Paleoecological and Carbon Accumulation Dynamics of a Fen Peatland in the Hudson Bay Lowlands, Northern

Ontario, from the Mid-Holocene to Present

by

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

© Copyright by Benjamin Cody O’Reilly 2011

ii

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.

iii

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

iv

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!

v

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

vi

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

vii

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

viii

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

ix

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

x

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

xi

List of Appendices Appendix A: VC01 raw pollen counts………………………………………………………….112

1

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

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.

7

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.

8

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.

9

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).

10

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

11

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.

Proxy Derived From Proxy Signal Controlled By Interpreted

Reconstruction

Pollen

Organic- walled

microfossils isolated from the

peat matrix

(1) Local-regional vegetation at/near

study site (2) Regional pollen

rain

(1) Hydroclimatology (2) Isostatic Uplift

(3) Air Masses

Vegetation Reconstruction/

Succession

Spectrophotometric Humification

Humic acids chemically extracted from the

dried peat matrix

(1) Moisture content

(2) Aerobic decomposition in

the acrotelm (3) Water table

depth

(1) Peat forming vegetation

(2) Differential resistance to

decomposition (3) Compaction of peat

Degree of peat decomposition and therefore, depth to water

table

C:N Ratios

C:N bulk content of the dried

peat matrix

(1) Peat forming vegetation

(2) Decomposition in catotelm

(3) Peat accumulation

(1) Hydrological inputs (ombrotrophy vs minerotrophy)

(2) Residence time of peat in acrotelm

Carbon accumulation

estimates; Isolation of successional

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

13

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.

15

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.

17

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.

18

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).

26

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.

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

Year (ka BP)

0 2000 4000 6000 8000

Rel

ativ

e se

a le

vel c

hang

es (m

)

0

100

200

300

400

32

(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).

34

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.

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.

36

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

40

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.

Sample

Depth

(cm)

No. of

species for

first 100

tree pollen

No. of

species for

second 100

tree pollen

New

Species in

2nd 100

grains

Proportion of new

species at level

Maximum proportion

of new species

found at any

analyzed depth

New

Species at

250 arboreal

grains

97 9 8 0 0 0 0

121.5 15 9 0 0 0 0

161.5 10 12 3

Ericaceae = 0.3% ;

Ambrosia = 0.3% ;

Larix = 1.1%

Ericaceae = 1.5% ;

Ambrosia = 3.9% ;

Larix = 1.6%

0

181 10 12 2 Acer = 0.18% ;

Ambrosia = 0.37%

Acer = 0.87% ;

Ambrosia = 3.9% 0

41

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.

44

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.

45

210Pb Total Activity (DPM/g)

10 100

Dep

th (c

m)

0

2

4

6

8

10

12

14

16

18

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.

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.

Depth (cm)0 50 100 150 200 250

Age

(cal

Yea

r BP)

0

1000

2000

3000

4000

5000

6000

48

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

Sample

Number

Laboratory

Number

Sample Depth

(cm)

Dated

Material

Conventional

Age 14C (years

BP)

2σ Calibration

(years BP)

Median Age

(years BP)

13C/12C Ratio

(‰)

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

49

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).

52

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.

53

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

54

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.

Depth (cm)

0 50 100 150 200 250

Dry

Bul

k D

ensi

ty (g

cm

-3)

0.10

0.12

0.14

0.16

0.18

0.20

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).

56

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

57

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.

% Carbon

36 38 40 42 44 46 48 50

Dep

th (c

m)

0

50

100

150

200

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).

% Nitrogen

1.0 1.5 2.0 2.5 3.0 3.5

Dep

th (c

m)

0

50

100

150

200

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.

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

C:N Ratio

10 15 20 25 30 35 40

Dep

th (c

m)

0

50

100

150

200

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.

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

LORCA (g C m-2 year-1)

0 50 100 150 200

Dep

th (c

m)

0

50

100

150

200

61

LORCA (g C m-2 year-1)

10 20 30 40 50 60

100

150

200

Dep

th (c

m)

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

LORCA (g C m-2 year-1)

0 50 100 150 200

Age

(cal

Yea

r BP)

0

1000

2000

3000

4000

5000

6000

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

62

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).

64

Age (cal Year BP)0 2000 4000 6000

Spec

troph

otom

etric

Abs

orba

nce

(%)

0

20

40

60

80

100

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.

Age (Cal Year BP)

0 2000 4000 6000

Det

rend

ed A

bsor

banc

e R

esid

ual (

A d)

-20

-10

0

10

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

65

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.

68

Age (cal Year BP)

0 2000 4000 6000

Ave

rage

Ann

ual T

empe

ratu

re (°

C)

-10

-5

0

5

10

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.

69

Age (cal Year BP)

0 500 1000 1500 2000

Aver

age

Ann

ual T

empe

ratu

re (°

C)

-5

0

5

10

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.

70

Age (cal Year BP)

0 2000 4000 6000

Aver

age

July

Tem

pera

ture

(°C

)

10

15

20

25

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.

71

Age (cal Year BP)

0 500 1000 1500 2000

Aver

age

July

Tem

pera

ture

(°C

)

8

10

12

14

16

18

20

22

24

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.

72

Age (cal Year BP)

0 2000 4000 6000

Tota

l Ann

ual P

reci

pita

tion

(mm

)

200

400

600

800

1000

1200

1400

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.

73

Age (cal Year BP)

0 500 1000 1500 2000

Tota

l Ann

ual P

reci

pita

tion

(mm

)

0

200

400

600

800

1000

1200

1400

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.

74

Age (cal Year BP)

0 2000 4000 6000

Ave

rage

JJA

Pre

cipi

tatio

n (m

m)

20

40

60

80

100

120

140

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.

75

Age (cal Year BP)

0 500 1000 1500 2000

Ave

rage

JJA

Pre

cipi

tatio

n (m

m)

20

40

60

80

100

120

140

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.

76

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.

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).

78

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.

79

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

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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

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(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

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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

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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

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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

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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

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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).

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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

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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

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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

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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

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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.

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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

96

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

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

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