J value of 6.6 expected for marine plankton (Redfield et al., 1963).

ôr5N values in surface POM varied from 2.460/ooto 5.97o/oo (Table 5-2) and were corelated with surface water salinity (R2:0.50, p<0.0001, n:19) and oxygen isotope composition 1ô180, R':0.+6, a better tracer of river runoff because it is less sensitive to

210 the influence of sea ice melt (cf. Tan and Strain, 1996), although not with ðr3C or CAI.

The upper part of the range of ôrsN values is comparable to typical ôl5N values of about

50/oo-60/oo in marine nitrate in the world's oceans (Altabet, 19S8). The measured POM ðr5N

values may be lower than actual ôlsN values in the organic fraction because of the

influence of inorganic N. Assuming an ammonium ôr5N value of I.8%o-4.1%o (Schubert

and Calvert ,2001), suggests that a contribution of up to 20o/o from ammonium would

decrease POM ðr5N values by less than 7o/oo (e.g., from 5.5%o to 4.8%o).

Composition of subsurface POM and implications for sources and processes

affecting OM composition

ðr3C values in subsurface POM varied from -26.9%o to -20.8%o (Figure 5-2) and were similar to values in surface POM, therefore probably also comprising a mixture of marine and terrigenous OM. There was high variation but no obvious trend in ôr3C with depth in the water column (Figure 5-2). POM previously deposited in shelf areas and subsequently resuspended and transported offthe shelfrepresents a possible source of this variation (e.g., Folest et al., 2008; O'Brien et aL.,2006). Indeed, contribution from resuspended ice algal POM is a plausible explanation for the relatively high ôr3C value of

-20.8o/oo at75-80 m depth in profile I (Figure 5-2). This value exceeds all the õr3C values in surface POM (maximum -21.5%o; Table 5-2). Ice algal POM generally has at least2o/oo

(cf. Hobson et al., 1995; Schubert and Calvert, 2001), and at times (e.g., late in the season) as much as 72%o higher (Tremblay et a1.,2006) ôr3C values than POM derived fronr pelagic phytoplankton. If we assume a ôr3C signature of -21.5o/oo for phytoplankton and -14o/oo for ice algae (consistent with observed end of season values, Tremblay et al.,

2006), then a composition of -20.8%o in POM implies about l\Yo ice algal origin.

211 ilt

A} C'N 6 o

100

t50

200 1s0 1 250 t50 i s00

1 300

200 a,,I 350

B) òilc .2g -26 .2a -22 -2D .28 -26 -21 .92 ,20 o¡*;l;*---"-- d'.ffi 0 o t%"1 t\ i\ t\ i)1 50 ,'l 'I O0 c 'ol .t g !0c \ 150 -* o i1 ô roo¡ ,*l Ò o 2oo G il 6û i ,/ _f tt il 1 i .,ro I eso I ,-oi I i I I 1 :t 90 200 IJ ì *,1 ' 'l "1 :¿r: : ìì: : j ,*j ?00..: i ?5G sso c1 ôr5u 34 67891C t0 ô 0 Ð {o*)

fo0

4Û :.150 r00 200 60

: 250

:: 300

tr roo Ë sso

<> POII4 I St¡rface sed¡rnent

Figure 5-2. Vertical profiles of subsurface POM OC/TN ratios 1a¡, ô13C values (b), and ôrsN values (c). The location of the fTve profiles (labelled I-Iz) is shown in Figure 5-1. Depth of the\'vatercolumn is indicated by grey shaded area. Properties of nearby sedimentary OM are indicated with filled squares.

An ice algal contribution of at most I0o/o to the total deposited marine POM in

Hudson Bay seems reasonable contribution considering that ice algal production

212 represents generally about 3%o of total annual production in Arctic shelf areas (Gosselin et a1.,7997) and a number of factors (e.g., limited zooplankton grazing) favour sinking and preservation of ice algal cells relative to phytoplankton (e.g., Michel et al., 1996;

Wassmann, 1998). Early results from a coupled bio-physical model for Hudson Bay suggest 4%-10% ice algal contributions to total primary production depending on the modelled year (Sibert et al., 2008). Clearly, ice algal contributions will represent a source of uncertainty (e.g., about lo/oo for a 15Yo contribution) in the appropriate marine end- member ôl3C value for Hudson Bay until we learn more about the ice algal signature in

Hudson Bay or obtain some other proxy to distinguish ice algal inputs from those of phytoplankton.

In contrast to ô13C, ðl5N values show a nearly ubiquitous pattern of subsurface enrichment in the vertical profiles of POM (Figure 5-2). ð'sN values reach as high as

8.0o/oo-9.7o/oo in each profile, compared to surface values of only 3.5o/oo-5.2o/oo. The enrichments occur largely within the surface mixed layer at three sites and deeper in the water column at the remaining two (Figure 5-2). Although several processes (e.g., loss of rsN-enrichment, inorganic N) may contribute slightly to this microbial degradation and zooplankton scavenging are probably the most important processes modifying the composition of the OM produced at the surface. It is well documented that isotopic fractionation during bacterial remineralization of suspended or slowly-sinking POM in

lsN-depleted the water column preferentially releases dissolved nitrogen, thus leaving the remaining PoM enriched in 'tN lHolmes et al., 1999; Lehmann et a1.,2002; Saino and

Hattori, 1980). Upward shifts of about20/oo-6%o between ôr5N signals generated in phytoplankton at the sea surface and ðr5N values in underlying sediment have been

2t3 attributed to these kinds of post-production effects in many previous studies (e.g.,

Altabet, 1988; Altabet and Francois, 7994;Francois et a1.,1992 Holmes et al., 1999).

lsN-enrichment The mechanism of is believed to involve kinetic isotope fractionation during protein hydrolysis and thus preferentially removes nitrogen over carbon (Holmes et a1.,1999; Lehmann et a1.,2002), which may account for corelation between the CÆ.,1 ratios of the POM (range 5.2-11.4; Figure 5-2) andtheir ôr5N values (R2:0.23, p:0.018, n:24). Although post-production processes can also affect ô13C, this does not seem to be significant here because preferential microbial degradation of proteins would tend to decrease ôr3C signatures in residual POM by 1o/oo-2%o (Lehmann et al., 2002;Macko and

Estep, 1984), while we see, if anything, increases in the ôl3C values in subsurface POM compared to the values at the surface (Figure 5-2).If, as proposed, ice algal contributions

(up to -10% of total POM) explain the relatively high ôr3C values in subsurface POM

(i.e., > -21.5o/oo, the maximum in surface POM), then the ice algal ðlsN signature may be another source of variation in the ôlsN values of subsurface POM. However, we assume that ice algal contributions would be a source of non-systematic variation, in view of the variable õrsN values in ice algal POM reporled in previous studies (cf. Iken et a1.,2005;

Tremblay et al., 2006).

Elemental and isotopic composition of Hudson Bay sediments

C/lli ratios, ôr3C values and ðlsN values in surface sediments (0-1 cm) in Hudson

Bay are shown in Figure 5-3. Because of the range in sedimentation rates among the cores (0.03 9-0.23 g cm-2 a-r; Table 5-3), the surface sections represent between -4 and 30 years of sedimentation. We did not attempt to adjust for this because sedimentary diagenetic effects on the bulk proxies are generally minimal and the time interval

214 ïepresented in the section was not correlated with CÆrl or ôr3C and only weakly (perhaps coincidentally) with ô'5N 1R':0 .34, p:0.04). The C/1.,1 ratios are calculated from OC/TN with the exception of core 8, in which the OC - TN relationship showed a significant positive TN intercept(0.017%) af 0%o OC, indicating the presence of inorganic N. CA{ values for that core were calculated using TN minus the intercept value. The OC - TN relationships in the other cores either passed through 0,0 or had small positive OC intercepts af 0o/o TN, indicating the presence of N-poor (likely terrigenous) OM.

The surface sediment CAI ratios varied from 7 .9 to 1 1 .5 in Hudson Bay (Figure 5-

3), with trends towards lower values with increasing water depth (R2:0.63), latitude

(R2:0.44) and distance from shore (R2:0.29). ôr3C ratios varied from -24.2o/oo to -20.4%o and showed a parallel spatial pattern, increasing with increasing water depth 1n2:O.Zt;, latitude (R3:0.67) and distance from shore (R2:0.57). The sediments outside the mouth of the Bay and in west Hudson Strait both had relatively high ôr3C values (>-21.5o/oo)but tlre former had a low C/lrl ratio (7 .9) and the latter a relatively high ratio (10.2; Figure 5-

3). Kelp, which is characterizedby high (marine¡ ðr3C signatures combined with high

CÆlI ratios (ca. 17, Goni and Hedges, 1995; Naidu et al., 2000), may represent a source of

OM to the latter core, as suggested by biomarkers specific to kelp (Chapter 3). Overall, the spatial gradients in C/1.{ and ôl3C are consistent with a large portion of the sedimentary OM composition being governed by mixing of marine and tenigenous OM, with other factors providing the rest of the variance. A plot of ôr3C vs. N/C (here N/C is used with ôr3C because this pair is linear under sirnple mixing) for all the sediment samples (n:I72; Figule 5-44) shows considerable scatter but a clear linear relationship between the properties. The N/C ratios and ôr3C values in the sediments also fall mostly

215 between the low values of river SPM and the high values of surface water POM (see outline boxes in Figure 5-44). Slightly higher õ13C values in a few samples may reflect contributions of ice algal POM, as previously discussed.

Independent support for the use of elemental ratios and ôl3C values to trace the distribution of terrestrial vs. marine OM in Hudson Bay sediments is the good correlation observed between lignin (48, which represents the combined yields of syringyl, vanillyl and cinnamyl lignin phenols; see Hedges and Mann, 1979) and N/C (R2:0.48, p<0.0001, excluding two outliers) or ô13C (RÍ:0.78, p<0.0001) (Chapter 3). The strength of the lignin (^8) - ôt3C relationship provides confidence that the ð13C values in particular are generally applicable to tracing the distribution of terrestrial OM in Hudson Bay.

Table 5-3. Properties of cores and surface sediment sections (0-1 cm) Core Dist. from Water r TN OC ô.,C ô.,N Fr^rc' ô"N.rr' shore (km) depth (m) (g (%) (%) (o/*) (oÁo) (%) (o/*) cm -,-a') -1. 365 39s 0.17 0.17 1.13 -2r.50 7.40 85 1.6 4 100 153 0.03 0,16 1 .18 -21.49 7.94 85 8.2 510 tt2 0.10 0.21 1.55 -22.43 7.16 IJ 7.5 62s tt9 0.r2 0.10 0.'79 -23.00 6.28 66 6.7 7 100 106 0.05 0.06 0.58 -23.214 7.ß 63 7.8 I 200 150 0.13 0.17 1.28 -21.88 7.30 80 7.6 925 34 0.23 0.04 0.40 -24.22 6.21 50 6.9 10 260 200 0.04 0.15 r.t2 -20.44 9,81 98 9.8 11 100 86 0.22 0.08 0.66 -21.81 8.23 81 8.6 t2 160 116 0.16 0.15 1.3 I -21.7 4 1 .98 82 8.3 13 290 145 0.08 0.13 0.95 -21.48 8.s6 85 8.8 14 410 244 0.03 0.17 1.18 -20.82 9.34 93 9.5 15 70 430 0.07 0.10 0.85 -20.98 7.01 9l 7.1 'Sedimentation rates calculated from 't'Pb profiles (Chapter 3) 2Estimated proportion of total sedimentary OC that is marine-derived 3Estimated d15N value for marine component of sedimentary OM a1-2 cm section

2t6 O:, : ..zÐkm

13 o o 10

Figure 5-3. Surface sediment CÆlI ratios (a¡, õ13C values (b) and ôrsN values (c).

217 0.20 CORE e3 x4 0.16 -l 5 Â6 Y7 e <8 z 0,12 Þo íi 10 ô 11 }T12 0.08 .' tL) t14 - 15 0.04 -30 -28 -26 -24 -22 -20 ð13C (%o)

10 CORE @3 I X4 +5 ¡b ô o\(J Y7 z

Figure 5-4. Relationships between sediment õr3C values and N/C ratios (a) and ôrsN values (b). Lines show best fit line in a) and example two end-member mixing lines in b), the latter non-linear because they incorporate the difference in CÆ.{ ratios of marine and terrigenous OM (here 6.6 vs. 2l). The two end-member model still provides a poor fit to the õr3C - õrsN relationship. Surface samples of the cores show core label. Dashed outline boxes shorv ranges for river SPM and marine POM.

2t8 Estimates of marine and terrigenous source contributions

We estimated the fraction of marine OC (Fmarc) in sedimentary OM as: Fmarc:

(õt'Coo, - ð13C,",')/(õt'Cn,o,. - ðl3C,"rr), where ôt'Coo, is the value in a given sample, and

ð"C,.,. and ôl3Cn,,u, are the assigned terrigenous and marine end-member values, respectively. We estimated ðl3C:,",-, at -28.2o/oo,the discharge-weighted average ôl3C value in river SPM (Table 5-l), and ô13Cn,'u, at -20.3o/oo. the maximum observed ôl3C value across all the POM and sediment samples.

The fractions of marine OC in the surface sediments of Hudson Bay calculated in this manner vary between about 50Yo and I00yo, with lowest fractions in core 9 and highest fractions in cores 10 and 14 (Table 5-3). Regionally, the fractions of marine OC are about 60%-70% in the southeast, 80%-90% in the northeast and norlhwest, and 85olo-

90% outside the mouth of the Bay (core 3) and in west Hudson Strait (core 15). Thus, the general pattern within the Bay is for progressively larger contributions from marine sources with distance from shore and secondarily from south to norlh. This broadly corresponds to the distribution of river inputs (Dery et al., 2005) and opposes inshore- offshore gradients in surface phytoplankton biomass (Anderson and Roff,, 1980a; Harvey et al., 1997). Considering that the production of marine OM dwarfs inputs of terigenous

OM (Chapter 4), the gradient in the sediments implies intense recycling of marine OM in the water column and hence minimal preservation, combined with substantial preservation of terrigenous OM. This tendency for recycling of marine OM and preservation of terrigenous OM has been documented previously in other Arctic systems

(Macdonald et al., 1998; Schubert and Stein, 1996). However, the dominance of marine

OM in the offshore sediments in Hudson Bay (e.g., cores 10 and 14) contrasts with

219 dominance of tenigenous OM in the sediments of the central Arctic Ocean (Schubert and

Stein, 1997) and most Arctic shelf seas (Gebhardt et al., 2005; Goni et a1.,2000; Naidu et a1.,2000; Schubert and Calvert,200l; Stein and Fahl, 2004). A terrigenous to marine gradient with distance from shore is, indeed, more typical of temperate shelves (Stein and

Macdonald,2004b). Hudson Bay's shallow depth would likely favour preservation of marine OM in sediments, and its seasonally ice-free offshore waters are possibly more productive than the perennially ice-covered waters of the central Arctic Ocean. Ice- rafting of terrigenous matter into the interior of the Bay may also be less important as a mechanism of offshore particle transport in Hudson Bay compared with the Arctic

Ocean. Redistribution of shallow water deposits (largely glacigenic) uplifted into the zone of wave action due to ongoing continuous isostatic rebound in Hudson Bay (Chapter

4) may also represent a source of relatively recalcitrant marine OM, which is more likely to be preserved than the freshly produced OM. Relatively larger contributions of ice algal

POM to the sedimentary OM at deeper sites, which is not unreasonable considering its greater preservation during sinking (Wassman et al., 2004) and later ice-free conditions in the central parl of the Bay (Markham, 1986), would have the effect of enhancing the apparent inshore-offshore gradient in marine carbon. For instance, if ice algal POC supplied 50% of the marine POC in the core i0 surface sediments, the total marine carbon contribution derived from ôl3C values would be only 75%. Such a large contribution frorn ice algal POC would be exceptional, considering what ice algal POC appears to provide an important but not dominant component of the annual vertical carbon flux in other seasonally ice-free areas (cf. Forest el al.,2008; O'Brien el aL.,2006;

Wassman etaL,2004).

220 Stable nitrogen isotope ratios (AttN) and development of õrsN as a proxy for surface

water nitrate conditions

The ôr5N of sediments varied from5.J%oto 9.8%o (6.2o/ooto 9.3o/oo at surface;

Figure 5-3) with spatial patterns that generally corresponded to N/C ratios and ôr3C, including increases with distance from shore (R2:0.67),latitude ß2:0.63) and water depth (rÚ:0.60). These gradients imply that terrigenous OM strongly influences sedimentary ôl5N values in Hudson Bay. However, it is clear from inspection of the ðl3C vs. ôrsN relationship (Figure 5-4B) that a two end-member mixing model (i.e., marine and terrigenous OM) cannot account for much of the variation in ô15N, not even if the differences in CA{ ratio between marine and terrigenous sources are taken into account

(making the mixing line slightly non-linear; Figure 5-48).

Another potential source of variation in ôrsN is variability in the ð15N of phytoplankton in the Bay's surface waters due to differences in relative nitrate utilization

(Altabet and Francois, 1994; Francois et al., 7992; Holmes et al., 1999).In a 'closed system' with respect to nitrate (e.g., stratihed marine surface waters), the ôl5N in remaining nitrate and consequently the marine OM produced from this source progressively increase as nitrate containing the lighter isotope is preferentially used for algal growth. Following Waser et al. (2000), where ôrsN-NO: is the õr5N of nitrate at any point in time, ðrsN-NO¡i the initial ôr5N of the nitrate supply, v the fi'action of nitrate remaining at any time during the drawdown, and E the isotope fractionation during nitrate uptake by phytoplankton, POM with an infinitely long residence time under these conditions has a ðr5N value described by: õr5N-POM : ôrsN-NO zi - E x vl(I-v) x ln v

(i.e., accumulated product equation), and POM produced under these conditions and

22t exported immediately has a ôr5N value described by: õI5N-POM : ôrsN-N Ot - E: ôr5N-

NO¡i - E x ln | - E, (instantaneous product equation). A 'closed system' contrasts with

an 'open system', where nitrate is replenished and thus ðl5N values in nitrate and POM remain at low and constant values: ðl5N-POM: ðl5N-NO3 - E. In the context of these

equations, relatively low ôlsN values are indicative of productive (e.g., nutrient upwelling) areas and high values of oligotrophic areas (Waser et al., 2000).

To evaluate variation in phytoplankton ôl5N values in Hudson Bay, we estimated the ôlsN value of the marine-derived portion of sedimentary OM using the following equation: ôttN*ur: [ôlsN x CA{n*rx (1-F.ur") + ôlsN x C/l.J1"rrX Fn,,ur" - ôl5N1.rrx CÆ.{n,o.x 5N t'N ( 1 -Fn'ur")] / F',u," x CÆlI¡.,r, where ôl and C/NI.-,^. refer to the marine component, ô and CÀtr are the observed values, and ðlsN,.r'. and C/N1",. refer to the terigenous component. F,,,0,. is the fraction of marine (vs. terrigenous) OM contributing to the sample as estimated from ðl3C (e.g., see surface sediment values in Table 5-3). CA{n,o,is assumed to be 6.6 (i.e., Redfield ratio, Redfield et al., 1963) and ôr5N,.,, is estimated at

3.9%o,the average ôr5N value in Hudson Bay river SPM (Table 5-1). Selection of an appropriate value for C/l{1.,, is more difficult because of the variable CÆ.,1 ratios in river

SPM (7.6-24.1; Table 5-1). We thus estimated an appropriate CÆl11.,, value fi'om the observed CA{ ratios in the sediments and the equation: C/NI'.,,: (CN - Fmarc x

C/l'{mar)/(1-Fmarc). This approach yields an apparent average C/1.J1.,. of about 2l in sedimentary OM, which is similar to the average C/ÌlI ratio of 16.1 in river SPM (Table 5-

1).

Derived ð'tNn,u,values for surface sediments vary from 6.5o/ooto 10.3%o (Table 5-

3). It should be noted that these values do not represent the original surface-generated

222 phytoplankton ðrsN but rather the ðrsN of phytoplankton-derived OM after it has gone

through post-production processes on the way to burial (see section 4.2).If we derive

ôl5N.u.values for surface water POM in a similar manner, we obtain values mostly

between 3o/oo and 60/oo.

The derived ðl5N-u, values in the sediments show a very similar spatial patteïn to

that of the bulk ðr5N values, with progressively higher values with distance from shore

(R2:0.63, p<0.001; Figure 5-5), water depth 1R2:0.S4, and from south to north

1n-2:O.SO;. We infer that river inputs not only contribute terrigenous POM with a low

ôr5N signature, which affects the bulk õr5N value by mixing, but also influence the bulk

ðlsN value indirectly by affecting the nitrogen conditions (open- vs closed-system) of the inshore surface waters and hence the ôl5N value of phytoplankton. The progressively increasing ðlsNn,u,values with distance from shore (Figure 5-5) imply that inshore waters have enhanced nitrate supply or renewal and hence probably productivity. The estimated

ôr5Nn.o.values outside the mouth of Hudson Bay (7.6, core 3; Table 5-3) and in west

Hudson Strait (7.1, core l5) are comparable to the values in the Hudson Bay inshore regions, and thus imply similar (relatively high) nutrient supply. We find the highest

ôl5N,.,,'n,values inthe offshoreportions of HudsonBay (9.8, core 10; 9.5, core 14), which suggests oligotrophic conditions.

A critical assumption here is that the influence of post-production processes on

ôr5N values is relatively consistent within Hudson Bay. Our vertical profiles of POM

(Figure 5-2) do not reveal any obvious systematic variations in the degree of post- lsN-enrichment production with, for instance, water depth or among different regions of the Bay. To date, the bulk of the evidence from numerous other studies also supports a

223 constant offset, within a given area, between the surface-generated ôl5N signal and the underlying signal preserved in sediments (Francois et al., 7992; Holmes et al., 1999).

Some recent observations spanning seasons and transitions from spring blooms to summer conditions, suggest larger diagenetic õlsN alterations of sinking POM during low productivity and low POM flux periods (Lourey et a1.,2003). That kind of variation may weaken the connection between ôl5N in sediments and surface water nitrate supply.

Howevet, for systems in which nitrate limits primary productivity, the same kind of regional patterns would be expected from flux-related variations in the effects of post- production processes and those produced by differences in nitrate supply (i.e., low ðl5N in high productivity regions, high ð'5N in low productivity regions).

Nitrate distribution and biogeochemical budget

Enhanced nitrate supply and productivity in the river-influenced inshore region of

Hudson Bay, compared to the offshore, may be related to a number of different processes including nutrient supply to surface waters by river inputs or replenishment from below through upwelling of deep, nutrient-rich waters as a result of winds, tides or physical entrainment (estuarine circulation) etc. A positive correlation between ôt'Nn.,u, values in surface water POM and surface water salinity ç*:0.++,p:0.002; Figure 5-6) or oxygen isotope ratio (ð180, R2:0.40; a better tracer for river water because it is less affected by sea ice melt) supports river inputs (and not simply water depth or other inshore characteristics) being associated with lower ôr5N values in phytoplankton. Neveftheless, the association may be coincidental, with liver inflow being constrained to inshore areas where some other processes enhance nitrate supply.

z-¿-+a1 A CORE rì i; 6Þ3 't x4 ,, -r5 ãq- u:ì ,r6 ¿€oo tr t$ v7 -:. t- * 9 Ë Ê¡tr .:'l n 10 roz8 x3 o 11 -iþ .rl .a:, y.12 ,a 13 114 - 15

300 400 Hudson Distance from shore (km) Strait

Figure 5-5. Relationship between estimated õrsN values in marine-derived sedimentary OM (õttNr,,.) and distance from shore.

y=0.26x - 2.60 R2=0.44, p=0.002

eÐ o t èR o l* oo tÐ o û Ë o roz d 'ln t o c

2 2A 25 30 35 Surface Salinity

Figure 5-6. Relationship between estimated ôlsN values in marine-derived surface water POM (ôttN',or) and surface lyater salinity.

225 Nitrate (+nitrite) concentrations measured during the September-October cruise

(unpublished data of J.-E. Tremblay,2005) were generally between 0.1 and 1.0 pM in

surface waters (top 20 m) and showed no clear spatial pattern nor correlation with either

surface water salinity or the ôl5N values in surface water POM. This seems inconsistent

with river inputs directly supplying surface waters with nutrients. However, it is not

surprising considering the relatively low nitrate concentrations measured in river waters

during the cruise (0.29-2.02 pM with the exception of the Nelson River, 6.31 pM; Table

5-1). Although data are scarce, seasonal samples from the Great Whale River suggest that

nitrate concentrations remain relatively consistent and low tluoughout the year (Hudon et

aL.,1996).

Nitrate concentrations were much more variable in the near-surface water layer

(20-80 m) in Hudson Bay in fall 2005 and, in particular, were generally higher in this

layer at inshore stations, compared to offshore stations. For instance, most inshore

stations had nitrate concentrations in the order of 5.0-7.5 pM at 50 m depth, whereas

offshore stations had concentrations of 1.5-4.5 pM (Figure 5-7). Similarly, nitrate

inventories in the upper 50 m of the water column were about three-fold higher at inshore

stations than at offshore stations (roughly 50 vs. 150 mmol.m-'¡. These differences in vertical distribution of nitrate lend support to the hypothesis of nutrient enhancement of inshore surface waters by upwelling rather than river nutrient input to surface waters.

226 E o G z

Salinity

Figure 5-7. Nitrate (+nitrite) concentrations (pM) plotted against salinity for samples collected in September-October 2005. Sample locations are shown in Figure 5-1. (Plot produced using Ocean Data View, R. Schlitzer, http://odv.awi.de.)

To evaluate whether or not the nitrate control on primary production, as inferred fi'om ôr5N distributions, and supply by upwelling, as inferred from nitrate distributions, is plansible, we constructed a simple nitrate budget for the Bay using LOICZ biogeochemical modelling guidelines (Gordon et al.,1996). This model rests on a salt balance, which defines water exchanges between sub-basins using average net freshwater supply and assuming a representative salinity for each sub-basin over time. Effective volume flows and exchanges implied by the salt balance then allow calculation of

227 interbasin transfers of nutrients. Differences between export and import from the sub-

basins imply sinks (e.g., nutrient uptake) or sources (regeneration) (e.g., Wulff and

Stiegebrandt, 1989). We constructed the budget using a two-layer stratified water column

(0-50 m surface layer) and separate inshore and offshore regions (after Anderson and

Roff, 1980a) to produce four boxes (i.e., four sub-basins). Consistent with the Bay's large

scale estuarine circulation (Saucier et al., 2004), saline waters flow into the system

(specifically the offshore deep basin) at depth from Hudson Strait/Foxe Channel and less

saline waters flow out from the surface part of this basin into Hudson Strait/Foxe

Channel. A representative salinity was estimated for each sub-basin and outside the

mouth of the Bay (Table 5-4) based on Prinsenberg (1986b;1987;1979). River input

enters the inshore surface water compartment, and net precipitation less evaporation is

lost from both surface compartments (Figure 5-8). Freshwater inputs from sea ice melt

are not included because we assume these are balanced over an annual cycle by

freshwater withdrawal during ice formation (Prinsenberg, 1984).

Representative annual average nitrate concentrations were estimated for each sub- basin (Table 5-4) from the 2005 measurements (Figure 5-7) and previously published

data, which included measurements made in the winter in several coastal areas (Gosselin

et al., i985; Irwin et al., 1988; Legendre et al., 1996;Legendre and Simard, 1979; Welch

et a1., 1991). No data are available for offshore areas in winter and thus we assumed that more saline coastal samples were most representative of winter conditions offshore. The influx of nitrate from rivers was estimated from the volume of river discharge (713 x 1 0e r.r' u-'; Prinsenberg, 1984) and the discharge-weighted average of the measured nitrate values (3.7 ptM; Table 5-1). This value is within the range of nitrate concentrations

228 observed in Hudson Bay rivers during the winter (I.3-2.1 ¡rM in the Great Whale River

(Hudon et a1.,7996), ca. 6 pM in the Churchill River (Chapter 2). A loss of nitrogen to

sediments for the system as a whole was estimated assuming net annual sediment

accumulation of 138 x 106 t a-t (Chapter 4) and an average of 0.13Yo sedimentary

nitrogen content. Finally, inflowing seawater from Hudson StrailFoxe Channel was

assigned a concentration (3.6 pM) to balance the overall nitrogen budget.

Table 5-4. Choices for'best estimates' (anel ranges) for representative salinity and nitrate in budget Region Salinity Nitrate (¡rM) Offshore surface 31.5 (31 .2-31 .8) 2.1 (1.2-2.6) Offslrore deep 32.9 (32.7-33.1) 10.8 (9.4-12.2) Inslrore surface 30.5 (30-31) 2.6 (1.4-3.4) Irrshore deep 32.2 (31.9-32.5) 9.3 (7.5-11.1) Inflow from Hudson Strail 33.0 (32.9-33.1) 3.6+ Foxe Channel +Set at value which balances the budget

The apparent total water exchange time implied by the budget (total volume of the

Bay divided by total inflow) is about 9 years, which is within the range of previous

estimates (4-14 years, Pett and Roff, 1982). The exchange time for the surface

compartments is less than2 years, while previous estimates are less than one year

(Harvey et al., 1997). This general agreement in water mass exchange rates provides confidence in the overall structure of the model.

229 Offshore lnshore

Water -103 0 -95 713 (10'mta') T/tt +/ 2.'15x10" m' 2.01xl 0'' m' Surface

+2973 +3234

4.30x10" r¡t' 1.94x10" m3 Deep

Salt (10nm'at) +258429

52=33 +97 797 +25293't

Niträte+nitr¡te -0 2.62 (10'mol a')

+12.6f +24.8 +21.7 |

Figure 5-8. Biogeochemical model for Hudson Bay for freshwater, salt and nitrate (+nitrite). The Bay is represented as a two-layer stratifïed water column (0-50 m surface layer) and separate inshore and offshore regions to produce four boxes (i.e., four sub-basins). Net volume flows are represented by one-way arrows and mixing exchanges by two-way arrows.

The export of low salinity surface waters from inshore to offshore surface waters and ultimately out of the Bay drives roughly 2.5-fold greater upward fluxes (vertical entrainment) in the inshore region compared to the offshore (Figure 5-8). Similarly, modelled vertical mixing exchanges related to processes such as tidal mixing are about

230 four-fold higher in the inshore region compared to the offshore. Consequently, the nitrate budget shows the inshore surface waters receiving about 75.2x 10e mol nitrate a-lvia entrainment of nutrient-rich deep inshore waters and an additional upward flux of 22.6 x

10e mol nitrate a-' fro- vertical mixing. Altogether, this implies an upward transport of nitrate in the inshore region more than 2.5 times that in the offshore region. River inputs of nitrate supply only about 2.62 x 10e mol a-l to inshore surface waters. Thus increased mixing and physical entrainment of deep, nutrient-rich waters in the inshore region is the significant source of nitrate for primary production, and not river inflow. The same conclusion was reached for the interior Arctic Ocean shelves (Macdonald and Anderson,

2008). Large differences in net nitrate uptake (AN) between inshore surface waters (AN:

-76.9 x l0e mol a-r) and offshore surface waters (AN: -25.3 x 10e mol ar) (Figure 5-8) also imply greater primary production in inshore waters. The uptake values correspond to production of about 7 .5 gC m-' a-l of marine organic matter in the inshore region

(assuming a C/l\ ratio of 6.6), compared to 2.4 gC m'2 a-' in the offshore. Net releases of nitrate in both inshore and offshore subsurface compartments (+20.6 x 10e mol a-l and

+71.6 x l0e mol a-I, respectively) show net nutrient regeneration in those areas.

Only major changes in the nitrate or salinity values chosen for the box model would alter the fundamental conclusion that vertical entrainment enhances nutrient supply in the inshore region. Increasing river nutrient fluxes, for example by including dissolved organic nitrogen at concentrations similar to those observed in other Arctic rivers (e.g., l3 pM in the Lena River, Lara et al., 1998), still leaves this supply at only one-sixth the vertical flux of nitrate into inshore surface waters from deep water' upwelling. Varying salinity or nitrate concentrations within the estimated ranges (Table

231 5-4) changes the estimated fluxes and AN values generally by less thanZlo/o. The model was most sensitive to lowering the surface salinity (which changes the upwelling flux by up to 50%) and to lowering the nitrate concentration in inshore deep waters, with concentrations less than 8.5 ¡rM resulting in a negative AN, implying net uptake rather than release of nitrate in that compartment.

Implications of the box model and õrsN distributions

Our data, together with the box model, suggest that in inshore regions of Hudson

Bay, upwelling of deep, nutrient-rich waters replenishes surface nitrate, resulting in 'open system' conditions, which tend to maintain nitrate ôl5N at low and constant values, and these values are reflected in the sinking detritus. Tides, winds, buoyancy differences and seasonal convective instabilities are all known to contribute to vertical mixing throughout the Hudson Bay system, with probably variable contributions from one sub-region to another, especially where large freshwater inputs circulate through a coastal cunent in southern and eastern Hudson Bay (Saucier et al., 2004). Certainly, there are also seasonal variations, with tides being of greater relative importance to vertical mixing under the ice

(Gosselin et al., 1985). Entrainment and consequent upwelling where river runoff circulates though the coastal current would be potentially most signif,rcant during summer and early fall because the waters from southern Hudson Bay reach Hudson Strait by late fall (Saucier et al., 2004). Nutrient entrainment related to the large river inflow into Hudson Bay has not yet been studied but previous studies in estuarine systems have shown that the process of entrainment can generate brackish-water flows that exceed 30 times the freshwater input (Greisman and Ingram , 1977; Sutcliffe, 1972 Yin et al., I 995) and consequently entrain large quantities of nutrients, e.g.,75Yo of total supply to

232 euphotic zone in the St. Lawrence Estuary, compared to <25o/o from freshwater inputs

(Greisman and Ingram, I97l).

The ðr5N data and budget suggest that offshore regions of Hudson Bay, which are characteÅzed by strong stratification and slow circulation patterns, fit a 'closed system' scenario with respect to nitrate. Our interpretation is that the nitrate pool in the interior basin of Hudson Bay, established by mixing over winter, may go without resupply until the following winter due to stratification. An additional implication is that the subsurface chlorophyll maximum (SCM), which was widespread in the interior of the Bay

(Granskog et al., 2007), was perhaps mostly a recycling community utilizing primarily old nitrogen, rather than nitrate supplied from below. We had expected that the SCM would maintain low ôl5N values and transfer these to the sediment in the offshore, representing a large proportion of the annual new production in that stratified region, as found elsewhere (cf. Weston et al., 2005).

Hudson Bay has been proposed to be an oligotrophic sea because it receives too much fresh water (Dunbar, 1993). The primary production estimates derived from the nitrate budget certainly give an image of a low-productivity ocean (2.4 - 7 .5 g C m-' u' new production from interior to margin). However, our data together with the box model suggest that entrainment by the large river inflow (about 713 km3 a-'; muy actually contribute, along with wind- and tidally driven upwelling, to supporting primary production in the Bay and organizing it into the rnargin. This organization perhaps in part explains the importance of the inshore region as habitat for high trophic-level animals

(Stewart and Lockhart, 2005). Additional data are needed to assess the relative importance of winds and river runoff-related entrainment to upwelling because each

./.J 3 process will be affected independently by climate change, past, present and future.

Decreases in the total annual river inflow to Hudson Bay during the past 40 years (Dery

et a1,2005), or changes in flow between winter and summer (for example due to

damming; Dery et a1.,2005; Prinsenberg, 1983; Prinsenberg, 1980), must be viewed

critically in their potential to decrease biological productivity, especially in the inshore

region, or alter the timing of nutrient supply so that it does not match the light cycle. If

wind-driven upwelling is most important in supporting inshore primary production, we

may f,rnd a record in sedimentary OM of the system's response to wind changes during

the unusual positive phase in the Arctic Oscillation in the 1980s-90s. The offshore

nitrogen regime is probably less sensitive to changes in river discharge and wind but

could be more affected (and productivity possibly enhanced) by reductions in sea ice

(melt water playing a larger relative role than river discharge in the vertical stratification

of the water column offshore).

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245 Chapter 6: General Discussion and Conclusions

Coastal seas are important sites of organic carbon cycling, with relatively larger

inputs of organic matter (OM) from marine primary production and allochthonous

(terrigenous) sources than open ocean areas, and more active physical and biogeochemical processes (Chen et al., 2003; Liu et al., 2000). Arctic shelf seas provide even more dynamic environments for organic carbon cycling because of their intense

seasonality and sea ice processes (Eicken, 2004; Rachold et al., 2004). As temperatures and sea ice conditions change rapidly (ACIA, 2005; IPCC, 2001; IPCC, 2007),the

Arctic's organic carbon cycles will be altered, with consequences for local ecosystems and global carbon cycles (e.g., Chen et al., 2003).In-depth knowledge of the sources and chemical composition of OM and the major processes controlling its transport, transformation and burial is needed to predict and measure changes in the years to come.

This thesis addressed knowledge gaps in our understanding of the major sources and processes forming the modeln organic carbon cycle in Hudson Bay, a large but poorly-known Arctic inland sea, which is undergoing change more rapidly than other

Arctic areas (ACIA, 2005). Through the application of various geochemical tools, the organic contents of dated sediment cores and suspended particulate samples fi'om marine and river waters were characterized to obtain new data on the origin, composition, distribution, and rates of transformation and burial of marine and terrigenous organic matter in the Hudson Bay system. The new data were synthesized with literature data to develop the first Bay-wide budgets for sediment and particulate organic carbon, and to construct a biogeochemical (LOrcD box model for nutrients (nitrate).

246 The new data together with the modelling provide significant new insights into the major sources and processes controlling OC cycling in Hudson Bay (Figure 6-1).

Based on ôl5N data, marine primary production, which represents the principal source of marine OM, is organized primarily into the inshore regions of the Bay, where there is enhanced nitrate supply to surface waters, compared to offshore regions (Figure 6-1). A nitrate box model indicates that the major source of nitrate is stronger upwelling of deep, nutrient-rich waters in the inshore domain, compared to the offshore domain. By comparison, river inflow represents a relatively minor source of nitrate; however, the large volumes of river water circulating through the coastal cunent in the inshore region may support the enhanced primary production indirectly, by contributing to nutrient entrainment and consequent upwelling.

River inflow represents the most important source of allochthonous (terrigenous)

OM to Hudson Bay. Bulk organic composition data and specific biomarkers (lignin) indicate that the terrigenous OM inputs are heterogeneous, including both plant debris and liighly-altered, mineral-associated OM. The distribution of OM components in the

Bay's sediments reveals a strong influence of hydrodynamic sorting of the terrestrial material in the coastal zone, with coarse material largely retained near river mouths and a fine-grained fraction transported more widely by marine cuments (Figure 6-1).

Disposition of the landfast sea ice and the rubble ice at its seaward margin also affects the dispersal of terrigenous OM during winter and spring.

A very important process supplying both marine and terrigenous OM to the interior of Hudson Bay is resuspension and lateral transport of fine-grained sediments from shallow-water areas (Figure 6-1). An unusual importance of this process in Hudson

247 Bay compared to other Arctic areas is attributed to the Bay's history of continuous isostatic rebound (relative sea-level fall) throughout the postglacial period and continuing today atarate of about 1 wt/a (Peltier, 1998). Similar to the process described in the northem Baltic Sea, where rates of rebound approach those in Hudson Bay (as high as 0.9 cmla), ongoing relative sea-level fall in Hudson Bay may effectively raise former areas of sediment accumulation (e.g., seabed from the postglacial Tyrrell Sea) up into the zone of wave action and ice processes, and thereby renew the supply of fine sediments available for resuspension and transport (Figure 6-1).

I ns h ore-offsho re tra nse ct

E c o oo)

0 Distance from shore (km) Figure 6-1. Schematic illustrating major sources and processes in the Hudson Bay OC cycle

248 The old, likely glacigenic, organic matter transported with the apparent resuspension sediment supply is significant in the Bay's overall organic carbon budget, with the terrigenous fraction representing about one-half the Bay's total terrigenous OM supply (Figure 6-2), and the marine fraction likely contributing to efficient burial of marine carbon in Hudson Bay's sediments, compared to other Arctic areas.

Consistent with an unusually active resuspension process, patterns of sediment accumulation and OC burial on the Hudson Bay seafloor are complex. However, careful account of mapped postglacial deposits, together with sedimentation rates derived using sediment-core models incorporating both sedimentation and mixing, reveals a sediment sink of comparable size to that in other Arctic areas. Organic carbon burial rates in

Hudson Bay (1.03 x 106 t C a-l and0.23 x 106 t C a-l for marine and tenigenous OM, respectively; Figure 6-2) also lie within the range of rates in other Arctic coastal seas.

Rates of oxidation in the water column and surface sediments are also generally consistent with Arctic systems elsewhere, perhaps slightly lower than average for marine

OM because of recalcitrance of the resuspension supply compared to modern inputs, and higher than expected for highly-degraded terrigenous OM, related to repeated cycles of deposition and resuspension.

249 Figure 6-2. Inputs and fluxes of marine and terrigenous organic matter in Hudson Bay.

Below, the major fìndings of the research are discnssed in greater detail, placing the new knowledge about Hudson Bay into the context of previous work in the area and our understanding of the other Arctic coastal seas. A final section highlights the sensitivities to climate change brought to light by the research, and the fundamental similarities and differences between Hudson Bay and the Arctic Ocean marginal seas, which must be understood in detail if Hudson Bay is to be used as a sentinel for change in

Arctic coastal marine systems.

250 MARINE PRIMARY PRODUCTION AND ITS CONTROLS

Hudson Bay has been widely considered an oligotrophic region because of strong

stratification due to large freshwater input from river runoff and sea-ice melt (cf.

Anderson and Roff, 1980a). Low or incomplete nutrient (nitrate) regeneration in the

Bay's deep waters has also been hypothesized (Legendre and Simard,1979; Pett and

Roff, 1982). However, Hudson Bay's oceanography has been the focus of only very

limited study, and nutrient sources and processes controlling nutrient availability remain poorly understood (Anderson and Roff, 1980b; Harvey et al., 1 997;Petf and Roff, 1982).

Annual production estimates extrapolated from regional datavary almost four-fold and the regional patterns of production and relative contributions to total production have yet to be directly assessed (Lapoussière et al., 2009; Sibert et al., 2008).

Organic compositional (ð'sN) data together with the biogeochemical box model for nitrate developed in the thesis (Chapter 5) reinforce the general view that Hudson Bay is an oligotrophic system (Anderson and Roff, 1980a; Pett and Rofl 1982; Roff and

Legendre, 1986). Primary production estimates derived from the nitrate model suggest

Bay-wide annual new production of about 10 g C m-' a,'. This estimate is lower than a previous estimate Qa g C --' u-') derived fi'om carbonate data from the nofthem part of

Hudson Bay (Jones and Anderson, 1994) but this difference is not surprising in view of the large regional variations in water mass characteristics, biomass parameters, etc. (cf.

Anderson and Roff, 1980a; Harvey et al, 1997), which complicate extrapolation from regional data. The new estimate (10 g C m-2 a-'¡ is comparable to estimates for oligotrophic interior shelf seas of the Arctic Ocean (Wallace et al., 1987), including the

Beaufort Sea, where new production was recently estimated at 12-16 g C m-2 a-l

251 (Carmack et a1.,2004). The value is much lower than the -30 g C m-2 a-l estimated for the Canadian Archipelago (Michel et al., 2006). Considering that Hudson Bay partly shares source waters with the Archipelago (i.e., Arctic Ocean outflow), the implication is that the controls on annual primary production in Hudson Bay differ fundamentally from those in this neighbouring sea. Recent observations of very low organic matter sinking fluxes in fall in Hudson Bay, comparable to background winter values in highly productive Arctic seas (e.g., Barents Sea), offer further support that late-summer primary productivity in Hudson Bay is at the low end of the range for Arctic areas (Lapoussière et aL.,2009).

The õr5N data suggest that the principal source of new production in Hudson Bay is phytoplankton production in surface waters, with little contribution from the subsurface chlorophyll maximum (SCM), despite its widespread occurrence through the interior waters of the Bay (Anderson and Roff, 1980b; Granskog et al., 2007). This result was unexpected because the SCM has been found to represent a large proporlion ofannual new production in other stratif,red regions (cf. the North Sea, Weston et al., 2005). The

SCM in offshore Hudson Bay is perhaps mostly a recycling community utilizing primarily old nitrogen, rather than nitrate supplied from below. Low contributions to total

(new) production from the SCM are consistent with recent observations of low fluxes of biogenic silica and diatom-associated carbon in fall in the offshore regions of Hudson

Bay (Lapoussière et a1.,2009). Contributions from ice algae of up to about 10% of the total marine particulate OM input are consistent with the compositional data.

The overall picture that emerges from the organic composition data together with the nitrate box model (Chapter 5) is one of about 70o/o of the new primary production in

252 Hudson Bay being concentrated in the surface waters of the coastal or inshore domaln

(within 100-150 km from shore; Figure 6-1). Organic compositional data indicate that this production is related to enhanced nitrate supply to surface waters in the inshore region, compared to the offshore, with upwelling of deep, nitrate-rich waters providing the primary source of nitrate (Figure 6-1). Upwelling, together with increased tidal- mixing exchanges, results in about 2.5-times the vertical transport of nitrate to inshore surface waters, compared to what occurs in the offshore.

This picture of the spatial organization of primary production and nutrient availability in Hudson Bay is significant in that it implies that river water supports rather than suppresses the Bay's annual new primary production. It has been proposed previously that Hudson Bay is oligotrophic because it receives too much fresh water, especially river inflow (Dunbar, 1993). River inflow adds the equivalent of about 0.8 m of freshwater to the surface of Hudson Bay annually (Prinsenberg, 1984) and thereby supports, together with sea ice melt, the surface freshening and strong stratification that charuclerizes Hudson Bay during summer (Figure 6-1, Granskog et a1., 2007;

Prinsenberg, 1988). However, recent studies indicate that river waters are tightly constrained to the inshore domain, producing a freshwater inventory there which exceeds

5 m in places (Granskog et a1.,2009). The river waters, together with wind stresses, establish the Bay's strong cyclonic coastal curuent (Figure 6-1, Saucier et al., 2004). In this context, then, it seems likely that river runoff contributes to supporting the concentration of new production and enhanced nitrate supply in the inshore domain, specifically, through entrainment and upwelling of nutrients due to the strong brackish- water surface flows in the coastal cunent. This process would be potentially most

2s3 significant for replenishing nutrients in inshore surface waters during surnmer and early fall, because the river waters from southern Hudson Bay reach Hudson Strait by late fall

(Saucier ef a1.,2004). Previous studies in estuarine systems have shown that the pÍocess of entrainment can generate brackish-water flows that exceed 30 times the freshwater input (cf. Greisman and Ingram,1977; Sutcliffe, 1972;Yin et al., 1995) and consequently entrain large quantities of nutrients (e.9.,75o/o of total supply to the euphotic zone in the

St. Lawrence Estuary, Greisman and Ingram,1977). This would mean that river inflow contributes, albeit indirectly, to the new production that occurs in the inshore surface waters, which appears to be the major proportion (-70%) of overall new production in

Hudson Bay.

For offshore regions of Hudson Bay, ôl5N data and the nitrate box model suggest that nutrient resupply to surface waters occurs infrequently; indeed, the nitrate pool, which is established by vertical mixing during winter, may go without resupply until the following winter due to stratification. Therefore, in contrast to the inshore domain, the major limit on new production in the offshore domain would be the depth of winter mixing, which depends on conditions during fall (freshwater expott, fall winds and storms), winter ice production, and winter conditions in estuaries, which govem the distribution of winter river inflow. This nutrient regime resembles that in the Amundsen

Gulf, adjacent to the Beaufort Sea, where primary production in a given summer depends primarily upon the nutrient renewal in surface waters during the previous fall and winter

(Tremblay et al., 2008). However, whereas large-scale circulation patterns and conditions genelated remotely (e.g., on other Arctic shelves) affect the depth of winter mixing in the

Amundsen Gulf, variations in offshore Hudson Bay are probably driven by processes

254 operating within the Hudson Bay system itself. Large interannual variations are therefore unlikely, although regional variations are probably large, e.g., reflecting a generally deeper winter mixed layer in the northwest, compared to the southeast (Prinsenberg, re81).

The nitrate box model suggests that the major ultimate source of nitrate to the

Hudson Bay system is oxidation and nutrient regeneration within the Bay's own deep waters (Figure 6-1). This process supplies an estimated two-thirds of the nitrate necessary for production, while Foxe Basin/Hudson Strait inflow supplies most of the remainder and river inflow contributes only a minor amount. The dominant role for nutrient regeneration, relative to external supply, is consistent with the Bay's semi-enclosed state and likely limited import and export of OM (cf. lignin distributions in Chapter 3) and the apparently long residence time of the deep waters (-4-14 years, Pett and Roff, 1982).

Nutrient supply primarily through nutrient regeneration is characteristic of 'recycling' seas like the Baltic Sea and Norlh Sea (Chen et al., 2003) but uncommon among the

Arctic shelf seas, where inflows from the interior Arctic Ocean or from the Pacific or

Atlantic Oceans, generally provide the principal sources of nutrients for primary production (Chen et a1.,2003; Macdonald and Anderson, 2008).

The nitrate box model, which implies regeneration of about 9.5 gC m-2 a-r of marine OM annually in the deep waters of the Bay, therefore provides no evidence for low or incomplete nutrient regeneration processes in Hudson Bay, as previously proposed

(Legendre and Simard, 1979; Pett and Roff, 1982). Scarce and perhaps regionally-limited data (concentrated in the southeast) in previous studies may explain the different results, considering that the deep basin in the southeast may be partially isolated from the main

255 (western) Hudson Bay basin (Prinsenberg, 1986). Differences in vertical carbon flux

(Lapoussière et al., 2009) and possibly ventilation (Saucier et aI.,2004) may contribute to

large variation in deep water mass characteristics between the west and east sides of the

Bay.

Notably, the box model shows the reservoir of regenerated nutrients located

principally (>90%) in the offshore deep waters. Considering that only 30%o of the new

production occurs in the offshore, the nutrient reservoir in this area must partially result

from advective or convective processes, rather than local particle settling and oxidation

alone. If regeneration actually occurs inshore, then regenerated nutrients must be

transported seawatd, perhaps via dense water formation during winter (Figure 6-l).

Dense shelf overflow is an important process for transporting nutrients from the Beaufort

Sea shelf to the interior Arctic Ocean (cf. Melling and Lewis, 1982). Circumstances are

favourable for this process in western Hudson Bay (extensive flaw leads over shallow

shelves) and observations of high winter salinities (up to 33.8) near Churchill support

extensive fì'eezing in the flaw lead (cf. Chapter 2; Saucier et al., 2004); however, the

evidence for this process in Hudson Bay remains unclear.

Although the ôl5N and nitrate studies in the thesis emphasize the importance of upwelling rather than river inflow as the source of nitrate for new primary production in

Hudson Bay, it is important to note that river nitrogen inputs to Hudson Bay are

significant for production locally, e.g., in estuaries, and seasonally. Indeed, studies in the

Chut'chill River estuary (Chapter 2) showed that the river's nitrogenous nutrient supply in late winter/early spring contributes to more favourable conditions for primary production in estuaries, compared to the sunounding coastal (marine) areas. Similar conclusions

256 were reached in the Beaufort Sea, where inflow from the Mackenzie River alleviates

nitrate limitation of primary production in the inner shelf region (Carmack et al., 2004)

but supplies only a small fraction of the nutrients for primary production shelf-wide

(Chen et al., 2003; Macdonald and Anderson, 2008).

TERRIGENOUS OM: SOURCES, COMPOSITION AND DISTRIBUTION

Studies in the thesis applying bulk (CÀtr, ðr3C, õr5N¡ and specific organic

biomarkers (lignin) to characterize sediments from rivers and the seafloor provided the

first assessment of the nature of terrigenous OM delivered to and distributed within

Hudson Bay. Similar to other river-dominated Arctic coastal systems (cf. Fernandes and

Sicre,2000; Goni et a1.,2000,2005; Peulve eta1.,7996;Zegouaghet al., 1996), the terrigenous organic matter in Hudson Bay sediments and river inputs is comprised of a heterogeneous mixture. The ultimate origin of the terrigenous OM is C3 vascular plants; howevet, lignin compositional data indicate contributions from angiosperm and gymnosperm, woody and non-woody tissue sources (Chapter 3). The terrigenous OM inputs are also heterogeneous in terms of form and degree of degradation, reflecting variable mixtures of relatively undegraded plant debris and highly-degraded mineral- associated soil OM. Some of the cornpositional variation is organized regionally, with boreal forest (woody gymnosperm) vegetation an important source in the south, vs. tundra (non-woody angiosperm) in the north. Undegraded plant debris contributes more to the inputs of the relatively large rivers Q.Jelson and Hayes Rivers) draining the Hudson

Bay Lowlands, compared to rivers draining taiga or tundra regions in the southeast and north (Chapters 3, 4). The organic signatures suggest possible specific sources including peat, organic surface soil, deep mineral soil (likely eroded from river banks), and marine

257 clays or old fluvial deposits associated with the postglacial Tyrrell Sea. A highly- degraded, mineral-associated, non-woody angiosperm-derived material is a particularly prominent component of terrigenous OM throughout the Bay, with the pool of terrigenous OM in the sediments in the central basins almost exclusively comprised of this material. Although there are several possible sources of f,rne-grained sediments with this kind of signature, emerged Tyrrell Sea deposits, which surround southern Hudson

Bay, and emergent deposits in coastal areas are probably important sources. These inputs are significant in that they represent material likely to be highly resistant to further degradation in the marine environment and consequently more likely to be buried.

In addition to the regional watershed (vegetation) control on terrigenous OM composition, several processes operating within the marine environment place important controls on the quantity and composition of terrigenous OM distributed throughout

Hudson Bay. Striking compositional differences between the composition of the terrigenous OM presently delivered by rivers and the OM distributed in the Bay's sediments (Chapter 3) indicate that hydrodynamic sorting of terrestrial material in the coastal zone is a very important process (Figure 6-1). There is no evidence of such mis- match between contemporary source composition and composition of the sedimentary

OM sink on the Beaufort Sea shelf (Goni et a1.,2000, 2005). The sorting in the coastal zone appears to resuit in the preferential dispersal of fine-grained sediments enriched in highly-degraded, non-woody angiosperm organic matter and retention of most of the coarse-grained sediments and modern plant debris, especially of gymnosperm origin, close to the river sources. The fine component is transported primarily by marine curtents within Hudson Bay, with southem-derived material redistributed to the northeast and a

258 small fraction into west Hudson Strait by the Bay's general cyclonic circulation (Figure

6-1). The distribution patterns indicate that sea ice is probably not an important direct means of transport of terrigenous OM: there is no evidence for large-scale transport of terrigenous OM from the north to the south, which is the general direction of ice transport

(Markham, 1986). Localized ice rafting may in part account for the presence of coarse- grained plant materials more than 30 km from shore along the shallow southwest inner shelf (Chapter 3). Sediments in west Hudson Strait also contain kelp remains, which are interpreted as reflecting inputs of ice-rafted debris transported from Foxe Basin.

The sea ice cover indirectly influences the transport of terrigenous OM in Hudson

Bay by placing a control on the distribution of river runoff and associated suspended particulate matter during winter and spring (cf. Chapter 2). The nature of this control depends upon the spatial extent of landfast sea ice, the stability of the ice cover during winter and spring, the volume of river discharge, and the oceanographic conditions (e.g., tides). Regional differences in winter estuary structure and function in Hudson Bay probably contribute significantly to variations in the distribution of terrigenous (and possibly marine) OM. In areas with a nanow band of landfast sea ice bordered by rubble ice, such as the Churchill River estuary in western Hudson Bay, winter and spring river discharge may be partially impounded and diverted, forming an along-shore plume, which consequently deposits its suspended particulate load inshore of where it would be deposited in open water. Areas of relatively wide, stable landfast ice cover, which are common in southeast Hudson Bay, provide a different scenario in which the plumes that spread laterally under the ice cover are larger than those developed during open-water conditions. These conditions may promote non-directional dispersal of terrigenous OM

259 around river mouths. A third coastal ice regime, which occurs in places in western

Hudson Bay (e.g., around the Nelson River estuary), is characterizedby relatively unstable landfast ice cover, and consequently tends to export rather than retain winter river discharge and associated particulate matter. This winter export of river water may represent a second factor, in addition to ice rafting, explaining the wider distribution of coarse-grained plant materials along the southwest Hudson Bay shelf.

ROLE OF RESUSPENSION

Based on the studies in the thesis, one of the most imporlant processes controlling

OC cycling in Hudson Bay is resuspension and lateral transport of fìne,largely glacigenic, particles from coastal deposits and topographic highs and subsequent deposition in deeper sedimentary basins (Figure 6-1). The importance of the resuspension process in supplying sediments broadly to the Bay's sedimentary basins has not been previously well-recognized or quantified, although regional studies (Lavoie et al., 2008;

Zevenhuizen et al., 1994) proposed that redistribution of fine particles fi'om coastal deposits provides a principal sediment source for accumulation in local basins.

Henderson (1989) ploposed that erosion and remobilization of nearshore glacigenic deposits also contribute to postglacial sediment accumulation in central Hudson Bay.

Preliminary sediment and OC budgets developed here (Chapter 4) suggest that the resuspension process supplies about 80% (-l l6 x 106 t a-r) of the ca. 138 x 106 t ar of accumulating sediments in Hudson Bay.

Although the proposed resuspension term in the sediment budget (116 x 106 t a-l) is uncommonly large compared to shelf seas in the Arctic (Stein and Macdonald, 2004), or subarctic and temperate regions (e.g., Dunieu de Madron et al., 2000; Johannessen et

260 a1.,2003), the continuous relative sea-level fall in Hudson Bay (rates of almost 1.2 crnla in the southern part of the Bay; Peltier, 1998) is exceptional among Arctic areas. Falling

RSL represents a fundamentally different control on coastal geomorphology and sedimentation than the constant or slowly rising RSL characleúzing most other Arctic areas (cf. Beaulieu and Allard, 2003).In the northern Baltic Sea, resuspension of shallow water sediments is proposed to account for about 80% of the modern accumulating sediments (Håkanson et al., 2004; Jonsson, 1992).It therefore seems plausible that resuspension of shallow water sediments, supported on century to millennial time scales by falling RSL, represents the major source of accumulating sediments in Hudson Bay.

In addition to supplying the majority of the sediment for accumulation and burial, a second important role for resuspension in Hudson Bay's OC cycle is as a mechanism for redistributing and recycling terrigenous carbon. According to the sediment and organic carbon budgets constructed for the Bay (Chapter 4), about 0.58 x 106 t ar of terrestrial carbon is associated with the resuspension sediment supply, which represents as much terrigenous carbon as is supplied by river input and subaerial coastal erosion combined (Figure 6-2). Resuspension of emergent coastal deposits is therefore likely the most important source of the fine-grained, highly-degraded terigenous OM found widely distributed throughout the Bay's sediments.

The apparent resuspension supply would also deliver about 0.46 x 106 t a-r of marine carbon (Figure 6-2). This amount is insignificant compared to the estimated input from contemporary marine primary production in Hudson Bay (16.1 x 106 t C u-'); however, given that the coastal sediments undergoing resuspension are primarily old

(glacigenic), the associated ancient marine carbon would be highly degraded and

261 recalcitrant relative to freshly-produced marine OM, and thus resistant to further degradation during lateral transport. The resuspension supply of old marine carbon therefore probably explains in part the dominantly 'marine' character of the Hudson Bay sediments (-80%). Tenigenous OM is dominant in the sediments in most of the other

Arctic shelf seas (Gebhardt et al., 2005; Macdonald et al., 1998) and the central Arctic

Ocean (Schubert and Stein, 1997). Although some uncefiainty in source apportionments is unavoidable because organic tracer data have not been previously used in the area, the use of a combination of traditional tools for distinguishing OC sources (e.g., C/lrl, ðr3C¡ and more specific proxies (e.g., lignin) provides greater confidence in the accuracy of the derived OM fractions than could have been obtained from individual organic geochemical parameters alone (cf. Eglinton and Repeta, 2004; Schubert and Calvert,

200I; Schubert and Stein, 1996). Furthermore, appropriate end-member values for marine and terrigenous sources, which are required to quantify the relative contributions of these two sources, \¡r'ere estimated primarily from site-specific measurements in

Hudson Bay rivers and the marine water column.

SEDIMENT ACCUMULATION AND OM BURIAL

Arctic sediments generally provide efficient sites for OM burial (Chen et al.,

2003; Stein and Macdonald ,2004) and because Hudson Bay is an inland sea with relatively closed boundaries, accumulation on the seabed was expected to be a primary sink for sediments and POM produced within or delivered to the system (Leslie, 1963;

Pelletier, 1986). This is indeed the case, with overall sediment burial estimated at -138 x

106 t a-1, which compares well with aveïage Holocene sediment burial rates of between

100 x 106 t a-r and,200 x 106 t a-l in the Beaufort, East Siberian and Kara Seas (Stein and

262 Macdonald , 2004). However, consistent with the active resuspension process described above, the sedimentation patterns in Hudson Bay are complex and OM burial is

2lOPb discontinuously distributed across the seafloor. Sedimentation rates derived from profiles in 15 sediment cores (Chapter 4) varied widely (0.032-0.23 g cm-2 a-r), consistent with a regional sedimentation map developed from seismic data (Josenhans et al. (1988) and Geological Survey of Canada Open File Report#2215), showing modern sediments generally restricted to scattered pockets and localized depressions on the Bay's pseudo- shelves, while glacial till and glaciomarine sediments remain exposed at the seafloor surface throughout much of the Bay's interior (see Figure 4-5). The widespread distribution of glacigenic sediments in Hudson Bay (Josenhans et al., 1988) necessitated very careful interpretation of sediment core records (see Chapter 4) because under conditions of low sedimentation, sediment mixing processes become an increasingly

2lOPb important control on the distribution of and other tracers (cf. Anderson et al., 1988; de Haas et al., 1997,2002).

The overall burial rates for marine and tenestrial OC in Hudson Bay derived from

I the estimated sediment sink, and OC content and composition data, are 1.03 x 106 t C a

and,0.23 x 10ó t C a-r, respectively (Figure 6-2). These rates fall within the range for the

Arctic Ocean marginal seas (0.23 x 106 t C a-r to 2.8 x 106 t C a-r; Stein and Macdonald,

2004). As mentioned, the available organic proxy data indicate that the OM is predominantly of marine origin. Terrigenous OM contributes less than 20o/o of the lrral sedimentary OM load in most of the Bay, with slightly higher contributions (30-40%) in the southeastern paft of the Bay (where rivel inflow is concentrated; Dery et al., 2005) and a contribution of up to 50%o in the southwest. These data provide a baseline for

263 detecting possible future changes in OM burial and attributing changes to alterations in processes affecting either marine and terrigenous sources.

Proportions of the marine and terrigenous carbon inputs lost to oxidation or leaching (estimated as the difference between input and burial; Figure 6-2) indicate that marine carbon is eff,rciently oxidized in the Hudson Bay water column (>93%). Slightly higher proportions (95% ->99%) of marine carbon recycling have been estimated for the

Beaufort Sea (Goni et a1.,2005; Macdonald et al., 1998), productive areas of the arctic

Alaskan shelf Q.{aidu et al., 2004) and the central Kara Sea (Gebhardt et al., 2005). The budget therefore supports the idea that the resuspension supply of marine carbon contributes to slightly greater preservation of marine carbon in Hudson Bay than in other

Arctic coastal seas, where the marine OM inputs are entirely modern. The apparent leaching or oxidation of terrestrial OM in Hudson Bay (-65%) seems large considering the highly degraded state of the terrigenous OM delivered by Hudson Bay rivers and associated with the resuspension sediment supply. However, a roughly 65% loss seems reasonable in the context of the 80-90% loss of modern terrigenous OM and 650/o loss of ancient temigenous OM recently observed on the Beaufort Shelf (Goni et al., 2005).

Degradation of old OM in Hudson Bay may be enhanced by repeated cycles of resuspension and deposition during transport to offshore sinks (Figure 6-1, Blair et al.,

2004; Hulthe et al., 1998). Progressive degradation of sedimentary lignin from the margin to the interior in Hudson Bay (Chapter 3) supports significant degradation of terrigenous material during transport to offshore sinks.

The significance of resuspension in the OC cycling of coastal seas remains a subject of some debate, with some studies suggesting that this process is insignificant

264 because it does not add to the burial of new OM but simply transfers already-present OM from one place to another (de Haas et al., 1997),but other studies suggesting that the process may be significant because it relocates nutrients (Jonsson et al., 1990), or perhaps interacts with and influences the deposition of modern OM (O'Brien et a1., 2006).It has been proposed that Arctic ecosystems seasonally shift from an algal to a detrital mode, in which they may actually rely on old organic carbon associated with resuspended materials (Forest et al., 2008). Extensive removal of terrigenous OC from suspended particles and partial replacement with marine carbon at the river-ocean interface (Hedges et al., 1997) at least supports possible interactions between the resuspension process and the cycling of modern OM. The apparent extensive degradation of terrestrial OM, including some of that associated with the resuspension sediment supply (Figure 6-2), implies that the resuspension process does play a role in active OC cycling in Hudson

Bay and does not simply represent a passive transfer.

IMPLICATIONS FOR HUDSON BAY'S RESPONSES TO CLIMATB CHANGB

The findings of the studies in the thesis imply important differences in terigenous

OM sources and transport pathways, coastal processes, nutrient sources and primary production, and sedimentary regime between Hudson Bay and marginal seas of the Arctic

Ocean. These differences will need to be clearly understood before using Hudson Bay as a sentinel for change in the Arctic's marine environment.

The proposal that new production is concentrated in inshore surface waters of

Hudson Bay because of enhanced nitrate supply due (in part) to river inflow and entrainment, represents a new perspective on primary production and its controls in

Hudson Bay. Clearly, additional dataneed to be collected to properly assess the relative

265 contributions of nutrient entrainment related to river inflow vs. tidal- and wind-driven vertical mixing. Nevertheless, one may speculatethat changes in the volume or timing of river discharge related to previously-constructed dams and diversions (Prinsenberg, 1980) or ongoing climate change (Dery et al., 2005) may have leverage to adversely affect primary production, especially in inshore regions of Hudson Bay. The relatively minor direct role of river inflow as a nutrient source in Hudson Bay means that changes in river- water nutrient concentrations, which may come about due to changes in the watershed

(Rouse et al.,1997), will be less important for the Bay's total annual production.

It has been proposed that the most important control on marine primary production on interior Arctic Ocean shelves in the future will be upwelling of interior

Arctic Ocean water across the shelf break, which depends on sea ice distribution which, in turn, reflects atmospheric forcing (Carmack et al., 2006; Macdonald et al., 1987).

Although a major change in atmospherically-forced seawater exchanges (e.g., increased inflow from the Atlantic Ocean) would dramatically alter Hudson Bay's nutrient regime and consequently arurual production, variations in timing and volume of river discharge, local winds, and processes within the Bay, such as vertical mixing and nutrient transport by dense shelf overflows, will otherwise govem future production. The sedimentary record from the last few centuries (including the Little Ice Age) may contain clues as to the susceptibility of the system to major shifts in seawater exchange. Reductions in sea ice cover (Gagnon and Gough,2005) and consequent increasing wind-driven nutrient upwelling may enhance primary production in the inshore region in spring. Reductions in sea ice may also affect ploduction in the offshore region either positively, by virtue of decreased melt (reduced stratification), or negatively, through decreased winter mixing

266 and consequently a smaller nutrient pool in spring. Changes in the sea ice cover in coastal areas may also influence the Bay's overall nutrient supply by altering the winter buoyancy flux from estuaries and consequently where (or if.¡ dense shelf water formation occurs and redistributes regenerated nutrients. The instability of the ice cover in western

Hudson Bay implies that the estuaries along this coast may be particularly responsive to change, which is especially important if western Hudson Bay is presently a site of dense water formation (Saucier et a1.,2004).

Terrigenous OM inputs, composition and distribution in Hudson Bay all seem to have considerable capacity for change over time, perhaps greater than in other northern areas. The studies suggest that there are distinct reservoirs of terrigenous OM with different signatures in the Hudson Bay watershed; mobilization of these different materials will be affected by changes in the relative importance of surface vs. subsurface or river-bank erosion, which in turn depends on river ice conditions and the timing and quantity of runoff (cf. Guo and Macdonald, 2006). Many small rivers supply the river input to Hudson Bay, rathel than the large rivers characteristic of the interior Arctic

Ocean shelf seas, and the numerous rivers seem likely to respond in various ways to climate change (e.g., increased discharge in the north, decreased discharge in the southwest, Dery et al., 2005). The estuaries of small rivers may also be more variable and unstable during the winter than the estuaries of larger rivers, which implies they may respond more rapidly or to a greater degree to changes in Arctic ice regime. Together, these factors will make it more difficult to interpret future terigenous organic compositional changes in Hudson Bay than in other Arctic areas and attribute them to changes in permafrost, watershed hydrology, vegetation, ice-processes, etc. There is

267 clearly also capacity for major changes in terrigenous OM composition as a result of

altered processes within the Hudson Bay marine system itself. The importance of

hydrodynamic sorting for separating coarse and fine components of modem river inputs,

and the apparent large role of resuspension and lateral transport of coastal deposits, imply

that terrigenous OM distribution could be strongly altered by the increased storminess

and wave-base erosion that might result from delayed fall freeze-up. For other Arctic

shelves, the largest changes in sediment and terrigenous POC supply are expected to

come from accelerated subaerial coastal erosion, e.g., thawing and erosion of ice-bonded

shoreline cliffs, enhanced by ongoing relative sea-level rise and reduced sea ice cover

(Carmack et al., 2006). Sediment and organic materials moved about by resuspension

likely differ in particle size, OM composition and degree of degradation, etc. from those

supplied by subaerial coastal erosion. Therefore, increased inputs from these sources

would have different consequences for the regional OC cycles (e.g., enhancing

heterotrophic metabolism vs. enhancing burial). This emphasizes the importance of an in-

depth understanding of regional sources and OM composition and the specific processes

contributing to supply and transport.

Based on the studies in the thesis, it seems likely that sediment and OC supply

and burial are not at steady-state in Hudson Bay on century or millennial time-scales. In general, it takes a long time for geologic and geomorphologic processes to recover from deglaciation (Ballantyne,2002) and consequently for biogeochemical cycles to reach steady-state (cf. Der'ry ,2009). Continued relative sea-level fall like that occurring in

Hudson Bay represents a continued perturbation, which alters the timeline and even the trajectory of recovery (Ballantyne,2002). This kind of 'transitional' sedimentary and OC

268 regime will present a unique challenge for efforts to detect and interpret Hudson Bay's

responses to future climate change. Furthermore, changes in particulate transport likely

lag isostatic rebound and thus sediment core records may reflect a kind of time-lagged

compositional shift (i.e., compositional changes propagating slowly fi'om coastal reservoirs to central basin sinks). The time-scale of major processes controlling Hudson

Bay's OC cycle, such as erosion of coastal sediments and deposition and resuspension

(Figure 6-i), clearly needs to be better understood. For instance, it is possible that changes to present day inputs - e.g., increased inputs of plant debris related to accelerated permafrost degradation - would take decades (or longer) to detect in the organic composition of accumulating sediments. The advances in our understanding of OC cycling in Hudson Bay through the studies in this thesis represent an important first step, albeit down a long path, toward the in-depth understanding that will be needed to predict and detect the effects of climate change on the OC cycle of this complex Arctic coastal sea system.

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