Paleoecological and Carbon Accumulation Dynamics of a Fen Peatland in the Hudson Bay Lowlands, Northern Ontario, from the Mid-Holocene to Present

Benjamin Cody O’Reilly

A thesis submitted in conformity with the requirements for the degree of Master of Science Department of Geography University of Toronto

Paleoecological and Carbon Accumulation Dynamics of a Fen Peatland in the Hudson Bay Lowlands, Northern Ontario, from the Mid-Holocene to Present

Benjamin O’Reilly

Master of Science

Department of Geography University of Toronto

2011 Abstract

Pollen assemblages, peat humification and carbon:nitrogen stratigraphy were examined at high resolution in a core from a fen peatland in the Hudson Bay Lowlands, Northern Ontario, to interpret the factors that drive long-term peatland dynamics. Subtle changes in the vegetation community are evident over the record, suggesting both allogenic and autogenic influences, but a fen community appears to have been resilient to external perturbations including isostatic rebound and hydroclimatic changes between 6400 and 100 years BP. Paleoclimatic reconstructions from the fossil pollen assemblages indicate that precipitation increased 3000 years BP at the end of the Holocene Thermal Maximum, and that carbon accumulation in the fen was controlled more by effective surface moisture (precipitation) than by temperature. The pollen record suggests changes over the past century, including increases in shrub Betula, Alnus,

Ambrosia, and Cyperaceae and a decrease in Sphagnum spores, consistent with the observed

Pan-Arctic shrub increase.

Acknowledgments

I would like to start by thanking Dr Sarah Finkelstein for the guidance and assistance that she afforded me throughout my Masters. She was constantly keeping me motivated and excited about the next step, and was always so encouraging. I am very grateful for the opportunity that was given to me to work on this project.

This project relied substantially on the funding and support of a number of sources. I wish to thank the Ontario Ministry of Natural Resources, the Ontario Ministry of Training,

Colleges and Universities, the Natural Sciences and Engineering Research Council of Canada, the Wildlife Conservation Society of Canada and the Northern Scientific Training Program of the

Department of Indian Affairs and Northern Development. Once the roads end, the cost of doing field research really takes off, and it couldn’t be done without the generosity of these sources.

The logistical support provided by Brian Steinback and the rest of the staff at DeBeers

Victor Mine Environmental Lab is greatly appreciated. The stay at Victor was memorable, and I hope my torn pants were a lesson in proper field attire (or at the very least, a lesson in writing a proper near-miss card). I must admit, very few things cap off a day of walking around expansive muskeg like pulling a truck, so thanks for the staff at Victor for making us feel welcome!

HMS PGB would never have sailed without the careful construction of Mircea Pilaf.

Thank you for all your help over these two years Mircea! I also wish to thank Jim McLaughlin and Benoit Hamel for the core collection and supplemental site description.

To the others in the Paleoecology Lab – Carlos, John-Paul, Charlotte, Joan, Maara,

Kristen and Nikki and those already moved on – Jane and Jen, thanks for all the coffee breaks, patio beers, rants, discussions, assistance and good times. I owe you all a lot for the motivation

iii you afforded me, and for not laughing at my jokes resulting in me thinking of better ones! A special thanks to Kristen for helping sub-sample peat when the temperatures of the sediment lab approached solar-surface levels, and Joan for patiently sharing her vast knowledge of statistics with me.

I would also like to thank Charlie and Jock for the visits, interesting conversations and helpful suggestions.

I really need to thank my parents twice, mainly because I forgot to thank them in my undergraduate thesis acknowledgements, but more so because they encouraged me to take this opportunity and have been more supportive than I could have ever dreamed. I hope I can repay their kindness and goodwill!

To the rest of the folks of PGB, thanks for making movie nights, Fridays, Chinese New

Year and other events memorable. It really helped get through the tough parts of graduate school, and I will cherish this time forever.

Lastly I’d like to thank my girlfriend Tatiana for all her love and support during this time in my life. Thanks for all your encouragement and motivation, especially when I was at my grumpiest! I still think the Washington Redskins are better than the Philadelphia Eagles though!

Table of Contents

TITLE PAGE………………………………………………………………………………………i

ABSTRACT………………………………………………………………………………………ii

ACKNOWLEDGEMENTS………………………………………………………………………iii

TABLE OF CONTENTS………………………………………………………………………….v

LIST OF TABLES………………………………………………………………………………viii

LIST OF FIGURES………………………………………………………………………………ix

LIST OF APPENDICES………………………………………………………………………….xi

CHAPTER 1: INTRODUCTION……………………………………………………...………….1

1.1 GENERAL INTRODUCTION AND OBJECTIVES………………………………...1

1.1.1 Development of Northern Peatlands………………………………………2

1.1.2 Rationale…………………………………………………………………..4

1.1.3 Proxies Utilized and their interpretation…………………………………..5

1.1.4 Ecosystem Resilience…………………...…………………………………9

1.1.5 Peatlands as Complex Adaptive Systems…………………………………9

1.2 LITERATURE REVIEW……………………………………………………………12

1.2.1 Holocene Climatic Transitions…………………………………………...12

1.2.2 Past Paleoecological Studies……………………………………………..13

1.2.3 Carbon Accumulation in Peatlands………………………………………20

1.3 STUDY SITE………………………………………………………………………...23

1.3.1 Study Region……………………………………………………………..23

1.3.2 Site Description…………………………………………………………..24

1.3.3 Climate of the study area………………………………………………...25

1.3.4 Local and Regional Geologic Setting……………………………………27

1.3.5 Quaternary Glacial History of the Hudson Bay Lowlands………………27

1.3.6 Post-glacial Isostatic Adjustment………………………………………...29

1.3.7 Local and Regional Vegetation…………………………………………..31

CHAPTER 2: METHODS……………………………………………………………………….34

2.1 FIELD METHODS…………………………………………………………………..34

2.2 LABORATORY METHODS………………………………………………………..35

CHAPTER 3: RESULTS………………………………………………………………………...44

3.1 210Pb DATING OF VICM_T3_SP3………………………………………………….44

3.2 AGE-DEPTH MODEL DEVELOPMENT………………………………………….45

3.3 PALEOECOLOGICAL RECONSTRUCTION…………………………………….49

3.4 BULK DENSITY……………………………………………………………………54

3.5 C:N STRATIGRAPHY……………………………………………………………...54

3.6 LORCA………………………………………………………………………………59

3.7 PEAT HUMIFICATION…………………………………………………………….62

3.8 PALEOCLIMATIC RECONSTRUCTIONS………………………………………..65

CHAPTER 4: DISCUSSION……………………………………………………………………76

4.1 DRIVERS OF VEGETATION CHANGE…………………………………………..76

4.2 CLIMATE RECONSTRUCTION…………………………………………………...83

4.3 CONTROLS ON CARBON ACCUMULATION DYNAMICS……………………89

4.4 RESILIENCE OF THE VICTOR FEN ECOSYSTEM……………………………...92

CHAPTER 5: CONCLUSION…………………………………………………………………..95

5.1 CONCLUSIONS FROM THE VICTOR FEN RECORD..………………………….95

5.2 FUTURE WORK…………………………………………………………………….97

REFERENCES…………………………………………………………………………………..99

APPENDIX A: RAW COUNTS OF VC01…………………………………………………….112

vii

List of Tables

Table 1: The three study proxies and their interpreted reconstruction…………………………..11

Table 2: Grain counts for the rationale of a 200 arboreal pollen grain count……………………40

Table 3: AMS radiocarbon dates for the Victor Mine Fen Core (VICM_T3_SP3)……………..48

Table 4: Percent carbon and nitrogen data used to test the homogeneity of the peat matrix…….56

viii

List of Figures

Figure 1: A map of past paleoecological studies in relation to the Victor fen…………………..15

Figure 2: Postglacial emergence curves for the Victor fen site………………………………….31

Figure 3: The activity of 210Pb in the uppermost Victor fen core section………………………..45

Figure 4: Age-depth model derived for the Victor fen Core……………………………………..47

Figure 5: Percentage pollen diagram from Victor fen core……………………………………...52

Figure 6: Pollen Influx diagram for the Victor fen core…………………………………………53

Figure 7: Bulk density of the Victor fen core……………………………………………………54

Figure 8: Percentage carbon in the peat sequence of the Victor fen core………………………..57

Figure 9: Percentage nitrogen in the peat sequence of the Victor fen core……………………...58

Figure 10: Carbon/Nitrogen ratio of the peat sequence of the Victor fen core………………….59

Figure 11: LORCA estimates for the entire peat sequence of the Victor fen core………………60

Figure 12: LORCA estimates for the 60 cm to base section of the Victor fen core……………..61

Figure 13: LORCA estimates for the Victor fen core based on the age-depth model…………...61

Figure 14: Raw spectrophotometric absorbance results for the Victor fen core…………………64

Figure 15: Detrended absorbance values (Ad) for the Victor fen core…………………………...64

Figure 16: Reconstructed Average Annual Air Temperature for the Victor fen core…………...68

Figure 17: Reconstructed Average Annual Air Temperature of the most recent 2000 years for the Victor fen core…………………………………………………………………………………...69

Figure 18: Reconstructed Average July Temperature for the Victor fen core…………………..70

Figure 19: Reconstructed Average July Temperature of the most recent 2000 years for the Victor fen core…………………………………………………………………………………………...71

Figure 20: Reconstructed Total Annual Precipitation for the Victor fen core…………………...72

Figure 21: Reconstructed Total Annual Precipitation of the most recent 2000 years for the Victor fen core…………………………………………………………………………………………...73

Figure 22: Reconstructed Total June, July, August Average Precipitation for the Victor fen core……………………………………………………………………………………………….74

Figure 23: Reconstructed Total June, July, August Average Precipitation of the most recent 2000 years for the Victor fen core……………………………………………………………………..75

List of Appendices

Appendix A: VC01 raw pollen counts………………………………………………………….112

Chapter 1 INTRODUCTION 1.1 General Introduction and Objectives

Peatlands store more carbon per unit area than any other terrestrial ecosystem (Dise

2009). However, most peatlands are located in the Boreal and Subarctic zones of the

Northern Hemisphere, where the climate has been warming faster than anywhere else on

Earth; this is a trend which is projected to continue (Meehl et al. 2007). An alarming consequence of global climate change is a decreased ability of certain ecosystems to uptake and store carbon. A decrease in carbon storage acts as a positive feedback to global climate, accelerating warming during periods of carbon release, and this warming is expected to impact peatland carbon cycling (Beaulieu-Audy et al. 2009). High- resolution paleo-data retrieved from peat repositories indicate that the carbon sink potential of northern peatlands has varied by an order of magnitude or more in past millennia (Yu 2006), in response to hydroclimatic change.

Northern peatlands span an area of approximately 4 000 000 km2 and are thought to contain a carbon pool of between 270 to upwards of 547-621 Gt of carbon (C), more than one third of the world’s soil carbon (Beilman et al. 2009; Frolking et al. 2010;

Gorham 1991; Turunen et al. 2002; Yu et al. 2010). For regions with permafrost

(including continuous, discontinuous, sporadic and isolated zones), peat soils including

Histels (perennially frozen peatland soils) and Histosols (unfrozen peatland soils) are estimated to contain 114.5 Pg and 227.3 Pg of soil organic carbon for North America and the total Northern Hemisphere respectively (Tarnocai et al. 2009). The wide variation in estimates of the total carbon pool supports the role of the paleo-record as an integral

1 2 element in quantifying future carbon storage capacity of northern peatlands under projected hydroclimatic conditions.

1.1.1 Development of Northern Peatlands

The limited number of basal peat radiocarbon dates before circa 16 500 years BP (based on 1516 basal radiocarbon dates of peat initiation from high latitude Europe, Asia and

North America) suggests that there were none of the extensive peatland complexes (West

Siberian Lowland, Hudson Bay Lowland (HBL)) that characterize the modern northern circumpolar region during the Last Glacial Maximum (MacDonald et al. 2006). This is supported by the near absence of Sphagnum spores in peat deposits from 16 500 years BP

(Gajewski et al. 2001).

Sphagnum peatlands developed soon after deglaciation (15-11 000 years BP) in

North America, with initiation beginning in Alaska and the St. Lawrence regions, spreading eastward and westward respectively, in response to newly colonisable land

(Gajewski et al. 2001; MacDonald et al. 2006). The arrival of early Holocene warming at

11 500 years BP immediately following the Younger Dryas cold event is characterized by a rapid expansion of peatlands throughout the north (MacDonald et al. 2006). Carbon accumulation rates also peak at approximately 25 g C m-2 year-1 in the early Holocene

(between 11 000 and 9000 years BP, based on 33 northern peatland sites), concurrent with the peak in peatland initiation (MacDonald et al. 2006; Yu et al. 2010). However,

Yu et al. (2010) use the previous northern peatland radiocarbon synthesis of MacDonald et al. (2006) with very poor spatial coverage of initiation dates in the HBL. The HBL represents an important gap in the complete spatial coverage of northern peatland

3 initiation (Gorham et al. 2007), especially because peatland initiation in the HBL took place after 7000 years BP.

The North American initiation findings are consistent with rates from Alaska with the highest rate of peatland development (based on 284 basal peat dates) from 12 000 to 8

600 years BP (peak at 10 500 years BP) (Jones and Yu 2010). The rate of additional peatland development in North America was constrained by the activity of ice retreat and land exposure. In boreal North America, major development occurred after 9000 BP, in response to the retreating Laurentide ice (Gajewski et al. 2001; MacDonald et al. 2006).

The impact on today’s atmosphere due to the establishment and growth of northern peatlands over the Holocene is that of a carbon sink (net deficit) of between 40-80 GtC

CO2 (20-40 ppmv) and a source (net increase) of approximately 0.2 to 0.4 GtC CH4 (75-

150 ppbv) as northern peatlands accumulate carbon (as the vegetation uptakes CO2) and emit CH4 through microbial production under anaerobic conditions (Frolking and Roulet

2007; Klinger et al. 1994). These northern peatlands have resulted in a radiative forcing cooling impact of -0.2 to -0.5 Wm-2. However, early in the Holocene the radiative forcing impact would have been a net warming of 0.1 Wm-2 (Frolking and Roulet 2007).

The Hudson Bay Lowlands (HBL) of northern Ontario is the second largest peatland complex in the northern Hemisphere, after the West Siberian Lowland, and has been a significant contributor to the overall carbon pool that has accumulated in Northern peatlands during the post-glacial period (Beilman et al. 2009; Gorham 1991; Martini

2006). Due to the remoteness of the HBL, few stratigraphic reconstructions or carbon accumulation studies have been undertaken, pointing to a lack of understanding of the

Holocene dynamics of this large peatland basin. The ability to accurately reconstruct past

4 environments is necessary to understand the dynamics of Earth systems, and to test models used to predict future hydroclimatological changes (Belyea 2007). Given the uncertain estimates of the total carbon pool in northern peatlands, and how these systems have responded to hydroclimatic change in the Holocene, high resolution analysis of vegetation change through pollen analysis coupled with estimates of carbon accumulation in the peat deposits of the HBL are crucial.

1.1.2 Rationale

The objectives of this study were to reconstruct vegetation change and carbon storage in wetlands of the Attawapiskat River basin of the HBL and integrate these data sets with hydroclimatic changes inferred from paleoclimatic reconstructions. There remains considerable variability in the estimates of the carbon pool of northern peatlands. The variability in these estimates is due to possible inaccuracies in the basal dates of peat sequences as well as assumptions of average peat depth, average bulk density of peat and the proportion of carbon in peat (Gorham 1991; Turunen et al. 2002). Constraining these variables is important and this study aims to accurately characterize a poorly known region to refine estimates of the carbon pool, and how the pool has responded to climatic variability in the past. These objectives will be met by studying the paleoecological, paleohydrological and geochemical records retrievable from wetland sediments of the

HBL.

The records that were intensively studied include pollen assemblages isolated from the peat sediment, spectrophotometric humification of the peat matrix (the amount of humic acids at a given depth in the peat) and carbon:nitrogen ratios of the peat matrix.

The ability of peatlands to accumulate autochthonous (originating at the site) material in

5 a sequential order, to sequester carbon as peat for many thousands of years, and to contain a very detailed record of changes in local to regional vegetation makes peatlands useful for investigating environmental and climate changes over Holocene or longer timescales (Chambers and Charman 2004). The autochthonous process of peat accumulation also makes peatlands less susceptible to redeposition, which is more common in lake sediment sequences and can impact stratigraphic results (Chambers and

Charman 2004).

The ultimate goal of this research was to characterize the effects of climatic

(temperature and precipitation) and elevation (isostatic uplift) changes on vegetation communities and carbon accumulation in peat deposits of the HBL, and integrate these analyses with estimates of peat and carbon accumulation (Clymo 1984; Yu et al. 2003).

Multiple paleoecological and paleohydrological techniques were employed to get a holistic picture of the history of climate change and carbon accumulation in the fen peatland. A multi-proxy approach is used to avoid erroneous interpretations from a single proxy, resulting in more robust reconstructions (Blundell and Barber 2005).

1.1.3 Proxies Utilized and their Interpretation

Pollen assemblages will be used to highlight and separate the influence of the allogenic

(hydroclimatological) and autogenic (local biotic processes) factors on the carbon accumulation of each site. As ecotones, wetlands usually respond strongly to allogenic forcing (Mitsch and Gosselink 2007), highlighting the need to use proxies that can separate the biotic and abiotic drivers. Fossil pollen assemblages can be used to construct quantitative estimates of past environments using the modern analog technique (MAT)

(Williams and Shuman 2008). The MAT is an established, robust procedure that assists

6 in the reconstruction of past climates and vegetation from quaternary fossil pollen assemblages when combined with modern, spatially extensive calibration datasets

(Jackson and Williams 2004; Overpeck et al. 1985; Williams and Shuman 2008).

The climate reconstructions from the pollen assemblages of the Victor fen will provide important paleoclimatic information for the study area, where it is lacking. The reconstructions chosen were average annual temperature (°C), mean July temperature

(°C), total annual precipitation (mm), and average June, July, August (JJA) precipitation

(mm). Each of these climatic values was chosen for a specific reason. Rates of Carbon sequestration in peatlands depend on the ambient hydroclimatic conditions (Belyea and

Malmer 2004). The temperature values were chosen because the addition of carbon at the top of the acrotelm reflects the imbalance of fixation and aerobic decay; this relationship is affected by surface temperature of the peatland (Clymo et al. 1998). Precipitation values were chosen because past work has shown that carbon accumulation in fen peatlands responds strongly to even small changes in moisture conditions even if no change in dominant species is found in a paleoecological reconstruction (Yu et al. 2003).

A “summer” subset of both temperature and precipitation was used because of the continental climate of the site. The winter period (being moist and cold at mid to high latitude) has been deemed less important to long-term surface wetness changes in mires

(peatlands), with the exception of snow melt input in the spring possibly extending the season of surface saturation (Charman et al. 2009). Thus, precipitation reconstructed for the summer season was important. Also, given that humification values are surface humidity dependent and therefore, can exhibit a temperature or moisture signal, both reconstructions were necessary.

Studies on the mechanisms governing the vegetation dynamics of wetlands in the

Holocene, through the analysis of pollen assemblages, have indicated that responses to climate-induced hydrological changes (allogenic) and within-wetland species change

(autogenic) combine to facilitate succession (Singer et al. 1996; Winkler 1988). For example, moisture changes to the Portage Marsh basin (Indiana, USA), especially the transition from open-shallow lake to marsh, were coincident with changes in upland vegetation suggesting climate is the dominant mechanism driving the evolution from lake to marsh at that site (Singer et al. 1996). However, the progressive shallowing of the basin by the accumulation of autochthonous sediment has dampened to some extent the responses to climatic change, showing that both allogenic and autogenic influences determine wetland dynamics at this site (Singer et al. 1996).

In Washburn and Hook Lake bogs of south-central Wisconsin, the major hydrological and aquatic vegetation changes were synchronous after 6500 years BP with the change to a dry-warm climate as shown through upland vegetation changes indicative of regional warming resulting in a lowering of the water table at both bog sites (Winkler

1988). A later transition to Sphagnum occurred at both sites, and the growth of established Sphagnum has been found to intensify the acidification process (Glaser et al.

1981) resulting in a greater influence of autogenic forcing. The synchronicity of changes points to climate being an important factor in influencing hydroseral change (sequence of ecological communities at a saturated site) in wetland ecosystems (Winkler 1988). The complex nature of the combination of allogenic and autogenic factors acting to force vegetation succession in peatlands necessitates proxies sensitive to both factors and this makes pollen analysis of the peat deposit useful.

Ratios of carbon:nitrogen will be combined with the pollen stratigraphy to accurately assign the various development stages (fen versus bog) in the cores, to assess the degree of decomposition and to calculate the carbon accumulation rate of each peatland development stage. The degree of peat decomposition (an analog for moisture) at each portion of the core is estimated through spectrophotometric measurement of peat humification. Measuring the absorbance of an alkaline extract of dried peat returns a result proportional to the amount of humic matter dissolved, with less absorbance indicating less humified peat (Aaby and Tauber 1975). A trend towards less humified peat (lighter coloured) suggests increasing mire surface humidity, which can be due to higher water table position and/or a more positive surface moisture balance, driven by either higher precipitation or lower temperature, or a combination of the two factors

(Aaby 1976).

At levels where subsamples for both humification and C:N ratios are possible, the correlation between the two variables will be assessed to determine how accurately the

C:N ratios capture the decay signal. Given past work indicating that high N proportions and low C:N ratios are indicative of greater peat decay, the correlation is expected to be high (Belyea and Warner 1996; Borgmark and Schoning 2006; van der Linden and van

Geel 2006). Table 1 is a summary of each proxy studied, the influencing factors acting on each proxy, and variables reconstructed by each proxy. As indicated by Loisel and

Garneau (2010), the purpose of utilizing a suite of proxies is to attempt to isolate the mechanisms that drive peatland development, which include for the Hudson Bay

Lowlands isostatic uplift, hydroclimatological variability and autogenic successional processes.

1.1.4 Ecosystem Resilience

Past studies in the HBL have indicated that at many locations, a transition from a fen- type ecosystem to a bog takes place over a long period of time. In light of this observation, the study of a long-lasting fen ecosystem provides a useful test of ecosystem resilience. Ecosystems are considered resilient when ecological interactions combine to strengthen one another and reduce disruptions (Peterson et al. 1998). This resilience denotes the maximum perturbation that can be “absorbed” by the ecosystem without causing it to shift to an alternate stable state (Scheffer et al. 2001). It has been defined as the capacity of a system to absorb a disturbance and reorganize while changing to retain the same structure, function, identity and feedbacks (Folke et al. 2004). The combination of proxies that capture vegetation and climate signals will aid in testing whether or not the fen is a true resilient ecosystem.

1.1.5 Peatlands as Complex Adaptive Systems

Recently, peatlands have begun to be treated conceptually as complex adaptive systems

(CAS) due to the important scale-transcending spatial and temporal linkages between the relatively fast near-surface processes and the slower processes occurring deeper in the deposits (Belyea and Baird 2006). The general properties of CAS that peatlands exhibit are spatial heterogeneity, localized flows, a self-organizing structure and non-linearity

(Belyea and Baird 2006). The internal peatland dynamics and external forcing mechanisms both act to cause variability in hydroclimatological conditions and micro- relief patterns, and the allogenic and autogenic forcings impact hydrological conditions influencing peatland carbon cycling and development (Belyea and Baird 2006).

The Victor fen exhibits characteristics of peatlands as complex adaptive systems as defined by Belyea and Baird (2006). The surface of the peatland being at or near the depth of the water table differs from the surrounding peatland micro-relief in hydro- physical and ecological characteristics. The water table being so close to the surface would influence the peat accumulation rate, the redox conditions in the acrotelm, and the local vegetation that could thrive under these conditions. This feature represents the spatial heterogeneity component of a complex adaptive system. The minerotrophic input represents the localized flow feature of a complex adaptive system, with the litter and peat layers interacting through the flow of water and nutrients (Belyea and Baird 2006).

The size and shape of peatlands constrain processes operating at smaller scales.

Thus, the peatland as a whole would influence the ecology and hydrology of the fen throughout its existence. This influence is referred to as the self-organizing structure component of complex adaptive systems (Belyea and Baird 2006). Lastly, the hydrological conditions at the surface of the peatland change with external forcing and varying minerotrophic inputs, and the changing conditions would constrain surface structure and composition as well as peat accumulation rates. This is the non-linearity component of a complex adaptive system (Belyea and Baird 2006).

The consistent theme of a coupling of allogenic and autogenic influences acting on peatland dynamics provides further support for methods capable of separating the two dominant mechanisms. Establishing the relationship between hydroclimatic conditions and carbon dynamics is important because high LORCA (long term apparent rate of carbon accumulation, found by dividing the accumulated mass of C in a peat deposit by the age of the basal peat) (Korhola et al. 1995; Tolonen and Turunen 1996) values have

been found to be correlated both with wet or dry conditions (Loisel and Garneau 2010).

Paleoecological analysis of pollen can isolate local and regional vegetation and climate

changes (Chambers and Charman 2004) and together with peat humification, can be

linked to carbon accumulation estimates through bulk density, age-depth model

calculations and C:N ratios to determine the influence of hydroclimatic conditions on

carbon dynamics in peatlands.

Table 1: The three proxies utilized in this study, how they are isolated from the peat matrix, the signal they express, the external influences that force change in the signal and the interpreted reconstruction of each.

Derived Interpreted Proxy Proxy Signal Controlled By From Reconstruction

Organic- (1) Local-regional walled vegetation at/near (1) Hydroclimatology Vegetation microfossils Pollen study site (2) Isostatic Uplift Reconstruction/ isolated (2) Regional pollen (3) Air Masses Succession from the rain peat matrix

(1) Moisture Humic acids (1) Peat forming content Degree of peat chemically vegetation (2) Aerobic decomposition Spectrophotometric extracted (2) Differential decomposition in and therefore, Humification from the resistance to the acrotelm depth to water dried peat decomposition (3) Water table table matrix (3) Compaction of peat depth

(1) Peat forming Carbon (1) Hydrological inputs C:N bulk vegetation accumulation (ombrotrophy vs content of (2) Decomposition estimates; C:N Ratios minerotrophy) the dried in catotelm Isolation of (2) Residence time of peat matrix (3) Peat successional peat in acrotelm accumulation periods

12 1.2 Literature Review

1.2.1 Holocene Climatic Transitions

One of the key allogenic factors influencing peatland development is climate. Large scale climatic and hydrological changes during the period of peat deposition can strongly influence the rate that peat is deposited and sequestered (Zoltai and Vitt 1990). The

Holocene Epoch is formally defined to have begun approximately 11 700 cal year before

2000 AD, based on an abrupt shift in deuterium excess values, changes in δ18O and dust concentration changes, found within the NGRIP ice core from Greenland (Walker et al.

2008). The Holocene can be considered in three phases. The first phase coincides with the Boreal and Pre-Boreal zones, lasting from approximately 11 700 to 9000 years BP, with insolation at a maximum at 10 000 years BP due to the additive effect of the precession and obliquity orbital cycles but with some cooling effects from the remnant

Laurentide Ice Sheet (Wanner et al. 2008). The second phase is known as the Holocene

Thermal Maximum, Hypsithermal, Holocene Climatic Optimum or Atlantic zone and spans the period between approximately 9000 and 6-5000 years BP (Wanner et al. 2008).

The Holocene Thermal Maximum was a period of continuing high summer insolation

(lower than the 10 000 year BP peak) in the Northern Hemisphere and a negligible climatic influence of the Laurentide Ice Sheet on a hemispheric scale (Wanner et al.

2008).

The Holocene Thermal Maximum began approximately 10 000 years BP in the westernmost regions of Arctic and Subarctic North America. However, its onset was delayed in the Hudson Bay Lowlands (perhaps spanning 6000-3000 years BP,

McAndrews et al. 1982; McAndrews and Campbell 1993) as the remnant Laurentide Ice

Sheet had kept the proximal region cool through its impact on the surface energy balance

(Ritchie et al. 1983; Kaufman et al. 2004).

The third phase is the Subboreal and Subatlantic zones lasting from the terminal

Hypsithermal to the Pre-industrial (100 years BP), and is commonly referred to as the

Neoglacial period during which summer insolation declined in the Northern Hemisphere

(Wanner et al. 2008). The conditions during these major subdivisions of the Holocene influenced the establishment of northern peatlands as well as their expansion and succession throughout the Holocene (MacDonald et al. 2006). The terrestrial system is sensitive to these changes in global insolation but the climate response to this forcing is dependent on the amount of radiation, its seasonal distribution across the planet and feedback mechanisms (including ice cover, albedo, ocean and atmospheric circulation)

(Beer et al. 2000).

While the HBL is underreported in relation to the coverage of peat basal dates

(Gorham et al. 2007; MacDonald et al. 2006; Yu et al. 2010) and carbon accumulation estimates, studies have been undertaken on the evolution of the landscape during the

Holocene by focusing on stratigraphic studies of peat profiles. These pioneer works will be discussed to indicate what is known about the HBL, and to illustrate knowledge gaps in the paleoecology of the region, providing a further rationale for this thesis.

1.2.2 Past Paleoecological Studies

Previous paleoecological studies from the HBL (see Fig. 1 for locations) provide some insight into the relative importance of allogenic and autogenic processes in determining peatland vegetation changes. Terasmae and Hughes (1960) developed a pollen diagram

14 for a section along the Attawapiskat River, approximately 90 km west-northwest from the fen study site (Fig. 1), and it is the closest paleoecological reconstruction available for comparison. The section is approximately 150 cm in length, with 100 cm of Sphagnum peat overlying 30 cm of strongly decomposed woody-fen peat which in turn grades into peat with clay, brown clay and finally marine clay (Terasmae and Hughes 1960). The clayey peat contains foraminifera, indicative of brackish water at the site, and the high proportion of Cyperaceae pollen signifies a salt marsh (McAndrews et al. 1982; Sjors

1963; Terasmae and Hughes 1960). The basal peat was dated to 5430 ± 160 cal year BP at a depth of approximately 131 cm (Sjors 1963; Teramae and Hughes 1960) (all dates henceforth are expressed as calibrated years before present; if dates were not calibrated by original authors, they were calibrated using the program CALIB (ver 6.0.1) and the

INTCAL09 calibration curve) (Reimer et al. 2009; Stuiver and Reimer 1993). The peat section begins as a fen, followed by a period of bog development (Terasmae and Hughes

1960). The authors do not make any climatic inferences from the diagram and instead focus on a successional change from fen to bog. However, given the single basal date to develop a rough chronology, the peaks in Sphagnum spores between approximately 4480 to 1950 years BP may correspond to the high proportions of Sphagnum found by other authors, beginning between 3400 and 2500 years BP, and interpreted as evidence of

Neoglacial cooling (Kettles et al. 2000; Klinger and Short 1996; McAndrews et al. 1982).

Sjors (1963) proposed (based on the diagram by Terasmae and Hughes 1960) that the landscape evolved from a brief intertidal salt marsh to a swamp forest and then to a woody fen and finally developed into a bog, following a direction of increasing wetness and a decrease in minerotrophic inputs.

Figure 1: A map of the locations discussed in this section together with the location of the Victor fen site of this study and the surrounding communities of the Hudson Bay Lowlands.

This pioneer record needed to be improved upon because it was relatively short, coarsely dated and did not capture the important influence that climate has on peatland evolution. McAndrews et al. (1982) developed a pollen and macrofossil diagram from R

Lake (approximately 180 km north of the fen site, Fig. 1) in order to develop a longer record. An estimate of lake emergence from the former Tyrrell Sea based on a chronology developed from two radiocarbon dates and the modern sediment surface

(extrapolated to the basal depth) was 8 200 years BP. The pollen record indicated that there was a succession from sparse coastal tundra, dominated by Dryas, willow, sedges and grasses, to shrub tundra, dominated by shrub birch (Betula pumila), to the modern woodland between 8 200 and 6 500 years ago, in response to the decreasing influence of

16 the retreating Tyrrell Sea (McAndrews et al. 1982). The presence of forbs including

Najas flexilis (seed macrofossils) between 6500 and 3000 years BP was interpreted as evidence for the Holocene Climatic Optimum, recorded contemporaneously on the eastern shores of Hudson Bay (Gajewski et al. 1993; Kaufman et al. 2004). However since 2500 years BP, the macrofossil record indicates a decrease in tree abundances, and both Sphagnum bogs and bog forests have become more dominant suggesting some evidence for Neoglacial cooling and heightened rates of paludification (McAndrews et al.

1982). Paludification is defined as the process of bog expansion caused by a gradual rise in water table as the accumulation of peat impedes drainage (National Wetlands Working

Group 1988).

While the record of McAndrews et al. (1982) shows a strong role for climate in driving vegetation change, Klinger and Short (1996) found that hydrological changes driven by isostatic rebound and autogenic processes were important at the Kinosheo Lake bog site in the southern HBL (Fig. 1). Regional pathways for vegetation change over time were proposed based on land cover types and abundances from Landsat imagery, aerial photographs and ground vegetation surveys. These land cover types were used to reconstruct successional pathways through time as changes may be inferred from the sequences identified of different age communities in a spatial array (Klinger and Short

1996). Thus, these successional pathways are based on substituting distance from the present coast for time before present. The moist site pathway represented mesosere

(sequence of ecological communities with a balanced moisture supply) primary succession in the more low-lying, extensive areas between beach ridges, leading to a black spruce bog forest in approximately 2000 years of landscape evolution.

The paleoecological reconstruction of a peak block profile also from the region of

Kinosheo Bog (basal date of 4110 ± 80 years BP) identifies three distinct periods: an early succession zone high in herbs, Pinus, and Picea lasting between 500 and 1000 years, a period of maximum development of Picea woodland and a increase of Sphagnum between 3400 to 2500 years BP, and finally a period of Sphagnum dominated peatland with abundant ericaceous shrubs and an increase in ferns from 2500 years BP to the present. The pollen influx patterns at the site were found to be very similar to expectations from patterns derived from the regional moist-site chronosequence (Klinger and Short 1996). The authors proposed that the mechanisms driving landscape development in the Hudson Bay Lowlands involve a coupling of succession, hydrology, topography and climate (Klinger and Short 1996). As succession takes place over a significant period of time, the factors of topography and physically controlled groundwater hydrology seem to become less important than biotic (autogenic) and climatic influences.

Subsequent work at the Kinosheo Lake bog determined that large scale Holocene climate variations had a greater role than isostatic rebound in the evolution of that peatland. Kettles et al. (2000) analyzed microfossil, macrofossil and geochemical stratigraphy in a peat core from Kinosheo Lake bog (Fig. 1). It was proposed that this bog formed by paludification processes as no evidence of an aquatic fen stage was found in the early peatland record (basal section dated to 4000 ± 80 years BP). A similar succession was put forth by Klinger and Short (1996) for the same peatland. In the period of 4000 to 2500 years BP, pollen assemblage diversity declines in response to the establishment of a Sphagnum-dominated peatland due to cooler conditions (Kettles et al.

2000). The subsequent decline in Picea pollen during the last 2500 to 2000 years was indicative of more open forest cover, and combined with the increased amount of

Sphagnum spores supports a regression of forest cover consistent with the cooling trend that was observed further east (Gajewski et al. 1993, Kettles et al. 2000). This finding is consistent with the increased proportion of Sphagnum found by McAndrews et al. (1982).

Kettles et al. (2000) contended that the major changes in the record are a function of

Holocene climate changes even if (as indicated by peat geochemical data) ecological succession(s) over time also shape peatland dynamics. This supports the common theme of allogenic and autogenic factors both influencing long term dynamics in vegetation change in peatlands.

Using multiple peatlands along a regional chronosequence of isostatic rebound

(akin to the chronosequence studied by Klinger and Short (1996)), Glaser et al. (2004a) sought to corroborate the importance of isostatic rebound on peatland evolution. Glaser et al. (2004a) investigated the stratigraphy of three raised bogs in the Albany River basin

(Fig. 1) along the regional chronosequence, which is reflected in the age of the sites from the youngest (Belec Lake Bog) nearest the coast to the oldest (Oldman Bog) furthest from the coast (Glaser et al. 2004a). The depth of the peat profile also increases inland from the shallowest at Belec Lake to the deepest at Oldman (Glaser et al. 2004a).

Analyses of pollen, plant macrofossils and carbon:nitrogen ratios of the peat deposit were all utilized to investigate the dynamics of bog development. The three bogs exhibit similar pollen stratigraphies (and the same stratigraphic units), with four distinct zones representing the succession of vegetation at each site. The basal zone is interpreted as a tidal marsh at all three sites. This zone is overlain by a fen forest followed by a bog

19 forest, the change supported by an increase in the carbon: nitrogen ratios between zones two and three. This succession is explained by the nitrogen-deficient nature of bog ecosystems, a condition prevalent until other nutrients become limiting (Kuhry and Vitt

1996). The final zone is interpreted as a non-forested bog. This succession from marsh to fen to bog at all three sites mimics that found by Terasmae and Hughes (1960).

This shared stratigraphy between the different bogs suggests that the peatland succession followed the same pathway at each site, driven by geological processes, primarily the isostatic rebound of the region. The authors concluded that the differential pattern of uplift, which reduces the regional gradient and raises water table levels, is the primary driving factor of peatland genesis in the Hudson Bay Lowlands. The bog development conformed to a simple predicted pathway indicating a conservative response of the local biota to the regional environment (Glaser et al. 2004a; Glaser and

Janssens 1986) but the influence of long-term variations in hydroclimatology (especially the climatic conditions during the major subdivisions of the Holocene) was ignored.

More recently, Loisel and Garneau (2010) investigated two peat bogs (Lac Le

Caron and Mosaik, Fig. 1) in the James Bay Lowlands of Northern Quebec using a multi- proxy approach (involving the analyses of plant macrofossils, testate amoebae, peat humification, bulk density and C:N ratios) in order to assess whether hydroclimatic changes resulted from autogenic or allogenic factors. The plant macrofossil based reconstructions provided a more robust understanding of peatland dynamics (than just the inferences made from the testate amoebae) through identifying the patterns of vegetation succession at the sites. However, the testate amoebae captured short-term (multi-decadal)

20 hydrological changes and were more sensitive indicators of moisture conditions than the macrofossils.

Two synchronous changes in hydroclimatology were isolated between the two peatlands with humid conditions around 1000 years BP and wetter conditions from 250 years BP to the present, interpreted by the authors as indicative of the Medieval Climate

Anomaly and the Little Ice Age respectively (Loisel and Garneau 2010). These two large climatic anomalies were not identified or described by the authors of the other studies; the resolution of this study is higher than that of Glaser et al. (2004a), Kettles et al.

(2000) and McAndrews et al. (1982), which may explain why it was able to capture shorter-term climatic changes. The synchronicity between the two sites indicates regional allogenic forcing on peatland development. Site-specific autogenic forcing was also identified through the differences between the cores taken from the ribbed sections of the peatlands and from those taken from the pool sections reflecting the local geomorphic and hydrological states (Loisel and Garneau 2010). The isolation of the relative contribution of both allogenic and autogenic influences on peatland dynamics reaffirms the importance of combining multiple proxies to separate potential drivers whenever possible (Blundell and Barber 2005).

1.2.3 Carbon Accumulation in Peatlands

Each successional study that has been conducted indicates that the HBL often evolves towards a Sphagnum dominated peatland. In this and other types of peatlands, each year’s cohort of litter undergoes some aerobic decay and is buried under the weight of younger material, until the main plant structure collapses. Eventually, the organic material becomes waterlogged and anaerobic, where decay happens a thousand times

21 slower than near the surface, thereby sequestering carbon on long time scales (Belyea and

Clymo 2001). Most peat-forming ecosystems consist of two layers (and are referred to as diplotelmic): the upper acrotelm, an aerobic layer of high hydraulic conductivity where decay is relatively high, and the lower, thicker catotelm, an anaerobic layer with lower hydraulic conductivity and much lower rates of decay (Clymo 1984; Ingram 1978). The boundary between the two layers corresponds to the mean depth of the minimum water table in the peat profile during the summer (Clymo 1984).

The above- and below-ground components of plants (litter) growing on the surface of the peatland decomposes rapidly in the acrotelm, due to such processes as the leaching of soluble organics (Belyea and Malmer 2004; Yu et al. 2001). During passage through the acrotelm, the peat becomes progressively more enriched in the more slowly decaying components, or recalcitrant components, and selective decay may continue in the catotelm, under anaerobic conditions. Thus, the specific composition of peat at depth becomes an increasingly inaccurate representation of the surface vegetation that formed the deposit (Clymo 1984). Litter decay is most rapid in the zone of water table fluctuation, least in waterlogged peat, and intermediate in the oxic acrotelm above the water table (Belyea and Clymo 2001; Ingram 1978). The decaying plant material transitions to peat and is submerged at the base of the acrotelm by the rising catotelm/water table and becomes anoxic as the consumption of molecular O2 by microbial life forms exceeds the rate at which O2 can diffuse down through the water from the air (Clymo et al. 1998). As peatlands can be very long lasting ecosystems, a long-term rate in organic carbon accumulation (LORCA) becomes a meaningful measure

22 to quantify how this sequestration mechanism is influenced by internal and external forcings.

Throughout the Holocene, estimates of the average LORCA for northern peatlands range between 16.2 g C m-2 year-1 - 18.6 g C m-2 year-1, and 44.1 g C m-2 year-1

(Beilman et al. 2009; Gorham 1991; Yu et al. 2010). However, some studies have found higher LORCA estimates for certain peatland types (fens and marshes) of upwards of 72-

80 g C m-2 year-1 (Botch et al. 1995). Similar factors that produce uncertainty in the total carbon pool estimates, including average peat depth, average bulk density of peat and the proportion of carbon in peat combine to result in uncertainty in LORCA measurements

(Botch et al. 1995; Gorham 1991; Turunen et al. 2002). Bogs typically have a higher

LORCA and accumulation is more uniform and predictable than fens, and accumulation tends to decrease from the more southerly peatlands (boreal) to the more northerly

(Subarctic) (Beilman et al. 2009; Tolonen and Turunen 1996; Turunen et al. 2002; Zoltai

1991). Past data sets have contributed to the range of uncertainty surrounding LORCA values because they are biased in the inclusion of profiles almost exclusively from the centre of mires, where peat was the deepest (thus under-representing shallow mires), and from terrestrialized basins at the expense of paludified mires (Turunen et al. 2002).

Carbon accumulation also tends to be more rapid at younger mires as opposed to older mires, with a clear increase in LORCA for peat columns younger than 5000 years

(Tolonen and Turunen 1996).

The generally cool, moist climate during the Holocene has tended to favour C accumulation and maintained the boreal and Subarctic sink of carbon in peatlands

(Turunen et al. 2002). In Canada, the major period of peat (and therefore carbon)

23 accumulation at the northern border of the boreal forest was the early to middle

Holocene, when summers were warmer than present (Ovenden 1990). Mid post-glacial climates were unfavourable for peat growth except in northern peatlands, while the accumulation rates have become lower towards the present (Sjörs 1980). LORCA is influenced by decay (the actual rate of carbon accumulation is lower due to some amount of plant decay in the anoxic zone of the peat), but LORCA still provides useful insight into the dynamics of carbon input and decay (Clymo et al. 1998; Korhola et al. 1995;

Turunen et al. 2002). The humification analysis of the Victor fen core was included to try to account for the influence of decay.

A subsequent study of the same two peatlands studied by Loisel and Garneau

(2010) determined that their Holocene C accumulation rate was 18.9 and 14.4 g C m-2 year-1, for Lac Le Caron and Mosaik respectively (Van Bellen et al. 2011). The late

Holocene reduction in long term C accumulation at these sites (which was a continuation of a gradual slow down) was attributed to both autogenic (local water table mound conditions) and allogenic (climate change) factors (Van Bellen et al. 2011). A new estimate of LORCA for the fen peatland studied is another objective of this research.

1.3 Study Site

1.3.1 Study Region

Peatlands cover approximately 12% of the present land area of Canada, with 97% of these peatlands occurring in the boreal and subarctic wetland regions (Tarnocai 2006), two ecoclimatic regions dominated in Ontario by the nearly unbroken extensive peatland basin of the Hudson Bay Lowland (Sjörs 1963). Extensive peat basins are unique regions

24 where the factors of climate, landscape and local biota produce high water tables that facilitate the expansion of peatlands into adjacent areas (Glaser et al. 2004b). More than

90% of the Lowland itself is a saturated peatland ecosystem, and these organic deposits range from 0.5 m to upwards of 4 to 6 m deep (Martini 2006; Pala and Weischet 1982;

Riley 2003). The depth of peat accumulation is a function of the length of time that the site has been exposed above water, the topography of the underlying material (glacial till or marine sediment) and the distance of the site from the present coastline of Hudson-

James Bay (Glaser et al. 2004a; Pala and Weischet 1982; Martini 2006). There appears to be a strong correlation of peat depth, elevation and distance from the coastline below 65 m a.s.l. for open and treed fens in the High boreal wetland region (Riley 1982).

1.3.2 Site Description

The immediate area of the study site is dominated by the near complete coverage of peatlands (~90%) (Tarnocai et al. 2000). This peatland cover comprises 55% bogs, and

35% fens (Tarnocai et al. 2000). Bogs are distinguished by having a water table at or near the surface, with the surface virtually unaffected by nutrient rich groundwater (and are therefore low-nutrient ecosystems) whereas the fens have a water table at or just above the surface with waters rich in nutrients originating from mineral soils, and a very slow internal drainage by seepage down low gradient slopes (Zoltai 1988). The dominant vegetation of the site was categorized according to the Canadian Forest Ecosystem

Classification. There was no coverage of trees taller than 10 m. The trees or shrubs between 2 and 10 m high were represented by Larix laricina with coverage of 40%. The trees or shrubs 0.5-2 m and <0.5 m height categories contained Betula pumila with coverage of 50% for both. There was herbaceous cover between 75 and 100% comprised

25 of Cyperaceae and Menyanthes trifoliate, and bryophyte cover of 40%. The vegetation community resembles that of the open fens studied by Sims et al. (1982) in the coastal

Albany and Moose River basins. The peatland falls within the region that is severely sensitive to future climate change (Kettles and Tarnocai 1999).

1.3.3 Climate of the Study Area

The peatland studied is located near the northern boundary of the humid high boreal wetland region, very close to the southern margin of the low subarctic wetland region

(Zoltai et al. 1988a). The humid high boreal wetland region experiences cold winters and short, warm summers with the northern and southern boundaries defined by the average summer position of the arctic frontal zone and the winter position of the arctic frontal zone respectively (Zoltai et al. 1988b). The low subarctic wetland region is characterized by very cold winters and short, warm summers and is the location for the most frequent encounters between arctic and temperate air masses (Zoltai et al. 1988a).

The closest Canadian climate station (climate normals period of 1971-2000) to the study site is at Lansdowne House (52°14' N 87°53' W, Fig. 1), at 254 m elevation and approximately 260 km west southwest from the study area (Fig. 1). The mean annual temperature is -1.3°C; the mean January temperature is -22.3°C; the mean July temperature is 17.2°C and the mean total annual precipitation is 700 mm (Environment

Canada 2011). The average annual number of growing degree days (with temperatures

>5°C) is 41 and the period with no snow depth at month’s end is May to September

(Environment Canada 2011).

The study site falls within the zone of discontinuous permafrost, in the scattered or sporadic sub-zone, where permafrost occurs as islands in mostly unfrozen terrain and varies in thickness between less than a few metres at the zone’s southern margin, and 100 metres at the boundary with the continuous permafrost zone (Hydrological Atlas of

Canada 1978).

Hudson Bay generates onshore winds across a temperature and pressure gradient.

Hudson Bay exerts a cooling effect on the surrounding Lowlands in the summer months

(up to 500 km from the coast of the Bay) (Rouse 1991). Temperature forcing has recently begun to change the sea ice regime in Hudson Bay. Trends in surface air temperature

(SAT) anomalies (relative to the 1980-2005 mean) have been found to be positive, expressing a warming of between 0.2-1.8 °C per decade, and were highly correlated with both Sea Ice Concentration (SIC) anomalies and Sea Ice Extents (Hochheim and Barber

2010).The anomalies indicate that temperatures have warmed significantly since the mid

1990s due to the change to the negative phase of both the East Pacific/North Pacific index and the North Atlantic Oscillation (Hoccheim and Barber 2010). Trends in SIC indicate reductions in Hudson Bay of between -36 and -50% during 1980-2005

(Hoccheim and Barber 2010).

Other work has shown that Northern Hemisphere cryospheric cooling has declined by 0.45 Wm-2 between 1979 and 2008, with near equal contributions from sea ice and from land surface snow cover, concurrent with hemispheric warming and representative of a positive feedback of surface reflectivity of climate (Flanner et al.

2011). This has important implications for the climate of the Lowland peatlands that

27 surround Hudson and James bays. These recent developments will impact large-scale climatic and hydrological processes.

1.3.4 Local and Regional Geologic Setting

The Hudson Bay platform consists of the roughly circular Hudson Bay and Moose River basins, separated by the Cape Henrietta Maria Arch (Suchy and Stearn 1993). The fen site is located at the northern edge of the Moose River basin of the Hudson Bay

Lowlands, a low lying, and flat bedrock plain that slopes gently toward Hudson Bay

(Dredge and Cowan 1989; Zoltai et al. 1988a).

The underlying bedrock at the Victor fen is characterized by the Attawapiskat

Formation of the Middle to Lower Silurian System that is predominantly sedimentary carbonate rock (dolostone and limestone) with some sandstone, shale and siltstone

(Ontario Geological Survey 1991). The limestones are predominantly composed of calcite (CaCO3) with minor amounts of calcium-magnesium carbonate (Hattori and

Hamilton 2008). These sequences of Paleozoic carbonate rocks occupy a permanent depression in the Precambrian terrain beneath and adjacent to Hudson Bay (Shilts 1982).

The sedimentary rocks unconformably overlie Precambrian basement rocks (~3 billion year old Archean granite-greenstone belts) of the Canadian Shield (Hattori and Hamilton

2008; Suchy and Stearn 1993).

1.3.5 Quaternary Glacial History of the Hudson Bay Lowlands

During the Last Glacial Maximum, the Laurentide Ice Sheet emanated from multiple plateau-centres of ice accumulation (Dredge and Cowan 1989; Dyke et al. 1989).

Multiple plateaus, including the high plateaus of Labrador-Ungava and Baffin Island, and

28 the lower plateaus of Keewatin and northern Ontario, facilitated the initiation of the

Laurentide Ice Sheet due to the lowering of the regional snowline to the plateau levels

(Dyke et al. 1989). Deglaciation began in the zones of convergence of Hudson and

Keewatin and Hudson and Labrador ice along the ice’s southern margin, which was bounded by glacial lakes prior to deglaciation (Dredge and Cowan 1989; Shilts 1982).

Paleogeographic deglaciation maps (developed from a radiocarbon chronological database) indicate the Victor site was deglaciated between 8600 and 8450 cal. year BP

(Dyke 2004). The site then likely became covered by the expansive Glacial Lake

Agassiz-Ojibway until its final abrupt drainage placed at 8205 cal year BP (Roy et al.

2011).

The Tyrrell Sea marine incursion resulted from marine waters entering the

Hudson/James Bay region along a break between Hudson and Labrador Ice and between

Hudson and Keewatin Ice (Dredge and Cowan 1989; Peltier and Andrews 1983). The

Tyrrell Sea reached its maximum extent between 7000 and 8000 years ago and regressed as upwarping of the land began due to the removal of the load of the ice sheet (Dredge and Cowan 1989; Lee 1960). The sea amassed fine grained sediments of predominantly silt and clay, which were deposited as quiet water sediments at the study area (Fulton

1995). These sediments underlie the peat profile studied at the Victor fen site.

Surficial Quaternary deposits of the fen site are characterized as organic deposits of undifferentiated peat, muck and marl (Pala et al. 1991). The site lies adjacent to raised beaches or bars of glaciolacustrine, glaciomarine, or marine origin. These features are

29 present northeast of the site (<10 km distant) and immediately southwest (<5 km distant, many oriented north-northwest – south-southeast) (Pala et al. 1991).

1.3.6 Post-glacial Isostatic Adjustment

The Lowlands was the geological province located the closest to the centre of the

Laurentide Ice Sheet outflow (Riley 2003). The effects on the elastic crust near the ice sheet margins would have been twofold: 1) a forebulge would have been produced due to the elastic bending of the lithosphere above the pre-glacial equilibrium level, and 2) at the ice margin the ground surface would have been forced below the equilibrium surface

(Walcott 1970). The loaded crust area would sink below equilibrium as the mantle flows outward (Henton et al. 2006). Upon deglaciation, the mantle, which behaves like an extremely viscous fluid, would flow back into the regions where the ice load forced its dispersal. This would cause the lithosphere to rebound in those regions and would also result in the collapse of the peripheral forebulge that had formed along the margins of the ice sheet resulting in subsidence of the lithosphere (Henton et al. 2006).

Relative sea level curves describe the relaxation conditions during deglaciation and are influenced by a combination of glacial-isostatic rebound and the rise in global sea level due to the melting of the Continental Ice Sheets (Andrews and Peltier 1989; Walcott

1972). Holistic isobase maps based on published relative sea level curves for North

America (Andrews and Peltier 1989) provide regional estimates of post-glacial isostatic adjustment. Figure 2 shows two differing modelled emergence curves (an exponential and a quadratic) for the Victor fen site based on the following isobase data. Both models are shown because some post-glacial recovery curves from Southern Hudson Bay do not exhibit an initial rapid emergence which is characteristic of curves drawn for regions

30 within the ice dispersal centres (Dredge and Cowan 1989), so a quadratic model was included because it does not exhibit this initial rapid emergence.

The isobase map (Andrews and Peltier 1989) of relative sea level change from

7000 BP to present indicates an emergence of approximately 210 m for the Victor site which equates to an adjustment rate of 3 m century-1. The isobase map of relative sea level change from 4000 to 3000 BP indicates an emergence of approximately 26.5 m or an adjustment rate of 2.65 m century-1. The isobase map of relative sea level change from

2000 BP to present indicates an emergence of 24 m for the study site, which is an adjustment rate of 1.2 m century-1 and is useful for determining the present rate of land emergence (Andrews and Peltier 1989). Lastly, the isobase map from 1000 BP to present exhibits an emergence of 10-11 metres, or an adjustment rate of 1.0 - 1.1 m century-1.

These isobase maps outline a central area of emergence over James Bay, extending out into the HBL. The maps indicate a slowing of emergence (from 3 to 1.1 m century-1, see Figure 2), just as many curves exhibit a rapid initial emergence, and then usually decline as a simple exponential curve (Dredge and Cowan 1989).

The isobase maps compare very closely with estimates of emergence for Fort

Albany (0.9 to 1.2 m century -1, calculated from aerial photographs and historical archives of the Hudson Bay Company), Cape Henrietta Maria (1.2 m century-1 derived from fitting a post-glacial emergence curve to radiocarbon dated marine strandlines) and the York Factory Peninsula (1.0 – 1.3 m century-1 , calculated from comparisons made between modern and historical maps) (Hunter 1970; Tarnocai 1982; Webber et al. 1970).

However, the Cape Henrietta Maria curve was used in the contouring of the isobase lines,

31 explaining the excellent agreement. The estimates also compare closely with observations from the Canadian Base Network, a network of stable pillar monuments fitted with GPS receiver antennae (Henton et al. 2006). The observed vertical rates from this network are between 10 and 12 mm year-1, or 1 to 1.2 m century-1 (Henton et al.

2006). However, the density of the CBN network is coarse with only two sites in the

HBL.

400

300

200

100 Relative sea changesRelative level (m)

0 2000 4000 6000 8000 Year (ka BP)

Figure 2: Postglacial emergence curves for the Victor Fen site based on the isobase maps in Andrews and Peltier (1989) and the modern surface. The solid line is an exponential model and the dotted line is a quadratic model.

1.3.7 Local and Regional Vegetation

The fen site is located in the peatland and woodland floristic zone (Riley 2003). The five most widespread arboreal species in the HBL are Populus balsamifera, Populus tremuloides, Larix laricina, Picea glauca and Picea mariana, and there are 40+ shrubs

(12 species of Salix, 6 species of Ribes, Betula pumila, Myrica gale, Alnus crispa, Alnus rugosa, Ledum groenlandicum, Kalmia polifolia amongst others) (Riley 2003).

Horizontal fen conditions are indicated by an open canopied forest with Larix laricina the most common tree species, and shrubs including Betula pumila dominating portions of the fen (Zoltai et al. 1988b). Mosses including Sphagnum teres, Sphagnum warnstorfii, and Sphagnum fallax are found in low hummocks or wet carpets (Zoltai et al.

1988b). Herbs including Scirpus caespitosus, Scirpus hudsonianus, and Equisetum fluviatile are characteristic of treed or shrub horizontal fens. In the wettest section of the fens, species including Carex exilis, Carex lasiocarpa, Scirpus caespitosus, Eriophorum viridicarinatum, Habenaria dilatata and Menyanthes trifoliata are common (Zoltai et al.

1988b). In the southern James Bay area, horizontal fens are dominated by Larix laricina and Sphagnum warnstorfii and contain small, streamlined islands of Picea mariana

(Zoltai et al. 1988b).

On drier uplands, both black and white spruce (Picea mariana and Picea glauca) occur in relatively pure stands or mixtures with balsam fir (Abies balsamea), while on sandy soils or following forest fires, jack pine (Pinus banksiana) grow in even-aged stands, occasionally mixed with white birch (Betula papyrifera) (Zoltai et al. 1988b).

The vegetation characteristic of salt marshes in the Attawapiskat area near the present coast is also important to discuss given the incidence of this ecosystem in the early part of the paleoecological reconstructions. At low tide, the dominant colonizing species observed was Hippuris tetraphylla along with a smaller proportion of Scirpus validus, and Carex paleacea. These three species together are representative of brackish

33 environments in southern James Bay (Glooshenko and Martini 1983). Away from the river bank, the assemblage changes to one typical of the James Bay salt marshes including first the “low salt marsh species” Puccinellia phryganodes and Scirpus maritimus followed by the “high salt marsh” species of Carex subspathacea with lower proportions of Festuca rubra and Triglochin maritima (Glooshenko and Martini 1983).

These salt marshes transition to a willow thicket including Salix candida, Salix cordifolia, Salix brachycarpa with various forbs and grasses (Glooshenko and Martini

1983).

2 Chapter 2 METHODS 2.1 Field Methods

Complete cores through the peat sequence and into the Tyrrell Sea marine sediment contact were collected from a fen in the Attawapiskat River watershed (52.7123°N,

84.1714°W; 100 m asl) of far northern Ontario in July 2009. The location was recorded with a handheld GPS. Permafrost was not encountered during the coring of the peatland.

The structure of the surface of the fen was uniform. The peat sequence extends to a depth of 245 cm, at which point the core grades into fine marine sediment.

Peat cores were retrieved using both a Jeglum corer (Jeglum et al. 1992) and a

Russian chamber corer (Jowsey 1966). The Jeglum corer is a surficial box corer that is driven down through the first ≤50 cm of the peat profile. Collecting peat samples at depths greater than 50 cm necessitated the use of extension rods attached to the Russian corer. The Russian corer is lowered to the desired depth of the peat profile and turned

180° against the resistance of the “fin”, thereby enclosing the 50-cm sample. The corer isolates the sample from the surrounding matrix once the “fin” is closed allowing the user to retrieve a sample free from distortion or contamination from other levels. The Russian corer collected the remaining profile (to a depth of 260 cm), divided into 50-cm drives.

A paired beta Russian core was taken adjacent to the primary core for replication.

The peat samples were wrapped in aluminum foil and drain pipe and stored in a cooler under refrigeration until they were shipped back to the Ontario Forest Research Institute in Sault Ste. Marie and then onto the University of Toronto. These samples were then deposited in a cold room for archiving until sub-sampling.

34 35

2.2 Laboratory Methods

A chronology of the fen core was constructed from radiocarbon dates on terrestrially derived material picked from the cores themselves. To start, the goal of four Accelerator

Mass Spectrometry (AMS) radiometric dates, or one for every 60 cm, was attempted. A 1 cm slice of the core was taken (avoiding the outermost material to avoid contamination) and put into a 250-ml beaker. To this beaker, 100 ml of 5% KOH was added to disaggregate the peat and the beaker was brought to a light boil on the hot plate for 10 minutes. Next, the sample was sieved through 90-µm nylon mesh and washed until the filtrate was clear, thereby removing all the KOH. The sample was then transferred to a

50-ml centrifuge tube, shaken well and then a small amount was poured into an unused, rinsed, disposable Petri dish. Using the stereomicroscope, each sample was scanned at

10x magnification and forceps were used to pick out larger pieces of wood, moss, leaves, or other organic remains; fractions of different materials were placed into separate 1.5-ml clear plastic vials. The entire centrifuge tube was picked through, to obtain as much organic material as possible for radiocarbon dating.

If insufficient material was recovered from the first 1 cm section of the core, an adjacent 0.5 cm section was processed and picked completely. Before submitting the vials to be dated, each one was re-picked and washed with distilled water to ensure organic material was free of other material. Material dated included conifer needles (three levels), twigs (two levels) and an unidentified piece of wood (one level). The conifer needles were identified with the aid of an illustrated guide (Lévesque et al. 1988).

Samples were sent to Beta Analytic Inc. for AMS dating. A total of six dates were retrieved for the fen core, including a basal date.

The goal was also to have recent (the past 150 years) dating done using 210Pb activity in the peat core. Using a utility knife, rectangular pieces of peat were cut from each depth increment that was to be dated. Each piece was kept intact while measurements of the dimensions to calculate volume were conducted, and then the sample was added to a metal cup. Wet weights were recorded immediately to prevent error associated with evaporative losses. Samples were placed in the drying oven at a maximum temperature of 60°C, dried to a constant weight, and ground to a fine powder using a mortar and pestle. The mortar and pestle were wiped clean between samples with a kimwipe. Seven samples were then transferred to 15-ml centrifuge tubes and shipped to

Flett Research Ltd. (Winnipeg, Manitoba) for measurement of 210Pb activity. The bulk density of the peat was calculated by dividing the dry weight by the volume of peat sampled to result in an amount in g cm-3, and sent along with the samples.

Wet samples were taken directly from the core for peat humification analysis, which followed a protocol modified from Blackford and Chambers (1993). Since previous work indicated a weight loss of 90%+ during drying, sample size could be adjusted accordingly for a desired dry weight of 0.2 g peat. Samples of 1 cm width were taken at 3 cm intervals (when possible), with the outermost material left intact to avoid contamination. These samples were placed in a metal cup and dried to a constant weight in the oven at 60°C for >24 hours. Dried samples were then ground to a fine powder using a mortar and pestle and 0.2 g of powdered peat was added to a 150-ml beaker. 50 ml of 8% NaOH was added to each beaker, and samples were kept well mixed with glass stir rods. Each solution was then warmed on a hot plate until boiling, at which point the heat was reduced and the samples simmered for 45 minutes (samples were monitored

37 carefully to ensure each remained at constant temperature). When the volume of the solution dropped below 30 ml in any one of the beakers, 20 ml of distilled water was added to each sample.

At the end of 45 minutes of heating, the volume of each solution was topped up to

120 ml using distilled water. These samples were stirred, and 50 ml of each solution was added to large centrifuge tubes. The samples were centrifuged at 2500 RPM for four minutes, and the supernatant was filtered through a funnel and Whatman No. 1 (150-mm) qualitative filter paper. 100 ml of distilled water was added to the filtered solution for dilution. A reference blank of 5 ml of 8% NaOH and 12 ml of distilled water was prepared to ensure the spectrophotometer was not drifting during analysis, which was checked both halfway, and at the end of each batch of samples. For each sample, 3 ml was transferred via pipette into a glass cuvette and the spectrophotometric absorbance was measured at a wavelength of 540 nm. This wavelength is best suited to peat-based climatic studies because maximum variability in absorbance is preferable for sensitivity to hydrological conditions (Blackford and Chambers 1993). The absorbance reading was repeated twice per sample, and the two readings were averaged for a single absorbance value. One replicate per batch of samples (11 samples and 1 replicate for a total of 12) was included in the processing to measure error. Humification values were then detrended using a quadratic model. These detrended absorbance values better represent the conditions of peat decomposition for interpretation by eliminating the depth- dependent trend of anoxic decay in the catotelm (Mauquoy et al. 2002a).

Wet samples were also taken directly from the cores for pollen analysis. A protocol derived from Faegri and Iverson (1989), using all the pertinent treatments for

38 processing peat, was used. Samples were taken at an average interval of 6.25 cm, and a maximum and minimum interval of 9 and 3 cm respectively. 1 ml of peat was sampled from the core using an open topped syringe and transferred to 15-ml centrifuge tube. 1-2 tablets of exotic Lycopodium, with a known number of spores per tablet (Stockmarr

1971), were added to each tube and dissolved in 10 ml of 10% HCl, to quantify pollen concentrations in the samples. These tubes were centrifuged (throughout the process, the tubes were centrifuged for 4 minutes at 2500 RPM), decanted, stirred and then washed with water. Next, 5 mL of 10% KOH was added to each tube and the tubes heated for 3 minutes in a hot water bath during which the samples were stirred intermittently. This treatment removes humic acids or unsaturated organic soil colloids, as well as acting as a deflocculation step (Faegri and Iverson 1989).

The contents of the tubes were then poured through a 150-μm (coarse) nylon mesh sieve. All material that passed through the sieve was kept, thus removing any coarse particles from the sample. With the use of an engraver, the contents were then passed through a 10-μm (fine) nylon mesh sieve, and all material that did not pass through the sieve was kept and transferred to a centrifuge tube with distilled water and washed.

Three samples necessitated treatment with hydrofluoric acid (HF) due to the very fine particulate matter of the marine sediment at the base of the core. HF dissolves silica

(including clay) while not appreciably attacking organic remains (Faegri and Iverson

1989). These three samples had 5 ml of HF added to them and were heated in a water bath for 1 minute and decanted. The samples were then washed twice each as a precaution to remove all the HF. All samples were then treated with 5 ml of glacial acetic

39 acid (GAA), centrifuged, decanted and stirred. Each sample then had 5 ml of acetolysis solution added to it and was heated in the hot water bath for 1 minute. The acetolysis solution was made with 9 parts acetic anhydride ((CH3CO)2O) and slowly adding 1 part sulphuric acid (H2SO4) to this solution. Acetolysis removes polysaccharides, such as cellulose, as the peat cores are organogenic deposits (and therefore contain abundant cellulose) (Faegri and Iverson 1989). After centrifuging, decanting and stirring, the samples were washed with GAA and then subsequently washed with distilled water.

1-2 drops of safranin stain were added to each tube, stirring well, followed by 5 ml of Tert-butyl alcohol (TBA) and stirred again. TBA is used in the dehydration of the remaining material in the sample. Following centrifuging and decanting, the contents of the tubes were transferred to small glass vials. These vials were levelled using TBA, centrifuged, and decanted. Finally, silicone oil (2000 centistokes viscosity) was added to each vial to completely cover the sample and each vial was stirred extremely well.

Silicone oil is more permanent, does not cause as much swelling of pollen grains and has a lower refractive index than glycerol, making it more advantageous to use as a liquid mount (Faegri and Iverson 1989). Each vial was left uncorked in the fumehood overnight to allow the TBA to evaporate, and then each sample was ready to be mounted on a microscope slide.

Slide preparation commenced after the samples were well stirred. Pollen identifications were made with the aid of reference keys (Kapp et al. 2000; McAndrews et al. 1973; Richard 1970), in addition to in-house reference slides from the Royal

Ontario Museum. Each sample was enumerated to a minimum tree pollen sum of 200 under 400 x magnification, using a combination of bright-field and differential

interference contrast (DIC) microscopy. This minimum sum was chosen because it was

found that further counting yielded no additional species that had yet to be documented

(Table 2). This pollen sum is consistent with those used by McAndrews et al. (1982),

Klinger et al. (1996), Kettles et al. (2000) and Glaser et al. (2004a) for other areas of the

Hudson Bay Lowlands. Transects across each microscope slide were chosen randomly

and the entire slide was enumerated at 1-mm intervals unless the minimum sum was

reached prior to slide completion. If the minimum sum was not reached, then additional

slides were made from the processed sample and subsequently enumerated. Rough size

measurements were made to attempt to separate Picea mariana and Picea glauca.

Following Lindbladh et al. (2002) and Klinger and Short (1996), unbroken grains

averaging over 100-110 μm in maximum diameter were counted as Picea glauca, and

anything smaller was treated as Picea mariana.

Table 2: Grain count chart emphasizing the rationale for a pollen sum of 200 arboreal pollen grains. No new undocumented taxa were discovered above 200 arboreal grains at these test levels.

No. of No. of New Maximum proportion New Sample species for species for Species in Proportion of new of new species Species at Depth first 100 second 100 2nd 100 species at level found at any 250 arboreal (cm) tree pollen tree pollen grains analyzed depth grains

97 9 8 0 0 0 0

121.5 15 9 0 0 0 0

Ericaceae = 0.3% ; Ericaceae = 1.5% ; 161.5 10 12 3 Ambrosia = 0.3% ; Ambrosia = 3.9% ; 0 Larix = 1.1% Larix = 1.6%

Acer = 0.18% ; Acer = 0.87% ; 181 10 12 2 0 Ambrosia = 0.37% Ambrosia = 3.9%

The total number of fossil pollen grains in the sample was computed by the formula: Total fossil pollen = (Fossil pollen enumerated x Total number of Lycopodium added) / (Total number of Lycopodium enumerated). The pollen concentration was then computed by dividing the total fossil pollen amount by the volume of the sample. Pollen influx for each taxon was calculated by the formula: Pollen influx = [(Fossil pollen of taxa Z enumerated x Total number of Lycopodium added) / (Total number of Lycopodium enumerated) / Years represented by sample]. The total pollen influx was calculated by dividing the total pollen concentration at a given depth by the number of years that the 1 cm depth represented in the age-depth model, yielding an influx in grains cm-2 year-1.

Statistical reliability demanded that the number of marker Lycopodium could not be less than 20% of the expected fossil pollen total (Faegri and Iverson 1989) and this constraint was met with all but one sample, which corresponded with the peak in Sphagnum spore proportion for the peat sequence.

The proportions of each taxon recorded at each distinct level were input into the software C2 (Ver. 1.6.8) (S. Juggins, University of Newcastle), used in the analysis and graphing of ecological and paleoenvironmental information, to construct the pollen diagrams. Proportional abundances of fossil pollen were then input into Zone (Ver. 1.2)

(S. Juggins, University of Newcastle) and subsequently evaluated using CONISS, a program for implementing agglomerative, hierarchal stratigraphically constrained incremental sum of squares clustering (Grimm 1987). Zones were identified and delimited using the total within-cluster dispersion dendrogram (used to illustrate the hierarchal relationship of the clusters) that was cut at a given height (below the first branching). Total within-cluster dispersion is not subject to reversals and emphasizes the

42 progressive construction of the clusters and generally provides the most information when identifying zones (Grimm 1987).

The fossil pollen data were integrated with the North American surface pollen dataset (Whitmore et al. 2005; http://www.geography.wisc.edu/faculty/williams/lab/Downloads.html), a dataset of more than 4500 sites, in order to perform the paleoclimatic reconstructions. Reconstructions were performed in the software C2 using the Modern Analog Technique (MAT)

(Overpeck et al. 1985). The MAT functions by measuring dissimilarity between a fossil sample and each individual point in the calibration dataset, and the environmental variables of the most similar modern samples specified are averaged and assigned to the fossil sample (Overpeck et al. 1985; Williams and Shuman 2008). The dissimilarity coefficient used was the squared chord distance (Overpeck et al. 1985), the number of modern analogs specified per fossil sample was 3 (Williams and Shuman 2008), and 500 bootstrap attempts were made per sample. Multiple analog matches per sample reduces stochasticity and improves the precision of reconstructions (Williams and Shuman 2008).

The database was processed to include only those sites with ≥150 pollen grains counted

(3604 sites) and only those taxa that were found in the fossil assemblage that co-occur in the modern database (18 taxa). This was done to partly capture the rapid gain in reconstruction precision that occurs between 75 and 300 grains (Lytle and Wahl 2005) while still using a large number of modern sites. All the depositional environments that the modern database contains were accepted. Pollen assemblages from lake sediments provide regional-scale records of past communities (Gajewski et al. 1993), and therefore regional scale estimates of climatic change (Bradley 1999). When combined with the

43 local-regional vegetation signal in peatland records (Chambers and Charman 2004), it was felt that the climate reconstructions would highlight both local and regional hydroclimatic signals. White and black spruce (Picea glauca and P. mariana) were amalgamated together as many sites had not differentiated the spruce grains.

Wet samples were also taken directly from the core for bulk density and elemental carbon: nitrogen (C:N) analyses. These samples were taken at an interval of 3 cm or less.

1 cc of peat was sampled from the core with an open topped syringe and transferred to a metal sample dish. This sample was weighed to determine the wet weight and then placed in a drying oven. Samples were dried for >24 hours at 100°C and then the dry weight was taken upon removal from the oven. The bulk density of the samples was then calculated.

Once dry, the samples were ground to a very fine powder using a mortar and pestle and then transferred to a 1.5-ml microcentrifuge tube. These tubes were stored in a glass desiccator to ensure that the samples would not take on any moisture. The samples were analyzed on an ESC 4010 Elemental Combustion System for CHNS-O (configured only for C and N) (Costech Analytical Technologies, Valencia, CA). 5 mg of the ground peat was used per analysis, which returned the percentage of carbon and nitrogen in the sample. LORCA values were obtained by multiplying the percentage carbon data by the bulk density to retrieve a carbon mass (in g C cm-3) and then dividing the carbon mass by the number of years the sample represented as defined by the age-depth model, resulting in a value in g C cm-2 year-1, which was then converted to g C m-2 year-1.

3 Chapter 3 RESULTS 3.1 210Pb Dating of VICM_T3_SP3

The activity of 210Pb had a vertical profile, indicating that background 210Pb activity has not been achieved in this core (Fig 3), indicating either a mixed peat profile, or an extremely high peat accumulation rate (~16 cm in less than about 10 - 20 years), neither of which was likely in this peat core. The only explanation given by the analysts to the vertical and elevated 210Pb profile would be natural sources of 210Pb (such as radon) continually entering the core deep in the profile. As this core was taken in a fen environment, inflow of groundwater into the lower peat profile from the underlying marine sediment, or adjacent peatland is possible. The lab recommended collecting a sample of groundwater at the site to measure the radon levels and test the validity of this explanation but we were unable to acquire a sample for this purpose. Therefore, the chronology of the fen core will be derived solely from radiocarbon dates.

44 45

Depth (cm) Depth 12

18 10 100 210 Pb Total Activity (DPM/g)

Figure 3: The activity of 210Pb in the uppermost 20 cm of the Victor fen core. Note the x axis is on a logarithmic scale. The vertical profile does not follow an expected trend of decay in activity with depth, indicating that background 210Pb activity was not achieved.

3.2 Age-Depth Model Development

The radiocarbon dates were calibrated in the program CALIB (ver 6.0.1) using the

INTCAL09 Northern Hemisphere atmospheric radiocarbon calibration curve (Reimer et al. 2009; Stuiver and Reimer 1993) (Table 3). Development of an age-depth model was done by curve fitting in the software SigmaPlot (Ver. 11 Systat Software Inc.). Linear interpolation between the radiocarbon dates implies that sedimentation rates suddenly change at the depths of each date (Telford et al. 2004). The dates for the Victor fen core were spaced equally throughout the core, rather than at identifiable stratigraphic transitions, so the use of a linear model was unrealistic because it resulted in rapid

46 changes in accumulation rates at points in the core with no apparent stratigraphic change.

Thus, the age-depth model was derived from fitting a 3-parameter sigmoid function to the six radiocarbon dates and the modern peat surface (Fig 4). A sigmoid function was chosen because a cubic function returned age reversals and a quadratic function did not match the trend exhibited by the radiocarbon dates (a period of high accumulation, followed by a long period of low accumulation, followed finally by another period of high accumulation), and so an s-shaped age-depth model best reflected the trend suggested by the radiocarbon dates.

The r2 of the sigmoid function was higher than the quadratic (0.96 versus 0.9433), and both were higher than a linear regression, suggesting a slightly better fit for the radiocarbon dates. However, r2 is not the most useful guide due to the fact that if the sediment sequence is in stratigraphic order and deposited over an appreciably long period of time (two constraints met with the fen peat sequence), r2 will often be high and will increase with higher-order polynomial functions (Telford et al. 2004). Fitting a 3- parameter sigmoid to the six radiocarbon dates and the modern peat surface yielding the function f = 7731.4442 / (1 + exp(-( X-167.9763) / 40.2202)), where X is the depth in the core for which the function solves an age. The r2 of this function was 0.96. The coefficient of determination (r2) represents the proportion of variability in the data set (in this case the radiocarbon dates) that is accounted for or explained by the sigmoid model, suggesting that the model is a very good fit to the radiocarbon dates.

A correction of -175.9 years was added to each level for which a model inferred date was calculated (every 0.5 cm depth). This correction was deemed necessary because

47 the model did not pass through the only known date of -59 years BP for the modern surface of the peat core.

6000

5000

4000

3000

2000 Age (cal Year BP)Year Age (cal

1000

0 50 100 150 200 250 Depth (cm)

Figure 4: Sigmoidal age-depth model derived for the Victor Fen Core. The original model is the dotted line, and the solid grey line has been adjusted so that the model passes through the modern sediment surface (-59 yrs BP = 0 cm). Radiocarbon dates are indicated by diamond symbols with 2-sigma ranges as error bars.

Table 3: AMS radiocarbon dates recovered for the Victor Fen Core (VICM_T3_SP3), the type of material dated and the intervals from which these samples were taken

Conventional Sample Laboratory Sample Depth Dated 2σ Calibration Median Age 13C/12C Ratio Age 14C (years Number Number (cm) Material (years BP) (years BP) (‰) BP)

VC01 37 Beta-286595 37-38.5 Conifer needles 580 ± 40 577-653 600 -28.7

VC01 60 Beta-281774 60-61 Imbedded Twig 1250 ± 40 1076-1276 1196 -26.2

VC01 120 Beta-281775 120-121 Wood 1660 ± 40 1507-1633 1564 -26.8

VC01 157 Beta-286596 157-158.5 Conifer needles 2730 ± 40 2756-2890 2824 -28.5

VC01 194 Beta-281776 194-195 Wood 5050 ± 40 5710-5908 5815 -26.9

VC01 242 Beta-281777 241.5-243.5 Conifer needles 5640 ± 40 6315-6494 6421 -28.6 48

3.3 Paleoecological Reconstruction

A total of 44 distinct levels were counted for pollen analysis, but only 42 were included in the paleoecological reconstruction (Figs. 5 and 6). The core section from 224-226.5 cm has unusually high Betula pollen percentages and two levels (226 and 226.5 cm depths) were rejected due to overrepresentation of Betula because it is hypothesized that a catkin fell onto the site where the core was taken from and was incorporated in the peat deposit.

The number of pollen grains remaining in catkins that fall to the ground is very high and may even be greater than that which was released into the air at the time of flowering

(Faegri and Iverson 1989, referencing Rempe 1937), potentially explaining these anomalous samples. The levels have 79% and 76% Betula proportions respectively, compared to the highest included proportion (35% at 224 cm depth) and the average proportion of Betula at all 42 levels (6.8%). It is possible that some of the Betula grains that form the peak level at 224 cm are from this hypothetical catkin and were displaced via upwash and subsequent decomposition as the peat sank through the acrotelm, which has been experimentally demonstrated in the unsaturated surface layer of Sphagnum dominated peat (Clymo and Mackay 1987). Total arboreal sums ranged between 200 and

379 grains. Total palynomoph sums for the paleoecological reconstruction ranged between 367 and 2044 grains. 27 distinct taxa were identified in the samples that were counted but a subset of the 15 most abundant taxa was used in the reconstructions (Figs.

5 and 6).

The percentage pollen diagram was split into three principal biostratigraphic zones using cluster analysis: zone 1 (256-240 cm), zone 2 (240-20 cm) and zone 3 (20-0 cm). The basal assemblage zone is split into two subzones: zone 1a corresponds to the

50 samples of the uppermost marine sediment and zone 1b refers to the basal peat sequence.

Zone 1a is characterized by the highest values of the regional arboreal Pinus banksiana

(14-19%), the highest values of Salix (3%), Chenopodiaceae (~3%), and low values of

Sphagnum (16-17%). This transitions to zone 1b, which is dominated by Typha latifolia

(15-81%), and Cyperaceae (5-42%), has a paucity of Sphagnum (0.5-7%) and high pollen concentration (~150 000 grains ml-1) due to contribution from local Typha.

The basal assemblage zone is overlain by a pollen assemblage of zone 2, which is split into 3 subzones. Subzone 2a is characterized by the disappearance of Typha latifolia, a decrease in Cyperaceae (0.8-8%) an increase in Sphagnum to the peak abundance (from 55% to 81%), and the peak abundances of Betula (35%) and Ericaceae

(2%). The arboreal pollen is dominated by Picea mariana (13-30%), with Pinus (1-11%) and Picea glauca (<1%) subordinate. Pollen concentration also peaks in this zone (~340

000 grains ml-1), due to the large contribution of the Sphagnum peak. Subzone 2b contains the peak proportion of Picea mariana (upwards of 40%), Picea glauca (2%),

Larix (1.5%) and three cycles of increase-to-decrease in Sphagnum proportions. Pollen concentrations decline towards the top of the zone. Subzone 2c contains a rise in

Ambrosia pollen (2-4%, versus 0.1-1.4% throughout the rest of the zone), as well as high

Sphagnum (47-58%), and a decrease in Picea mariana (14-19%).

The uppermost zone is characterized by the peak in Alnus (8-9%), an increase again in Cyperaceae (15-27%), a decrease in Sphagnum (to 13%), the absence of Larix, and the lowest pollen concentrations (9 500 – 15 000 grains ml-1). Calculations of pollen influx have adjusted the very low concentrations of zone 2c and zone 3 to reflect the lack of compaction of the peat. Pollen influxes of between 3 300 and 10 700 grains cm-2 year-1

51 are high compared to the rest of the sequence (average 2 900 grains cm-2 year-1), and to the peak influxes associated with the peak pollen concentrations (6 800 and 7 900 grains cm-2 year-1 for the Typha and Sphagnum peaks respectively). Pollen concentrations at equivalent depths represent a mixture of influences including a higher contribution from local vegetation, greater peat decomposition (due to less surface moisture) yielding more pollen in a similar volume or greater peat accumulation (due to more humid surface conditions), yielding less pollen in a similar volume. Pollen grains attributable to long distance transport (500+ to 1000 km) include Acer, Juglans cinerea and J. nigra, and

Quercus (Fowells 1965).

The influx diagram has some notable differences from the percentage pollen diagram. In the marine sediment at the base of the core, there is low influx from arboreal and shrub taxa, despite the high relative proportions of Pinus, Salix, Picea and Alnus.

Subzone 1b is similar to the percent diagram, with the large peaks for Cyperaceae and

Typha evident as large influxes (3360 and 1145 grains cm-2 year-1 for both taxa respectively). Peak arboreal influxes occur in subzones 2b and 2c for Pinus (1390 grains cm-2 year-1) and Picea mariana (4180 grains cm-2 year-1) respectively, with high levels sustained into zone 3. Betula has high influx at the base of subzone 2a (533 grains cm-2 year-1) that corresponds to its peak proportion, but also has high values throughout zone

2, and into zone 3 with peak influx in the surface sample (630 grains cm-2 year-1). Other shrub species also have peak influxes at the boundary between zones 2c and 3 including

Alnus (1090 grains cm-2 year-1), Salix (435 grains cm-2 year-1) and Chenopodiaceae (405 grains cm-2 year-1) in addition to the peak influx of Sphagnum spores (12605 grains cm-2 year-1).

Figure 5: Percentage pollen diagram from Victor fen core. Also included are pollen concentration and pollen influx. The primary vertical axis is depth, and the secondary is in years before present (scaled based on the age-depth model). X-axis scaling varies. 52

Figure 6: Pollen influx diagram for the Victor fen core. X-axis scaling varies. 53

3.4 Bulk Density

Bulk density increases with depth, and nearly doubles (0.101 to 0.196) through the peat sequence (Fig. 7). A strong relationship has also been found by other authors between dry peat mass and the time that the peat has accumulated in bogs and fens (Turunen et al.

2002; Zoltai 1991). The drop in density for the last three samples (236, 240 and 242 cm) is due to the fact that the peat at those depths when the samples were taken had been drying out due to being improperly sealed.

0.20

) 0.18 -3

0.16

0.14

Dry Bulk Density (g cm (g Density Bulk Dry 0.12

0.10 0 50 100 150 200 250 Depth (cm)

Figure 7: Bulk density of the Victor fen core.

3.5 C:N Stratigraphy

The instrument available for C:N analysis requires small sample sizes (~5 mg). Because peat is a heterogeneous mixture of materials, a pilot project was conducted to determine

55 if our lab procedure of homogenization was effective. Variance is expected to be high among replicates if the sample was inadequately homogenized; in this case, variance between samples was also expected to be similar to the within-sample variance determined from the replicates. If the sample was properly homogenized, low variance among replicates and higher variance between samples was expected. Six depths were sampled, each with a replicate sample, and each sample was run 5 times, for a total of 10 runs per depth. The mean values and mean standard deviations of the elements for the pilot study were found to be 46.82% and 0.23 for carbon and 2.76% and 0.048 for nitrogen respectively. The average variance of the carbon and nitrogen proportions for the within sample treatment was 0.058 and 0.0041 respectively. The average variance of the carbon and nitrogen proportions between samples was 2.28 and 0.053 respectively.

Due to the very low standard deviations, and the within-sample variance being substantially lower than the between-sample variance, the homogenization of the peat matrix was deemed successful.

A one way ANOVA was conducted on the % carbon data to determine if the differences in the mean values between depths (groups) were significant. The differences in the mean values amongst the six treatment groups (N= 10 each) were statistically significant (P<0.001) at α = 0.05 (F = 451.5; df = 59). In a pairwise multiple comparison procedure (Holm-Sidak method), 13 of the 15 comparisons were statistically significant

(P<0.001) with two (depth 33 cm versus 30 cm and depth 30 cm versus 115 cm) being non-significant (P = 0.056 and 0.341 respectively).

Table 4: Percent carbon and nitrogen from six samples with replicates that were used to test the homogeneity of the peat matrix. The mean and standard deviation (stddev) have been calculated for each sample and its replicate. Sample codes are the core name (VC01) and depth in cm. A and B replicates are shown for each level.

SAMPLE C (%) N (%) SAMPLE C (%) N (%) SAMPLE C (%) N (%)

VC01 15(A) 43.669 2.552 VC01 30(A) 47.761 3.094 VC01 33(A) 47.539 3.085

43.961 2.556 47.797 3.069 47.688 3.026

43.69 2.568 47.813 3.077 47.638 3.072

43.707 2.56 47.592 3.077 47.73 3.064

44.538 2.595 47.702 3.067 47.224 3.07

VC01 15(B) 44.054 2.817 VC01 30(B) 47.163 3.07 VC01 33(B) 47.966 3.066

44.107 2.846 47.388 3.063 47.552 3.071

44.254 2.833 47.283 3.067 47.848 3.04

44.238 2.83 47.506 3.051 48.059 3.047

44.224 2.802 47.096 3.068 47.968 3.057

MEAN 44.0442 2.6959 47.5101 3.0703 47.7212 3.0598

STDDEV 0.2884 0.1376 0.2664 0.011 0.2512 0.0175

SAMPLE C (%) N (%) SAMPLE C (%) N (%) SAMPLE C N (%)

VC01 65(A) 48.414 2.601 VC01 94(A) 45.274 2.603 VC01 115(A) 47.42 2.587

48.739 2.586 45.412 2.641 47.357 2.584

48.649 2.574 45.459 2.624 47.307 2.58

48.748 2.566 45.628 2.635 47.32 2.585

48.229 2.572 45.914 2.636 47.409 2.565

VC01 65(B) 48.542 2.657 VC01 94(B) 45.634 2.544 VC01 115(B) 47.466 2.514

47.986 2.645 45.806 2.545 47.635 2.494

48.297 2.641 45.664 2.549 47.425 2.508

48.204 2.635 45.815 2.537 47.254 2.495

48.015 2.631 45.793 2.541 47.471 2.491

MEAN 48.3823 2.6108 45.6399 2.5855 47.4064 2.5403

STDDEV 0.2815 0.0346 0.2044 0.0458 0.1072 0.0429

Percentage carbon for the peat core (Fig. 8) begins low in the basal peat (242-240 cm depth), and then increases, interrupted by a decrease between 211 cm and 196 cm, until a depth of 36 cm. At this point, values decline rapidly toward the surface of the peat.

This zone of rapid decrease corresponds to the acrotelm. The average percentage carbon was found to be 45.9%, which is lower than values found by Gorham (1991), Turunen et al. (2001), and Vitt et al. (2000) for other boreal peatlands (51.7%, 52.7% and 51.8% respectively) and these published estimates also include values from complete peat sequences, including the acrotelm.

100

Depth(cm) 150

200

36 38 40 42 44 46 48 50 % Carbon

Figure 8: Percent carbon in the peat sequence of the Victor fen core with errors given as ± 1 standard deviation.

Percentage nitrogen (Fig. 9) varies from 1.19 - 3.07%. Relatively stable values through most of the core are interrupted by a decrease of over half in % nitrogen after a depth of 226 cm (2.46) to the low of 1.19. After this low point, values recover and stay high until a depth of 30 cm when they begin to decline towards the top of the core,

58 indicating passage through the acrotelm. High N percentages suggest an increased amount of peat decay, as N becomes concentrated in peat at an inversely proportional rate to that of organic matter loss (Belyea and Warner 1996). The average percent nitrogen value of 2.39 is close to the average value of the intermediate fens (2.52%) studied by Bridgham et al. (1998) and close to the mean given for sedge peats (2.2% and

2.1%) from Northwestern and Northeastern Ontario by Riley and Michaud (1989) and

Riley (1989).

100

Depth(cm) 150

200

1.0 1.5 2.0 2.5 3.0 3.5 % Nitrogen

Figure 9: Percent nitrogen in the peat sequence of the Victor fen core with errors given as ± 1 standard deviation.

The carbon:nitrogen ratio (Fig. 10) varies between 15.19 and 36.60. The drop in nitrogen after 226 cm depth is shown by a large increase in C:N from 18.38 at 226 cm depth, to 36.60 at 205 cm depth. The ratios exhibit an increase in the uppermost peat from 18 cm to the surface, as the proportion of N in the uppermost peat declines more rapidly than the proportion of carbon. Low C:N ratios when combined with high N

59 proportions suggest an increase in the amount of peat decay at that level (Borgmark and

Schoning 2006; van der Linden and van Geel 2006). This can clearly differentiate the acrotelm from the catotelm in the stratigraphy.

100

Depth (cm)Depth 150

200

10 15 20 25 30 35 40

C:N Ratio

Figure 10: Carbon:Nitrogen ratio of the peat sequence of the Victor fen core

3.6 LORCA

LORCA estimates (Fig. 11) for the Victor fen core ranged between 14.54 and 196.59 g C m-2 year-1. The average LORCA for the 81 sample depths was 49.85 g C m-2 year-1. The sharp rise in LORCA estimates beginning at approximately 65 cm depth are due to fast peat accumulation near the surface and an incomplete decay process (van der Linden and van Geel 2006). Estimates in the deeper peat (60 cm to base) ranged between 14.54 and

55.98 g C m-2 year-1 (Fig. 12), with an average of 24.59 g C m-2 year-1. The emphasis of the sharp rise being a recent phenomenon (1/10th of the record when plotted against the

60 age-depth model versus 2/5ths when plotted by depth) is clearer when LORCA is plotted against age of the peat deposit (Fig 13), as the recent upward trend surpasses the highest

LORCA value for the preceding 6000 years approximately 600 years BP.

100

Depth (cm) Depth 150

200

0 50 100 150 200 -2 -1 LORCA (g C m year )

Figure 11: LORCA estimates for the entire peat sequence of the Victor fen core

100

150 Depth (cm) Depth

200

10 20 30 40 50 60

-2 -1 LORCA (g C m year )

Figure 12: LORCA estimates for the 60 cm to base section of the Victor fen core

1000

2000

3000

4000 Age (cal Year BP) Year Age(cal 5000

6000

0 50 100 150 200 -2 -1 LORCA (g C m year )

Figure 13: LORCA estimates for the Victor fen core based on the age-depth model

3.7 Peat Humification

Raw spectrophotometric humification absorbance values were detrended using a cubic polynomial (Figs 14 and 15). This function had the best fit (highest r2) to the data compared to a linear function, and a quadratic function. The equation that was fit is represented by the function f= 33.0394+ 0.0119 * X + -4.4880e-0062 + 5.3359e-010 * X

3 with an r2 of 0.6401. The raw humification absorbance values (Fig. 14) increase by approximately 30% with depth in the uppermost 50 cm of the peat profile indicating a passage from the acrotelm to the catotelm. In the next 150 cm sequence, absorbance values oscillate between increases and decreases of upwards of 18% (error calculated as

2.35% based on replicate differences) in response to changing humidity at the surface. In the lowermost 35 cm of the profile analyzed, absorbance increases 20% to a maximum before decreasing slightly at the lowest depth. The detrended absorbance values (Ad)

(Fig. 15) also clearly show the passage from the acrotelm to the catotelm and vary by a maximum of 26% absorbance in the next 150 cm of peat. The final decrease in absorbance in the final 15 cm is also evident in the detrended values, as the cubic line passes above the raw values at these depths. Both datasets are useful to discuss because the raw data exhibit a depth-dependent increase in humification (due to anoxic decay in the catotelm) while the detrended absorbance values remove this trend but yield instances where the function does not fit the raw data as closely because other factors are impacting the concentration of humic material in the peat (the climatic signals of interest).

For the 42 samples shared between peat humification and the carbon:nitrogen data, significant correlations exist. The detrended absorbance values (Ad) were

63 significantly correlated (P<0.001) with percent carbon data (r = 0.625) and percent nitrogen data (r = 0.539). The raw humification values were also significantly correlated

(P < 0.001) with percent carbon data (r = 0.759) and percent nitrogen data (r = 0.518).

This suggests that in addition to a peat-forming vegetation proxy, C:N ratios are useful to make inferences about decomposition at the Victor fen site. Significant correlations between C, N and humification in peat were also found by Borgmark and Schoning

(2006), and qualitative correlations were found by Mauquoy et al. (2002a). Significant negative correlations were also found between the LORCA values and the detrended absorbance values (Ad) (r = -0.313; P <0.05) as well as between the LORCA values and the raw humification results (r = -0.726; P < 0.001). As the LORCA values are derived from the percentage carbon data, this result is understandable. Interestingly, the raw humification results were better correlated to the percentage carbon data and the LORCA data than the detrended results. However, detrending is necessary because of the depth- dependent effect evident in the humification results (steadily increasing humification with depth).

100

20 Spectrophotometric Absorbance (%) Absorbance Spectrophotometric 0 0 2000 4000 6000 Age (cal Year BP)

Figure 14: Raw spectrophotometric absorbance results for the Victor fen core. The solid grey line is the best fit cubic polynomial that was used to detrend the absorbance values. ) d

-10

-20 Detrended Absorbance Residual (A Residual Absorbance Detrended

0 2000 4000 6000

Age (Cal Year BP)

Figure 15: Detrended absorbance values (Ad) for the Victor Fen core joined by a smoothed line, together with the calculated error.

3.8 Paleoclimatic Reconstructions

Four climate variables were reconstructed using the modern analog technique (MAT) and the pollen assemblages from the Victor fen peat core: average annual temperature (°C), mean July temperature (°C), total annual precipitation (mm), and average June, July,

August (JJA) precipitation (mm). Critical values for squared chord distances (<0.12) were met for all but one target fossil sample and the calibration dataset, suggesting that adequate modern analogs existed for all of the remaining fossil samples (Huntley 1996;

Overpeck et al. 1985). The third closest modern analog found for the basal salt marsh sample did not meet the critical value for squared chord distance (0.13), and the two closest modern analogs for the basal sample had values that were at or near the cutoff

(0.118 and 0.12). This depth was dominated proportionally by Typha (81%), which is not included in the Modern Surface Pollen Database (most aquatic taxa are excluded). The remaining assemblage for this one sample in question (244 cm depth) is likely biased because of this (as insufficient numbers of the remaining taxa may have been counted), so the high dissimilarity is unsurprising.

Two graphs per reconstruction are shown to illustrate both the whole record (0-

6700 yrs BP) and the past 2000 years in greater detail. The reconstructions appear in

Figures 16-23. The r2 of the reconstructions were 0.543, 0.523, 0.307, and 0.441 for average annual temperature, mean July temperature, total annual precipitation, and total

June, July, August (JJA) precipitation respectively. These r2 values are referring to the goodness of fit between the observed and predicted values of the modern dataset. The root mean squared errors (RMSE) of the reconstructions were 5.4 °C, 3.8 °C, 373.6 mm and 26.8 mm respectively. The RMSE are included in the figures. RMSE is an absolute

66 measure of model error scaled to the units of the reconstructed environmental variable, which can make inter-variable comparisons difficult (Williams and Shuman 2008).

Reliable reconstructions should have a high precision, meaning a high r2 value and a low

RMSE (Williams and Shuman 2008). Using r2 only to assess differences in precision among the same environmental variables, it suggests that average annual temperature was a slightly more precise reconstruction than average June, July, August temperature, but that total June, July, August precipitation was much more precise than total annual precipitation. The temperature reconstructions appear to be closer, but it is difficult to make inter-comparisons between both temperature and precipitation variables. The

RMSE of the reconstructions are large, and many (but not all) of these reconstructed changes are within the error specified.

The temperature reconstructions both begin (circa 6775 years BP) with an anomalously high value (7.6 °C and 21.5 °C for annual and July average temperatures respectively compared to the reconstructed averages of 0.77 °C and 16.45 °C and the modern values of -1.3 °C and 17.2 °C for Lansdowne House) (Environment Canada

2011) that immediately declines to the minimum values (-3.59 °C and 11.37 °C), a decrease of 10 °C. This decrease is followed by a slight upward trend in temperature that is punctuated by short periods of both cooler and warmer conditions. There is a large increase in temperature reconstructed for the most recent part of the record (circa 20 years BP), followed by a final decrease in the top two spectra (including the surface). The precipitation reconstructions also begin with high values at the base, with the highest reconstructed June, July, August total precipitation of 118.5 mm (mean of reconstructions

89.93 mm; modern value 291 mm at Lansdowne House) (Environment Canada 2011)

67 declining to a low value immediately much like the temperature reconstructions. A prolonged period of low precipitation is identified between 3500 and 2700 years BP, which corresponds to a period of low temperature. This cold, dry period is followed by a rise to higher precipitation that is especially pronounced in the total annual precipitation reconstruction (an increase to ~480 mm). The reconstructed precipitation for the two most recent samples gives low precipitation values relative to the Holocene average of

505 and 778 mm for total annual (mean of reconstructions 876 mm; modern value 699.5 mm at Lansdowne House) (Environment Canada 2011) and 47 and 65 mm for total June,

July, August precipitation. The comparisons between the reconstructions and the climate normals for Lansdowne House suggest that the reconstructions underestimate precipitation during the growing season, and overestimate precipitation over the entire year. This may suggest that the modern climatic regime (1971-2000) of the area is different from conditions that were prevalent during other periods of the preceding 6775 years, or that the pollen assemblages are not best suited to reconstruct this variable.

-5 Average Annual Temperature(°C) Annual Average

-10 0 2000 4000 6000

Age (cal Year BP)

Figure 16: Average annual air temperature inferred for the Victor fen site from the fossil pollen data with the root mean squared error of the reconstruction.

-5 Average Annual Temperature (°C) Average

0 500 1000 1500 2000

Age (cal Year BP)

Figure 17: Average annual air temperature of the last 2000 years for the Victor fen site as inferred from the fossil pollen data with the root mean squared error of the reconstruction.

10 Average July Temperature (°C) July Average

0 2000 4000 6000

Age (cal Year BP)

Figure 18: Average July temperature of the Victor fen site as inferred from the fossil pollen data with the root mean squared error of the reconstruction.

Average July Temperature (°C) July Average 10

8 0 500 1000 1500 2000

Age (cal Year BP)

Figure 19: Average July temperature of the last 2000 years of the Victor fen site as inferred from the fossil pollen data with the root mean squared error of the reconstruction.

1400

1200

1000

800

600

400 Total Annual(mm) Precipitation Total 200

0 2000 4000 6000

Age (cal Year BP)

Figure 20: Total annual precipitation of the Victor fen site as inferred from the fossil pollen data with the root mean squared error of the reconstruction.

1400

1200

1000

800

600

400

Total Annual(mm) Precipitation Total 200

0 0 500 1000 1500 2000

Age (cal Year BP)

Figure 21: Total annual precipitation for the last 2000 years of the Victor fen site as inferred from the fossil pollen data with the root mean squared error of the reconstruction.

140

120

100

40 Average JJA Precipitation (mm) Precipitation JJA Average

20 0 2000 4000 6000

Age (cal Year BP)

Figure 22: Average precipitation for June, July, August for the Victor fen site as inferred from the fossil pollen data with the root mean squared error of the reconstruction.

140

120

100

40 Average JJA Precipitation (mm) Precipitation JJA Average

20 0 500 1000 1500 2000

Age (cal Year BP)

Figure 23: Average June, July, August precipitation for the last 2000 years for the Victor fen site as inferred from the fossil pollen data with the root mean squared error of the reconstruction.

4 Chapter 4 DISCUSSION 4.1 Drivers of Vegetation Change

The pollen record from the Victor fen provides a clear picture of peatland development in the face of autogenic and allogenic forcing factors. The pollen assemblage in the uppermost marine sediment is characterized by high values of Pinus banksiana,

Chenopodiaceae, Cyperaceae and Salix. The pollen of P. banksiana is more regional than the others (it is predominant in forests south and west of the HBL), but northern populations do occur near the fen site today on drier river banks and on glaciomarine and glaciolacustrine beach ridges (McAndrews et al. 1982; Riley 2003). As the Attawapiskat

River has entrenched 30 metres into the underlying limestone and formed cliffs in the vicinity of the fen (Cowell 1983), conditions are conducive for growth of P. banksiana.

The low pollen influxes for the more local tree and shrub species indicate that much of the adjacent area was itself still below sea level at this time. Thus, at the time of the

Tyrrell Sea transgression, the pollen of P. banksiana would have originated from further away than today. The high values of Salix in the earliest portion of this record suggest that the location was adjacent to areas with an emerging substrate with a high pH (basic)

(McAndrews et al. 1982). The Chenopodiaceae pollen also supports the idea that this environment was both alkaline and brackish (Martini et al. 1980). The high values of

Chenopodiaceae may also be indicative of more disturbed conditions during the final stages of the Tyrrell Sea, similar to conditions found by Glaser et al. (2004a) at another site in the Hudson Bay Lowlands.

76 77

Following the retreat of the Tyrrell Sea transgression, the semi-emergent site became a salt marsh dominated by Typha latifolia and Cyperaceae. The hyper-abundance of Typha indicates that the waters had become much less saline than those covering the uppermost marine sediment. The radiocarbon chronology suggests the salt marsh stage lasted ≤ 100 years before the vegetation community changed to a shrub-fen assemblage.

The rapidity of change between the basal marsh zone and the establishment of the forest- shrub fen suggests a strong role for the rapid rate of isostatic uplift (between 2.65 and 3 metres century-1) in driving vegetation change (Andrews and Peltier 1989); the uplift separated the site from the waters of the retreating Sea at a rate likely too fast for autogenic factors to exert a strong influence (Glooshenko and Martini 1983). A time frame of 100-200 years for a salt and freshwater marsh stage before the influence of tidal waters is negated due to uplift, thus transforming the marshes into a fen ecosystem, was suggested by Glooshenko and Martini (1983) based on the rate of seaward advancement of marshes due to glacio-isostatic rebound. This marsh successional stage is comparable in terms of time span (~100 years) and in terms of abundance of Cyperaceae pollen to the basal salt marsh stage found by Glaser et al. (2004a) in three different raised bogs in the

Albany River Watershed, and also comparable in terms of some pollen taxa present to the salt marsh found by Terasmae and Hughes (1960). However the hyper-abundance of

Typha has not been reported before in palynological studies of the HBL. The growth and accumulation of autochthonous peat during the salt marsh stage supports the idea that the peatland genesis was spontaneous with land emergence, rather than being driven by lake- infilling or paludification, as was found in the Albany River basin by Glaser et al.

(2004a).

At approximately 6 300 years BP the dominant vegetation community changed to a poor fen community centered on Picea mariana, Sphagnum moss, Betula pumila,

Betula papyrifera, which was much less abundant than B. pumila (the two were amalgamated together), Alnus, and the continued local-regional contribution of Pinus banksiana. This assemblage includes smaller proportions of Picea glauca, Salix, shrubs in the Ericaceae, Larix laricina and Equisetum. This fen stage is similar in terms of the species represented to the assemblages found by Glaser et al. (2004a), Kettles et al.

(2000), Klinger and Short (1996), and Terasmae and Hughes (1960).

The fen community remained relatively stable (with oscillations in the assemblages reflecting both autogenic and allogenic signals) until approximately 100 BP, with a rise in Alnus, Cyperaceae, Ambrosia and Chenopodiaceae and a decline in

Sphagnum. The recent decline in Sphagnum and the rise in Cyperaceae as well as the rise in shrub cover of Alnus and Betula suggests drier conditions at the site, perhaps equivalent to the recent drier conditions of Kinosheo Lake Bog (Klinger et al. 1994;

Klinger and Short 1996). This drying is supported by the fall in total annual and total

June, July, and August precipitation in the most recent samples reconstructed from fossil pollen assemblages (Figs 21 and 23). Repeat photography has indicated that there has been considerable shrub (Alnus, Salix, and Betula) expansion in northern Alaska since

1945 (Tape et al. 2006). Combined with plot and remote sensing evidence from Canada,

Scandinavia and portions of Russia indicating shrub expansion, Tape et al. (2006) suggest a Pan-Arctic shrub expansion in response to recent climate warming. The expansion of shrub cover at the Victor fen may be part of this Pan-Arctic shrub expansion.

The rise in Ambrosia and Chenopodiaceae in the more recent peat is indicative of regional anthropogenic land disturbances, although these disturbances may be taking place at distances of several 100 km away from the study site (Klinger and Short 1996).

The peat profile does not show the succession from a fen to a bog that was documented by Glaser et al. (2004a) Kettles et al. (2000), Klinger and Short (1996) and Terasmae and

Hughes (1960). The persistence of a fen-type system is supported by the lack of the bog- indicating Rubus chamaemorus, and the low proportions of Ericaceae (maximum 2% versus upwards of 40-50% found by Glaser et al. (2004a), Klinger and Short (1996) and

Terasmae and Hughes (1960)) and the fact that nitrogen remains above 2% dry mass following the Sphagnum peak until values decline in the acrotelm (Bridgham et al. 1998;

Glaser et al. 2004a). These indicators suggest that the minerotrophic influence on the

Victor fen site has not disappeared completely, as would be the case in the transition to a bog recorded in the studies cited above. They also suggest that in this case the fen is the

“climax community” for the site (defined by long-term structural and compositional stability) if its perceived resilience is upheld, much like a bog was a climax community for the Kinosheo Lake site (Klinger and Short 1996). However, if the peat sequence continues to deepen with accumulated sediment, the minerotrophic waters of the local drainage could become cut off from the surface vegetation, limiting certain nutrients and eventually causing it to change to an ombrotrophic bog (Glaser et al. 1997).

The progressive accumulation of autochthonous sediment is an autogenic forcing, but it is modulated to an extent by allogenic factors, through the influence that climate and hydrology have on peat accumulation. Overall, the pollen assemblages of zones 2 and 3 are very similar to the contemporary surface pollen spectra of nine fens studied in

80 the southwestern James Bay Lowlands by Farley-Gill (1980). Surface sediments of those fen sites are dominated by pollen of Picea, Cyperaceae, Sphagnum, and Betula, with lesser amounts of Alnus, Pinus, Larix, and Ericaceae (Farley-Gill 1980).

Subtle changes in the pollen stratigraphy within the fen stage (Zone 2) include oscillations in the percentage of Sphagnum spores (with five sequences of increasing-to- decreasing proportions) suggesting changing hydroclimatic conditions in the fen.

Increases in the proportions of Picea mariana, Larix, Betula, and Alnus represent a change from a more open fen environment, to a more forested and shrub fen site, in response to lower surface moisture availability at the site. Similar peaks in Betula found by Terasmae and Hughes (1960) also represented temporary shifts towards a more shrub- dominated fen. The likely incidence of a Betula catkin falling on the surface of the peat and getting enveloped by the surface vegetation at depth 226 cm supports the assertion of greater shrub coverage at the peatland surface.

Long term floristic and ecological changes evident in stratigraphic peat layers indicate allogenic and autogenic succession and can be attributed to both external

(climate) and internal (local environmental conditions) forcings (Payette 1988).

Autogenic processes can be identified after the external forcings and the associated vegetation changes have been identified. The oscillations in Sphagnum and Picea proportions in the Victor fen record (Fig. 5; see depths 194, 181, 161.5-155, 129, 121.5,

64 and 13-7 cm.) suggest cyclic, self-perpetuating vegetation dynamics indicative of autogenic forcing, much like the patterns found in ombrotophic bogs in Northern Quebec by Payette (1988). Payette (1988) found alternating layers (macrofossils remains and wood) of Sphagnum-Picea that persisted for multiple millennia. The author contended

81 that the ecological traits of slow growth and stunted forms of Picea were important in forcing the long-term cyclic replacement of Sphagnum as follows: Picea maintains the growth of Sphagnum by facilitating increased snow accumulation (and thus, water and insulation), and Sphagnum is a suitable medium for Picea layering and thus the two species maintain a mutualistic relationship, provided the ecological threshold (the point at which conditions are no longer suitable for growth and survival) of either species is never reached (Payette 1988).

The growth of established Sphagnum has been found to intensify local acidification processes (Glaser et al. 1981; Kuhry et al. 1993) resulting in enhanced autogenic change. Precipitation is scarce in metal ions in non-maritime areas, leading to strong acidic reactions of both peat and water, resulting in the formation of some organic acids (Sjörs 1959). As Picea is never extirpated from the Victor fen site (based on the presence of the pollen throughout the profile, Fig 5), it is possible that the oscillations are a direct response to the cyclic variation in Sphagnum. The decline of Picea with increasing Sphagnum suggests that with higher a proportion of Sphagnum, the peatland surface is more saturated, resulting in conditions less conducive for the growth of arboreal species such as Picea.

This interpretation is supported by work from the Red Lake Peatlands in

Minnesota which suggests that in wetter channels of peatlands, the growth of Sphagnum is favoured at the expense of forest cover (Glaser et al. 1981). The Red Lake Peatlands contain mire complexes similar to those described by Sjörs (1963) for the Hudson Bay

Lowlands, especially the vegetation assemblages of the bogs and fens, suggesting this pattern may be similar. Once the ratio becomes too close to the ecological threshold

(either too abundant in Sphagnum or Picea), there is a response towards a more balanced proportion of the two forms. The cyclic succession of Victor fen is a special case according to Sjörs (1963), as the author claimed that minerotrophic peatlands do not normally exhibit cyclic succession. The high resolution of the Victor fen study may have been able to capture these changes more completely than past work. However, there are concurrent peaks in Picea and Sphagnum influx at 20, 55 and 97 cm depths suggesting that this relationship may not always be inversely proportional.

Mosses are excellent indicators of local peatland conditions (Kuhry et al. 1993) and are therefore useful to track autogenic factors. The large increase in Sphagnum spores from 5 900 to 4 900 years BP is also captured by the decline in percent nitrogen and the increase in the carbon:nitrogen ratio, as mosses have a relatively high C:N ratio

(compared to vascular plants) (Kuhry and Vitt 1996). The low values of N are in contrast to the higher percentage of N found elsewhere in sedge-dominated peat (Aerts et al.

1999). Low N values and high C:N ratios also mean less peat decay as Sphagnum decays less readily than other vegetation forms due to a high concentration of decay-resistant compounds, and waterlogged conditions (Aerts et al. 1999).

The change to more Sphagnum-dominated peat suggests a local environmental signal at this time; perhaps a decrease in minerotrophic inputs and the resulting acidification provided conditions conducive for the growth of Sphagnum. Nitrogen values for other northern wetlands support this change to a more acidified fen (Bridgham et al. 1998). A local environmental change facilitated the increase in Sphagnum mosses.

Once the mosses began to increase, autogenic changes resulting from the acidification promoted by the mosses themselves followed, eventually eclipsing the allogenic

83 influence. These autogenic changes included the further expansion of Sphagnum and potentially the arrival of new Sphagnum species. However, because the Sphagnum spores were only identified to genus rather than species, the coarser taxonomic information may be obscuring the environmental changes that occurred at this time, as the different

Sphagnum species respond to environment signals differently. For example, Sphagnum fuscum peat has low C:N ratios (intermediate between other mosses and vascular plants) as it typically contains the more decayed, finer fractions of peat that are enriched in nitrogen (Kuhry and Vitt 1996). Peat humification values between 5 950 and 5 070 years

BP also decrease suggesting less decomposition, due to the development of bog vegetation (more likely given the large changes in Sphagnum and C:N ratios) or the climate signals of either increased precipitation or decreased temperature.

4.2 Climate Reconstruction

As climate is one of the dominant forcing factors influencing the long term vegetation dynamics in peatlands (Chambers and Charman 2004), climatic signals should be inherent in the changing pollen assemblages even if autogenic factors are acting upon the system. The climate signal, however, may be more readily seen in the regional pollen assemblages compared to the local pollen assemblages, which are directly influenced by autogenic changes taking place within the peatland. The assemblages as a whole were used to derive the climate reconstructions, although the returned errors are high. Thus, the actual values of the reconstructed temperatures and precipitation (Figs 16-23) need to be interpreted with caution.

The basal fossil assemblage (circa 6800 years BP) resulted in a reconstruction with high average annual and mean July temperature (7.5 and 21 °C respectively), and

84 high June, July and August total precipitation (118 mm) relative to the succeeding time period. The high temperatures may be indicative of the Holocene Thermal Maximum, which initiated in the Hudson Bay region approximately 7-6 000 years BP (Kaufman et al. 2004). However, this interpretation is highly tentative as it is only one sample.

Conditions indicative of the Holocene Thermal Maximum have also been found in

Northern Quebec (circa 5780 cal year BP), suggested by a range extension of white pine

(based on fossilized wood, cones, leaves and a higher proportion of pollen) 100 km north of the present range (Terasmae and Anderson 1970). HTM conditions in the Clay Belt of

Northern Ontario have also been shown by an increase in Thuja between 6000 and 4500 years BP, marking the expansion of Thuja in lowland habitats due to warm and dry conditions (Liu 1990). Paleoclimate simulations have also indicated that July temperatures circa 6000 years BP (14C age) were warmer than the present day throughout

North America (Bartlein et al. 1998).

However, conditions change rapidly in the next fossil assemblage in the Victor fen record, with average annual and mean July temperatures falling to -3 and 11°C respectively, and total JJA precipitation to 68 mm. The rapid change is highly unlikely to be a true indicator of temperature and thus suggests that the upper sequence of marine sediments has been reworked and some of the pollen grains redeposited. McAndrews and

Campbell (1993) also reconstructed anomalously high temperatures at the base of the R

Lake record, a time of Tyrrell Sea coverage; the authors attributed that high temperature not to an actual high temperature, but to an assemblage composed of pollen grains recycled from other deposits. Palynological analyses of more samples from the marine

85 sediment at this base of the Victor fen core are needed to confirm that reworked pollen assemblages explain the anomalous basal assemblage.

The prolonged period of predominantly low total precipitation between 6775 and

3000 years BP reconstructed from the Victor fen record (Fig. 20) may serve as better evidence for the HTM. Changes in the status of lakes in North America has indicated many lakes exhibited drier to much drier conditions circa 6800 years BP (Wanner et al.

2008) and this period also coincides temporally to the HTM inferred by McAndrews et al. (1982) based on qualitative analysis of pollen taxa and macrofossils in the R Lake core.

The inferred decline in average annual temperatures between 3500 and 2800 years BP in the Victor fen record (Fig. 16) may be indicative of Neoglacial cooling.

Neoglacial cooling has been suggested by other work for a comparable time period in other regions of the HBL (McAndrews et al. 1982; Kettles et al. 2000), Northern Ontario

(Liu 1990), and for Northern Quebec (Filion 1984; Gajewski et al. 1993; Payette 1988).

There is a local low in reconstructed July average temperature from R Lake (~14 °C versus ~15 °C for the Victor fen) centered around 3000 years BP, which may indicate that the decline found in the Victor fen record is part of a more regional cooling trend

(McAndrews and Campbell 1993). Cooling is also suggested by the rise in Sphagnum and decline in Picea beginning approximately 2500 years BP found by Klinger and Short

(1996).

After 2700 – 3000 years BP, the precipitation reconstructions indicate an increase in precipitation over conditions reconstructed for 6775 to 3000 years BP (average of 925

86 mm versus 810 mm annual precipitation) (Fig. 20). The reconstructions suggest there are two distinct precipitation regimes: a period of low precipitation pre 3000 years BP and a period of higher precipitation post 3000 years BP. This increase in wetness may better indicate conditions of the Neoglacial than the <1000-year decrease in temperature.

However, this change in precipitation is still less than the range of the RMSE of the reconstructions, so the reconstructions must be interpreted with caution. Payette and

Filion (1993) studied a subarctic lake east of Hudson Bay in Northern Quebec and found low lake levels between 5400 and 3500 years BP and a predominantly high water level from 3500 years BP to present, suggesting greater precipitation and effective moisture.

Ali et al. (2009) found spatially heterogeneous fire history patterns from four lakes south of James Bay in central Quebec beginning approximately 4000 cal year BP. Prior to this time, synchronous fire episodes were identified at all sites, with the frequency predominantly controlled by climate. The authors attributed this pattern to an increased moisture regime and a rising water table that resulted from the cooler and moister conditions of the Neoglacial with a new local weather regime established that impacted fire ignition, propagation and extent (Ali et al. 2009). Cool and moist conditions beginning approximately 3900 years BP were also suggested by Garralla and Gajewski

(1992) with the increase of Picea, Sphagnum and Ericaceae at the expense of Betula in a lake near Chibougamau of central Quebec. The wetter conditions expressed in the Victor fen reconstructions may therefore be part of a more regional signal of increased moisture.

The rise in annual air temperature between 1150 and 800 years BP (to between 2 and 3 °C, from the preceding low point of 0.25 °C centered around 1390 years BP) may correspond to the Medieval Climate Anomaly, a climatic warming that began 1650 years

BP that was geographically widespread in North America and the Northern Hemisphere

(Crowley and Lowery 2000; Mann et al. 2009; Viau et al. 2002). However, the data from the Victor fen reconstruction is variable and this period of relatively high temperature is well within the other peaks of the last 2000 years. A small rise in reconstructed average

July temperature was found from the R Lake record beginning approximately 1000 years

BP, suggesting that this warming may have been a more regional climatic change

(McAndrews and Campbell 1993). Precipitation increases to a local maximum between

950 and 800 years BP reflected in the low values of humification, punctuated by a high value circa 930 years BP. This single high value may be reflecting the increase in temperature, rather than the precipitation signal.

The decrease in inferred annual and July temperatures of between 4-5 °C and 2-3

°C respectively in the Victor fen record, beginning 500 years BP and lasting until 360 years BP, may correspond to the Little Ice Age, a well known period of climatic cooling in the Northern Hemisphere (Mann et al. 2009; Moberg et al. 2005; Wanner et al. 2008).

McAndrews and Campbell (1993) also reconstructed low average July temperatures for

R Lake at this time period, but the decline was only in the order of ~0.5 °C. Cooler conditions are supported by low values for humification at this time (a decrease of 10% absorbance at this time) (Fig 15). Climate deteriorations (cooling) in European bogs approximately 500 and 350 years BP have been linked to variations in solar activity

(inferred from shifts in 14C and 10Be isotopes which are modulated by the solar wind), the

Spörer and Maunder sunspot minima respectively (Beer et al. 2000; Mauquoy et al.

2002b). The cooling recorded in the peat record supports climate as a driver of peatland

88 dynamics, and also suggests that peatlands that are very distant from each other respond similarly to wide-scale climatic events.

Between 500 to 350 years BP, there is a concurrent decrease in precipitation (both annual and June, July, August total, to a minimum of 680 and 70 mm respectively) and decrease in temperature. These results suggest drier, warmer conditions, similar to conditions found at Lac Le Caron by Loisel and Garneau (2010) as well as to lower lake levels found at 300 BP by Payette and Filion (1993). Once temperatures and precipitation increase from the minimum values at 360 years BP, humification values increase to the maximum found (13.87%) at 280 years BP. The increase in humification, however, is due to the complexity of the record in the acrotelm and should not be interpreted as a climate signal.

The climate reconstructions are influenced by the inclusion in the dataset of both locally-influenced pollen and regionally-influenced pollen. The more regional pollen

(especially Pinus) has a large influence on the climate reconstructions because of the inclusion of depositional environments favourable to these long-distance transported taxa

(especially lakes). However, the reconstructions are influenced from the local signals of

Sphagnum, Cyperaceae, Betula, Alnus, and Picea mariana. While the temperature reconstructions are heavily influenced by the regional pollen, the precipitation reconstructions are more influenced by the locally derived species in the assemblage.

This is because the locally influenced pollen is derived from species on the surface of the peatland, and these species are sensitive to changes in the moisture regime at the surface

(especially Sphagnum and Picea).

Peat humification results return a signal that is a derived response to climate: the surface wetness/humidity of the mire in question. The degree of humification depends on the assumption that the major peat-forming botanical remains have been relatively homogeneous throughout the peat profile, and therefore the influence of botanical composition cannot be discounted (Caseldine et al. 2000; Yeloff and Mauquoy 2006).

Corresponding vegetation and humification “shifts” have been found in other peatlands

(Chambers et al. 1997), suggesting that the alkali extraction method to measure humification should only be used when a vegetation proxy is also employed. As there is no evident change in the dominant species of the pollen assemblage at any depth until zone 3, a change in botanical composition of the peat forming vegetation seems unlikely.

However, the exception to this is the Sphagnum peak between 214 and 194 cm. During this period, the humification results decline from a high point at 210 cm, to a low between 204 and 198 cm before rebounding at 195 cm. Caution should be used in interpreting these results in relation to decomposition, because the humification data follows the Sphagnum peak so closely. The humification trends for the rest of the peat sequence are understood to be reflecting only temperature and moisture at the surface of the peatland.

4.3 Controls on Carbon Accumulation Dynamics

Carbon accumulation follows a similar trend to that of peat accumulation as suggested by the age-depth model, with a period of relatively high accumulation in the basal section, followed by a long period of low accumulation, followed again by a period of high accumulation. Rates are high from 6 500 to 5 800 years BP, then decline and are low between 5 700 and 1 850 years BP, and finally rise to their highest levels from 1 730

90 years BP to the present. This high rate for the late Holocene is due to fast peat accumulation near the surface and an incomplete decay process, as indicated by van der

Linden and van Geel (2006). The average rate of carbon accumulation decreases with increasing time as the peat is exposed to leaching and further anoxic decay in the deeper peat layers (Clymo 1984; Tolonen and Turunen 1996; Turunen et al. 2002).

The average LORCA for the entire peat sequence of 49.85 g C m-2 year-1 is high compared to published estimates from other northern peatland regions (various peatland types including both fens and bogs), such as Finland (22.5 g C m-2 year-1; 17.3 to 26.1 g

C m-2 year-1) (Tolonen and Turunen 1996; Turunen et al. 2002), West Siberia (17.2 g C m-2 year-1; 3.8 to 44.1 g C m-2 year-1) (Beilman et al. 2009; Turunen et al. 2001), the former Soviet Union (30 g C m-2 year-1) (Botch et al. 1995), and for a mix of boreal and northern peatland sites (23 g C m-2 year-1 and 18.6 g C m-2 year-1) (Gorham 1991; Yu et al. 2010). The average value for the deeper peat in the Victor fen core (>60 cm) of 24.6 g

C m-2 year-1 is closer to these previous estimates. As stated by Tolonen and Turunen

(1996), LORCA can be estimated with dry bulk density, carbon content and age, but only for the deepest peat layers. However, the accumulation rates of only fens and marshes in the former Soviet Union of upwards of 72-80 g C m-2 year-1 (Botch et al. 1995) is very large compared to the other estimates, and more consistent with the estimates of the

Victor fen in the uppermost 50 cm of the peat sequence.

The high estimates for LORCA from the Victor fen core could be due to error in the calculation of bulk density of the Victor core. The over-estimate of bulk density could have happened through attempting to force the peat subsample into the 1 cm3 syringe top, thereby adding more peat for 1 cm3 than the true amount. The average bulk density of

0.157 grams cm-3 is greater than the average of 0.112 grams cm-3 given by Gorham

(1991). A greater bulk density would yield a greater carbon mass, and therefore a higher

LORCA value. By contrast, the lower proportion of carbon (45.9% versus 51.7%, 52.7% and 51.8% for other boreal peatlands) in the Victor fen peat acts to reduce the LORCA relative to other estimates (Gorham 1991; Turunen et al. 2001; Vitt et al. 2000).

The basal salt marsh assemblage returned C accumulation values larger than the lowermost fen sequence. The greater accumulation may suggest that the peat matrix representing the salt marsh stage (abundant Typha remains) is perhaps more resistant to decay than the vegetation that formed the overlying fen peat (high proportions of

Sphagnum and sedge). Work in the boreal region of Alberta has indicated that litter from a freshwater marsh that included Typha latifolia and Carex spp. had a higher rate of decomposition (due to a higher litter quality) compared to the litter of poor (containing

Sphagnum teres and S. angustifolium, Carex spp. and Betula pumila) and moderately rich

(containing Tomenthypnum nitens, Carex spp. and B. pumila) fens (Thormann et al.

1999). This finding suggests that peat derived from the salt marsh vegetation assemblage would have decayed more readily than that of the fen stage. Thus, the rapid rate of accumulation during the salt marsh stage must be reflecting greater productivity, rather than reduced decay. By approximately 880 years BP the accumulation rate of the fen equals that of the salt marsh development stage and quickly surpasses it due to incomplete decay. For this period of the record, an alternative method of calculating carbon accumulation to LORCA would be preferable. One alternate approach could be calculation of the recent (apparent) rate of carbon accumulation (RERCA), based on a peat section between the surface and a given dated horizon in a surface core (Tolonen

92 and Turunen 1996). However, in this case the surface Jeglum core was not processed for any radiocarbon dates, as the top date at 37-38.5 cm is from the first drive of the Russian core, and the 210Pb dating did not return useable dates so RERCA is not a viable option.

Unless a radiocarbon date can be retrieved for the surface core, the options to account for the incomplete decay in the acrotelm seem limited.

LORCA appears to follow both temperature and precipitation, but not always in the same direction. High LORCA values tend to be associated with greater precipitation

(total annual and total June, July, and August) and lower LORCA values with lower precipitation (Figs 13 and 20-23). LORCA values often increase and decrease in phase with increasing and decreasing temperature (more so with July temperature) but this relationship is not always straightforward as there are periods of high LORCA with low temperatures (Figs 13 and 16-19). The resolution of the LORCA estimates (by depth) is twice that of the paleoclimatic reconstruction estimates derived from the fossil pollen assemblages, which may explain the periods of time during which the relationship between the climatic variables and the LORCA estimates are not so clear. Overall, the

LORCA and climatic reconstructions suggest that the surface moisture balance of the

Victor fen, controlled largely by annual and growing season precipitation, exerts a more powerful climatic control on carbon accumulation than annual or growing season temperature.

4.4 Resilience of the Victor Fen Ecosystem

Overall, the vegetation community remains relatively stable in a shrub fen state for approximately 6100 years. Even the hypothesized change in minerotrophic inputs to the fen does not cause it to permanently shift to an alternate stable state. Instead the

93 ecosystem retains the same vegetation structure according to the pollen assemblages despite some environmental changes, suggesting that this ecosystem benefits from some resilience to external perturbations, facilitating stability. In this case, stability would be the persistence of the fen ecosystem (Gunderson 2000). The resilience reflects the ability that a complex adaptive system has to self-organize versus the organization forced by external (allogenic) factors (Folke et al. 2004). The property of an ecosystem that defines the change in stability states and resilience is known as the adaptive capacity (Gunderson

2000). The alternate stability domains, or alternate stable states, of wetlands are characterized by the dominant plant species, dependent on key ecosystem processes and structures occurring at a variety of spatial and temporal scales (Gunderson 2000). The relatively stability of the fen ecosystem for multiple millennia supports the idea that the resilience has prohibited the change to an alternate stable. The true test of resilience is if the recent developments in the pollen spectra (rise in shrubs, disturbance indicators, sedges and a decline in Sphagnum spores) continue and remain firmly established. The resilience afforded to complex adaptive systems may be the reason that the fen system did not reach a tipping point and switch to an alternate stable state between 5800 and

4900 years BP, despite changing hydroclimatological conditions.

The idea of resilience is important because this site has been subject to two major external forcing factors. Isostatic uplift on the order of ≤210 metres from 7000 years BP to present would act to raise the piezometric surface (Glaser et al. 2004a) which would influence the expansion of peatlands through paludification as the elevation of the water table changed (Andrews and Peltier 1989). The uplift terrestrialized the site, initially facilitating the establishment of the salt marsh development stage. It continued to lift the

94 site above the level of the tides which acted to replace the marsh ecosystem with a fen.

The emergence of adjacent land from the tidal waters is reflected in the large increase in

Sphagnum spores in zone 2a, facilitated by widespread paludification. After this period of the record, isostatic uplift appears to have had less of an influence on the vegetation community of the fen. It may have facilitated further changes (ie. decreasing the gradient of rivers in the area and raising water levels facilitating further paludification (Glaser et al. 2004(a)), but the record is too coarse to isolate these factors.

Furthermore, the site may have experienced increases in temperature during the

Holocene Thermal Maximum (spanning 6000-3000 BP in the region, Kaufman et al.

2004; McAndrews et al. 1982), which for the arctic and sub-arctic region averaged a warming of 1.6 ± 0.8°C (summer estimates based on 16 sites in the western Subarctic and

Arctic) (Kaufman et al. 2004). Other models have suggested that at 6000 years BP, summer temperatures throughout the interior of continental North America were between

2 and 4 °C higher than present (COHMAP 1988). This increase in temperature could have influenced the moisture conditions at the surface of the peatland, implying that the vegetation of the fen can survive a range of surface moisture conditions. The site falls within the region that models predict is severely sensitive to climate change under a scenario of doubled pre-industrial atmospheric CO2 concentrations (Kettles and Tarnocai

1999). Given model predictions of an increase in average annual air temperatures of 3-4

°C by 2020, and 5-10 °C by 2050 (Tarnocai 2006), this resilience will be tested by future allogenic hydroclimatological forcing.

Chapter 5 CONCLUSION 4.5 Conclusions from the Victor Fen Record

The Victor fen contains a 6 500-year high-resolution record of peatland dynamics from the mid-Holocene to the present. Past paleoecological studies in the region have almost exclusively focused on bogs, so the complete fen record is useful to compare and contrast the two major northern peatland types. This region has been sparsely studied, so the new information surrounding vegetation history, carbon accumulation and responses to paleoclimatic changes will be very useful as more work is conducted in the HBL.

Post-glacial isostatic uplift isolated the fen site from the retreating Tyrrell Sea, and continued to exert an influence on the ecosystem, resulting in the succession from salt marsh to fen. Once the ecosystem became firmly established as a fen, it has never shifted to an alternate state, even though substantial changes in hydroclimatic conditions have taken place. However, the changes in these conditions are evident in the subtle oscillations in the pollen assemblages which returned variable conditions in the paleoclimate reconstructions.

The Victor fen record appears to have been influenced by major Holocene climatic transitions including the Holocene Climatic Optimum, the Neoglacial, the

Medieval Warm Period, the Little Ice Age and the recent 20th Century warming trend.

However, the changes are subtle and within the substantial error of the reconstructions. A more evident and sustained trend is the transition from dry Holocene Thermal Maximum conditions pre-3000 years BP, to more moist Neoglacial conditions from 3000 years BP to present. All of these climatic events seem to be captured in the pollen record at various

95 magnitudes, supporting the role that climate (both temperature and precipitation) has on influencing long-term northern peatland dynamics.

Varying hydroclimate conditions were captured by the peat humification data, which emphasizes the temperature and moisture signal of the fen depending on how strong an oscillation either forcing was undergoing. Hydroclimatological variability was also reflected in the carbon:nitrogen stratigraphy which captured a period of heightened acidity (represented by the peak in Sphagnum spores) suggesting some local change in hydrology (which may have promoted the autogenic rise in Sphagnum) and which also acts as an estimate for the degree of peat decay.

The ability that the fen has shown to remain essentially the same ecosystem while being exposed to internal and external forcings suggests that it is a resilient ecosystem as well as a complex adaptive system. However, with projected hydroclimatic trends caused by future global change, this resilience may be overpowered by allogenic forcing. The recent developments in the pollen stratigraphy suggest a change to a more sedge and shrub-dominated fen at the expense of Sphagnum consistent with trends from the Pan-

Arctic region (Tape et al. 2006). If this trend continues due to global change then the ecosystem may quickly reorganize and change to an alternate stable state.

The Victor fen underwent a period of rapid peat accumulation shortly after terrestrialization, followed by a slowdown and long period of low accumulation, followed more recently by another period of high accumulation. The carbon accumulation exhibits a similar trend, with a period of high accumulation from 6 500 to 6

000 years BP, followed by a period of low accumulation until 1 800 years BP, followed

96 97 lastly by a steadily increasing rate of carbon accumulation until the present day. The carbon accumulation rate is high compared to past published estimates of northern peatlands, but is consistent if the recent peat sequence is excluded, given the incomplete decay that has taken place in the uppermost depths and no alternative method to account for the incomplete decay. The basal salt marsh zone accumulated more carbon than the basal fen zone, which may reflect the high rate of primary productivity immediately following emergence from the sea rather than the differing resistances to decay between the dominant peat forming vegetation. The large scale climatic and hydrological changes also influenced the rate of peat deposition and carbon sequestration once isostatic uplift became less of an influence during the fen development stage. The rate of carbon accumulation was more closely related to the surface moisture balance (precipitation) suggested by reconstructions of precipitation, than it was to the average reconstructed temperature.

The response of the fen system to hydroclimatic change over the mid to late

Holocene will be useful for making predictions regarding how the vegetation and carbon accumulation of the ecosystem may adjust to future global change. Given the vast size of this ecosystem, constraining the response to forcing factors that may impact this substantial region is very important.

4.6 Future Work

Future work in the area could include the use of other proxies including plant macrofossils that may indicate different information regarding internal peatland dynamics than the proxies used in this study. The Sphagnum spores were only identified to genus, so the study of Sphagnum macrofossils which can often differentiate species as

98 well may yield more information. Stable isotope analysis could be conducted on

Sphagnum remains in a peat core from the area to act as another proxy climate method.

The use of Testate Amoebae (thecamoebians) to reconstruct local hydrological conditions could also be attempted. A core from an ombrotrophic bog near the location of the fen core has been analyzed in this manner (Bunbury and Finkelstein in prep.) and may provide greater insight regarding hydroclimatic change. A greater depth of the Tyrrell

Sea deposits could be studied to determine if the region did experience temperatures indicative of the Holocene Thermal Maximum while it was still inundated by the Sea, or if the Tyrrell Sea moderated changes in temperature. A study of carbon accumulation rates involving a greater number of peat cores may better constrain the LORCA estimates. The past responses to changes in precipitation and temperature could be used as a template for modelling the predicted future changes in these factors that may occur in the region. This would result in a better understanding of how the vegetation community may change, and how the rate of carbon sequestration may change. Lastly, if short-term climatic changes are of interest, then dendroclimatological studies of the Picea and Larix trees that are ubiquitous at and near the site may return local environmental signals that are annually resolved and thus at a much higher resolution than the perceived high resolution of this study. If a longer, annually resolved record is desired, then the search for a varved lake sediment sequence in the region may be of value.

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Appendix A: Raw counts of VC01

Level (cm) 0 4 7 13 20 Species Picea mariana 78 108 123 126 134 Picea glauca 9 8 4 6 0 Pinus banksiana 29 56 55 62 24 Betula 90 34 23 17 14 Carya 0 0 0 0 0 Quercus 0 1 0 0 0 Acer 0 0 2 2 3 Juglans 2 0 0 0 0 Tilia 0 0 0 0 0 Alnus 40 51 16 12 35 Larix 0 0 0 0 0 Salix 10 10 7 7 14 Artemesia 0 0 0 0 0 Ambrosia 16 9 8 3 27 Chenopods 2 5 3 5 13 Cyperaceae 131 117 74 89 16 Sphagnum 64 142 147 145 404 Typha latifolia 0 0 0 0 0 Ericaceae 1 3 0 0 1 Equisetum 0 0 1 0 0 Eupatorium 0 0 0 0 0 Helianthus 0 0 0 0 0 Unknown 0 0 0 0 0 Polypodium 0 1 0 0 0 Typha angustifolia 0 0 0 0 0 Potamogeton 0 0 0 0 0 Drosera 0 0 0 0 0 Lycopodium Spike 1832 1353 628 896 475

112 113

Level (cm) 31.5 38 43 51 55 Species Picea mariana 125 136 111 125 122 Picea glauca 4 3 1 3 4 Pinus banksiana 72 37 65 56 67 Betula 34 36 25 17 15 Carya 0 0 0 0 0 Quercus 4 0 0 0 1 Acer 1 2 1 1 0 Juglans 3 0 0 0 0 Tilia 0 0 0 0 0 Alnus 26 25 21 20 10 Larix 3 0 7 0 1 Salix 12 13 16 8 4 Artemesia 0 0 0 0 0 Ambrosia 7 6 5 4 1 Chenopods 2 1 1 4 0 Cyperaceae 69 115 38 70 24 Sphagnum 489 496 265 211 219 Typha latifolia 1 0 0 0 1 Ericaceae 2 3 1 1 0 Equisetum 14 0 1 0 0 Eupatorium 3 0 1 0 0 Helianthus 0 0 0 0 0 Unknown 0 0 0 0 0 Polypodium 0 0 0 0 0 Typha angustifolia 0 0 0 0 0 Potamogeton 0 0 0 0 0 Drosera 0 0 0 0 0 Lycopodium Spike 890 1088 1229 851 172

114

Level (cm) 64 70.5 77 83 90 Species Picea mariana 107 140 142 129 119 Picea glauca 0 0 0 1 4 Pinus banksiana 66 52 44 46 53 Betula 34 52 20 29 21 Carya 0 0 0 0 0 Quercus 0 0 0 0 0 Acer 0 2 2 2 1 Juglans 0 1 1 1 0 Tilia 0 0 0 0 0 Alnus 20 27 15 17 9 Larix 0 5 0 2 4 Salix 10 6 13 11 4 Artemesia 0 0 0 0 0 Ambrosia 3 6 2 3 0 Chenopods 2 6 0 1 2 Cyperaceae 41 40 73 37 53 Sphagnum 311 282 258 189 178 Typha latifolia 0 1 0 0 0 Ericaceae 2 2 0 2 1 Equisetum 0 6 0 1 2 Eupatorium 0 0 0 0 0 Helianthus 0 0 0 0 0 Unknown 3 0 0 0 0 Polypodium 1 0 0 1 0 Typha angustifolia 0 0 0 0 0 Potamogeton 0 0 0 0 0 Drosera 1 0 0 0 1 Lycopodium Spike 1112 1162 1080 515 408

115

Level (cm) 97 103 109.5 116 121.5 Species Picea mariana 161 142 137 158 160 Picea glauca 2 0 2 0 6 Pinus banksiana 52 40 74 16 51 Betula 16 54 59 58 35 Carya 0 0 0 0 0 Quercus 0 0 3 0 0 Acer 1 0 6 0 0 Juglans 0 0 1 0 1 Tilia 0 0 0 0 0 Alnus 9 11 18 18 18 Larix 0 10 3 7 0 Salix 1 0 9 1 10 Artemesia 0 0 0 0 0 Ambrosia 0 8 0 3 2 Chenopods 0 0 1 0 2 Cyperaceae 26 57 58 26 149 Sphagnum 140 296 346 262 414 Typha latifolia 0 0 1 0 0 Ericaceae 0 7 5 2 1 Equisetum 0 0 2 0 1 Eupatorium 0 0 0 0 0 Helianthus 0 0 0 0 0 Unknown 0 0 0 5 0 Polypodium 0 0 0 0 0 Typha angustifolia 0 0 0 0 0 Potamogeton 0 0 0 0 0 Drosera 0 0 0 0 0 Lycopodium Spike 134 362 313 391 316

116

Level (cm) 129 134 142 148.5 155 Species Picea mariana 204 125 120 141 134 Picea glauca 0 2 12 5 0 Pinus banksiana 23 61 51 26 52 Betula 12 26 26 39 51 Carya 0 0 0 1 0 Quercus 0 0 0 0 0 Acer 0 1 1 1 0 Juglans 0 0 0 0 0 Tilia 0 0 0 0 0 Alnus 11 15 15 12 34 Larix 4 4 0 4 3 Salix 0 8 8 5 11 Artemesia 0 0 0 0 0 Ambrosia 0 7 0 3 4 Chenopods 0 2 2 1 4 Cyperaceae 28 45 48 22 65 Sphagnum 217 192 241 266 478 Typha latifolia 0 0 0 0 0 Ericaceae 1 1 1 8 6 Equisetum 0 0 0 0 0 Eupatorium 0 1 0 0 0 Helianthus 0 0 0 0 0 Unknown 0 0 0 0 10 Polypodium 0 0 0 0 0 Typha angustifolia 0 0 0 0 0 Potamogeton 0 0 0 0 0 Drosera 0 0 0 0 0 Lycopodium Spike 733 761 272 517 201

117

Level (cm) 161.5 168 175 181 187.5 Species Picea mariana 199 183 132 162 106 Picea glauca 2 0 2 0 4 Pinus banksiana 54 25 55 38 47 Betula 108 40 56 76 49 Carya 0 0 0 0 0 Quercus 0 1 3 1 3 Acer 9 2 0 1 2 Juglans 0 0 0 0 0 Tilia 0 0 0 0 0 Alnus 31 12 20 36 22 Larix 7 3 2 0 3 Salix 12 9 4 8 6 Artemesia 0 3 0 0 0 Ambrosia 3 5 5 2 3 Chenopods 1 1 0 0 0 Cyperaceae 73 43 37 45 57 Sphagnum 730 345 206 157 242 Typha latifolia 0 4 0 0 0 Ericaceae 14 2 3 4 5 Equisetum 2 0 1 1 1 Eupatorium 0 0 1 0 0 Helianthus 0 0 0 0 0 Unknown 1 0 0 1 0 Polypodium 0 0 0 0 0 Typha angustifolia 0 0 0 0 0 Potamogeton 0 0 0 0 0 Drosera 0 0 0 0 0 Lycopodium Spike 1112 274 395 549 324

118

Level (cm) 194 200.5 205 214 220 Species Picea mariana 183 148 166 165 145 Picea glauca 0 5 0 5 0 Pinus banksiana 22 31 43 49 25 Betula 36 25 36 35 51 Carya 0 0 0 0 0 Quercus 0 1 0 0 0 Acer 4 3 0 0 0 Juglans 1 0 0 0 0 Tilia 0 0 0 0 0 Alnus 6 6 25 21 8 Larix 0 3 0 2 0 Salix 3 2 16 10 4 Artemesia 0 0 0 0 0 Ambrosia 4 1 5 4 0 Chenopods 2 0 9 3 2 Cyperaceae 12 19 22 27 48 Sphagnum 1205 455 784 530 281 Typha latifolia 0 0 0 0 1 Ericaceae 2 0 8 1 5 Equisetum 0 7 0 5 0 Eupatorium 0 0 0 0 0 Helianthus 0 0 0 0 0 Unknown 0 0 0 1 0 Polypodium 0 0 0 0 0 Typha angustifolia 0 0 0 0 0 Potamogeton 0 0 0 0 0 Drosera 0 0 0 0 0 Lycopodium Spike 162 257 352 580 467

119

Level (cm) 224.5 233 240 244 251 Species Picea mariana 49 146 192 132 87 Picea glauca 0 0 0 4 2 Pinus banksiana 41 21 59 57 78 Betula 127 90 27 19 40 Carya 0 0 0 2 0 Quercus 0 0 0 0 0 Acer 0 0 0 6 0 Juglans 0 0 0 0 0 Tilia 0 0 0 0 0 Alnus 11 16 12 6 33 Larix 1 1 0 4 2 Salix 1 13 12 14 18 Artemesia 0 0 0 1 0 Ambrosia 3 3 5 4 5 Chenopods 1 0 3 2 15 Cyperaceae 11 34 376 111 150 Sphagnum 113 435 65 12 95 Typha latifolia 0 5 128 1662 8 Ericaceae 1 18 5 1 2 Equisetum 0 0 0 0 4 Eupatorium 0 0 0 7 3 Helianthus 0 0 0 0 0 Unknown 0 0 0 0 0 Polypodium 0 0 0 0 0 Typha angustifolia 0 0 0 0 4 Potamogeton 0 0 0 0 1 Drosera 0 0 0 0 0 Lycopodium Spike 293 466 178 510 489

120

Level (cm) 256 Species Picea mariana 109 Picea glauca 5 Pinus banksiana 70 Betula 29 Carya 0 Quercus 0 Acer 0 Juglans 0 Tilia 1 Alnus 8 Larix 1 Salix 12 Artemesia 0 Ambrosia 9 Chenopods 5 Cyperaceae 44 Sphagnum 59 Typha latifolia 11 Ericaceae 0 Equisetum 1 Eupatorium 0 Helianthus 0 Unknown 0 Polypodium 0 Typha angustifolia 3 Potamogeton 0 Drosera 0 Lycopodium Spike 507