Postglacial Water Levels in the Great Lakes Region in Relation to Holocene Climate Change: Thecamoebian and Palynological Evidence

Home , Georgian Bay, Lake Stanley, Lake Tonawanda, Main Channel (Ontario), Wanapitei River

Adam Patrick Sarvis, B.Sc.

A thesis

submitted to the Department of Earth Sciences

in partial fulfillment of the requirements

for the degree of

Master of Science

December, 2000

Brock University

St. Catharines, Ontario

Abstract

Various lake phases have developed in the upper Great Lakes in response to isostatic adjustment and changes in water supply since the retreat of the Laurentide Ice Sheet. Georgian

Bay experienced a lowstand that caused a basin wide unconformity approximately 7,500 years ago that cannot be explained by geological events. Thecamoebians are shelled protozoans abundant in freshwater environments and they are generally more sensitive to changing environmental conditions than the surrounding vegetation. Thecamoebians can be used to reconstruct the paleolimnology. The abundance of thecamoebians belonging to the genus

Centropyxis, which are known to tolerate slightly brackish conditions (i.e. high concentrations of ions) records highly evaporative conditions in a closed basin. During the warmer interval (9000 to 700 yBP), the Centropyxis - dominated population diminishes and is replaced by an abundant and diverse Difflugia dominate population.

Historical climate records from Tobermory and Midland, Ontario were correlated with the Lake Huron water level curve. The fossil pollen record and comparison with modem analogues allowed a paleo-water budget to be calculated for Georgian Bay. Transfer function analysis of fossil pollen data from Georgian Bay records cold, dry winters similar to modem day

Minneapolis, Minnesota. Drier climates around this time are also recorded in bog environments in Southem Ontario - the drying of Lake Tonawanda and inception of paludification in

Willoughby Bog, for instance, dates around 7,000 years ago. The dramatic impact of climate change on the water level in Georgian Bay underlines the importance of paleoclimatic research for predicting future environmental change in the Great Lakes...... - ;,_.:;' 'if-Utiy ,-. .

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The last three years completing this thesis has strengthened my scientific beliefs and inner character. I have many people to thank for their contribution in allowing me to finish this thesis. First, I would personally like to thank Dr. Francine McCarthy in the Earth Sciences department for providing me the opportunity to study this unique topic on climate change and lake levels. Over the years she has shared her wisdom and experiences to broaden my field of

study and understanding a graduate student lifestyle. I would like to express my gratitude to

Steve Blasco of the Geological Survey of Canada (Atlantic) in providing support through the years for paleoecological research and understanding the importance of Earth Sciences at Brock

University. I'm grateful for being part of an extensive research program on Georgian Bay that I hope will continue through the years. Extended gratitude and special thanks to Dr. Jock

McAndrews of the Royal Ontario Museum for providing computer software and personal data to this project. The field trips to Midland and Willoughby Bog were very helpful and educational.

John Menzies in the Earth Sciences department for advising me through the years. Anne Yagi of the Ministry of Natural Resources for being a secondary editor and for being there when times got tough (Thanks big sister). Special thanks to Jim Pengelly for his fime and personal

resources, I enjoyed our talks about the geomorphology of the Wainfleet Bog. I hope you become passionate with your work again very soon. Kevin Gostlin, fellow graduate student in the Earth Sciences department for his advice and good times through the years. Howard Melville in the Jaan Terasmae Radiocarbon Lab in providing the dates for this study and for the lab preparation in the first year course. Mike Lozon the Department draftsman for his expertise and helpfulness in producing the diagrams for this thesis. Diane Gadoury, the Department Secretary

for keeping the photocopy room open and all the little things that don't get noticed. Dr. Rick

Cheel in the Earth Sciences department for providing a light-hearted atmosphere throughout the years. The librarians Colleen Beard and Sharon Barnes of the Brock University Map Library for their help with the maps and airphotos, and not charging me overdue fees. Roland Seehagel and !

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*^- V ( I. > / ^fivC/' -^.Mi .^ H-Ji IV his staff in the Machine shop at Brock University for constructing the soil probe. Dr. Keith

Tinkler and Loris Gasparotto in the Geography department for their time in providing the diagram of postglacial water levels in the Niagara Region. Jim Cook of Environment Canada in

1998 for being cooperative in providing climate information for this research.

The Niagara Peninsula Conservation Authority (N.P.C.A.) for allowing me access to their property at Willoughby Bog and Wainfleet Bog. The volunteer organization Friends of Fort

Erie's Creeks (F.F.E.C.) for allowing me to pursue field work in the Wainfleet Bog during the summer of 1998. Thanks to Elmer Miskolczi, Matthew Pilato and the crew of 1998 for their passion in environmental restoration. A special recognition to the memory of Kevin Hartman and Craig Doan, two good friends who were taken from us in 1997 and loved the F.F.E.C.

To my parents Fred and Helen, and the rest of my family and close friends I would especially like to thank you for putting up with my crazy moods the last three years. To my brother Fred for providing me a loan to purchase a new computer and Jerry Peters for his computer expertise. Thanks for your help. Steve Steingart and Carol Dohn for coming out with

me to Willoughby Bog in 1998 at the last possible minute. Thanks. 1 997 was a dramatic year for

my family and I with the passing of my grandmother Elsie. This episode of my life tested my inner strength to continue my academic career. She was beloved and her memories will be close to our hearts.

To graduate and undergraduate students who have made this experience worthwhile and enjoyable in the Earth Sciences department throughout the years (Jaime Oxtobee, Peter Kryger,

Jennifer Hopkins, Michael Johnson, Sandra Savage, Lee Ann Warrell, Michael Shuker, Christine

March, Simon Haynes and others). Thanks. An overall thanks to the faculty and staff of the Earth

Sciences department for their generosity through the years in finishing this thesis.

Table of Contents

Abstract ii

Acknowledgements iii

Table of Contents v

List of Figures ...viii

List of Tables vi

1 .0 Introduction 11

1.1 Previous Studies on Georgian Bay 14

1 .2 Previous Studies on Willoughby Bog 17

2.0 Study Areas 17

2. Georgian Bay 17

2.2 Willoughby Bog 22

3.0 Georgian Bay Background Information 22

3.1 Climate of the Great Lakes 22

3.1.1 Climate of Georgian Bay 26

3.2 Vegetation and Soils of Georgian Bay 28

3.3 Hydrology of Georgian Bay 29

3.4 Geology of Georgian Bay 32

3.5 Geomorphology and Quaternary Geology of Georgian Bay 32

4.0 Willoughby Bog Background Information 35

4.1 Climate of Willoughby Bog 35

4.2 Vegetation and Soils of Willoughby Bog 36

4.3 Hydrology of Willoughby Bog 37

4.4 Geology and Geomorphology of Willoughby Bog 37

5.0 Postglacial History of the Late Wisconsinan Period in the Great Lakes Basin 38

5.1 Glacio-Isostasy in the Great Lakes 46

6.0 Palynology 47

6. Palynology and Wetland Environments 49

6.2 Palynology and Lake Environments 50

6.3 Transfer functions 50

6.4 Holocene Climate of Southern Ontario 52 .!.,'' 1 1

7.0 Thecamoebians 55

7. 1 Problems of Thecamoebians as Paleoecological Indicators 58

8.0 Methodology 60

8.1 Field Methods 60

8.1.1 Georgian Bay 60

8.1.2 Willoughby Bog 63

8.2 Laboratory Methods 68

8.2. Stratigraphy and Thermal Analysis (LOI) 68

8.2.1.1 Georgian Bay 68

8.2.1.2 Willoughby Bog 68

8.2.2 Pollen Analysis and transfer functions 69

8.2.2.1 Georgian Bay • 69

8.2.2.2 Willoughby Bog 69

8.2.3 Thecamoebitm Analysis 73

8.2.3.1 Georgian Bay 73

8.2.3.2 Willoughby Bog 73

8.3 Water Budget Modeling 74

9.0 Results 76

9. 1 Georgian Bay 76

9.1.1 LH95-035(Manitoulin Island) 76

9.1.1.1 Transfer functions 76

9. 1 . 1 .2 Thecamoebians 78

9.1.2 LH95-038(Manitoulin Island) 78

9. 1 .2. Transfer functions 78

9. 1 .2.2 Thecamoebians 80

9.1.3 LH95-041 (Tobermory) 80

9. 1 .3. Transfer functions 80

9. 1 .3.2 Thecamoebians 82

9.1.4 97802-020 (Severn Sound) 82

9. 1 .4. Transfer functions 82

9.1.4.2 Thecamoebians 82

9. 1 .4.3 Loss-on-ignition 84 ' v 1

Vll

9.1.5 97802-026 (Severn Sound) 84

9. 1 .5. Transfer functions 84

9. 1 .5.2 Thecamoebians 86

9. 1 .5.3 Loss-on-ignition 86

9.1.6 97802-032 (Severn Sound) 87

9. 1 .6. Transfer functions 87

9. 1 .6.2 Thecamoebians 87

9. 1 .6.3 Loss-on-ignition 87

9.1.7 97802-035 (Severn Sound) 87

9. 1 .7. Transfer fiinctions 87

9. 1 .7.2 Thecamoebians 91

9. 1 .7.3 Loss-on-ignition 91

9.1.8 97802-036 (Severn Sound) 92

9.1.8.1 Transfer functions 92

9. 1 .8.2 Thecamoebians 92

9. 1 .8.3 Loss-on-ignition 92

9.1.9 Transfer functions from Wye Marsh, Porqui Pond and Shouldice Lake 92

9.2 Modem Water Budget for Georgian Bay 94

9.2.1 Historical Climate Records for Georgian Bay 95

9.2.2 Problems with the Georgian Bay Water Budget 100

9.3 Willoughby Bog 101

9.3.1 WB4 (Channel Environment) 101

9.3.1.1 Palynology 101

9.3. 1 .2 Thecamoebians 1 03

9.3. 1 .3 Loss-on-ignition 103

9.3.2 WB6 (Channel Environment) 105

9.3.2.1 Palynology 105

9.3 .2.2 Thecamoebians 1 07

9.3 .2.3 Loss-on-ignition 1 07

10.0 Discussion 108

11.0 Conclusions 119

References Cited 123

Vlll Appendices

Appendix A: Water Budget, Climate, Vegetation, Lake Phases and Modem Analogue of

Georgian Bay 135

Appendix B: Thecamoebian & Loss-on-ignition data from Georgian Bay 140

Appendix C: Climate, Water Budgets, Lake level & Runoff data from Georgian Bay 147

Appendix D : Microfossil, Loss-on-ignition & Field data from Willoughby Bog 1 63

List of Figures

Figure 1 : Great Lakes Region and study areas 12

Figure 2: Georgian Bay stratigraphy and characteristic thecamoebian assemblages 16

Figure 3 : Erie waters forming Lakes Tonawanda and Wainfleet 18

Figure 4: Georgian Bay basin and bathymetry 19

Figure 5: Georgian Bay drainage basin 21

Figure 6: Niagara Peninsula and local wetlands (close up of Willoughby Bog) 23

Figure 7: Topographic map of the Willoughby Bog and surrounding area 24

Figure 8: Climate maps of eastern North America and the Great Lakes basin 25

Figure 9: Water budget of Tobermory, Ontario 28

Figure 10: Quaternary geology of the Georgian Bay basin 33

Figure 1 1 : Bottom sediments of Georgian Bay 34

Figure 12: Water budget for Welland, Ontario 36

Figure 13: Postglacial lake levels of the Great Lakes and major drainage outlets 40

Figure 14: Late Hough (Georgian Bay) and Late Stanley (Lake Huron) lake phases 43

Figure 15: Seismic stratigraphy of core LH95-041 43

Figure 16: Nipissing transgression around 7,000 yBP 45

Figure 17: Nipissing Flood around 5,500 yBP 45

Figure 1 8: Paleoclimate curve for southern Ontario 53

Figure 19: Paleoclimate maps of eastern North America 54

Figure 20: Common thecamoebian assemblages in Georgian Bay and Willoughby Bog 56

Figure 21: Core locations from Tobermory and Manitoulin Island 61

Figure 22: Tobermory airphoto showing core LH95-041 62

Figure 23 : Core locations from Severn Sound 64

Figure 24: Severn Sound airphoto showing core locations 65

Figure 25: Airphoto of Willoughby Bog 66

Figure 26: Soil transects and core locations at Willoughby Bog 67

Figure 27: Soil transect S-S' of the organic basin and core locations of Willoughby Bog 67

Figure 28: Wye Marsh pollen diagram with transfer functions 70

Figure 29: Porqui Pond pollen diagram with transfer functions 71

Figure 30: Shouldice Lake pollen diagram with transfer functions 72

Figure 31: Pollen and thecamoebian diagrams with transfer ftmctions of core LH95-035 77

Figure 32: Pollen and thecamoebian diagrams with transfer functions of core LH95-038 79

Figure 33: Pollen and thecamoebian diagrams with transfer functions of core LH95-041 81

Figure 34: Pollen and thecamoebian diagrams with transfer ftinctions of core 97802-020 83

Figure 35: Pollen and thecamoebian diagrams with transfer fiinctions of core 97802-026 85

Figure 36: Pollen and thecamoebian diagrams with transfer fiinctions of core 97802-032 88

Figure 37: Loss-on-ignition for the Severn Sound cores 89

Figure 38: Pollen and thecamoebian diagrams with transfer functions of core 97802-035 90

Figure 39: Pollen and thecamoebian diagrams with transfer functions of core 97802-036 93

Figure 40: Modem water budget for Georgian Bay 95

Figure 41 : Tobermory climate versus Huron-Michigan lake levels 96

Figure 42: Midland climate versus Huron-Michigan lake levels 96

Figure 43 : Modem climate compared to Georgian Bay lowstand and highstand 98

Figure 44: Modem winter precipitation compared to Georgian Bay lake levels 99

Figure 45 : Pollen and thecamoebian diagrams with transfer functions of core WB4 1 02

Figure 46: Stratigraphy and loss-on-ignition for cores WB4 and WB6 1 04

Figure 47: Pollen and thecamoebian diagrams with transfer fiinctions of core WB6 106

Figure 48: Pollen and thecamoebian zones compared to postglacial lake levels 109

Figure 49: Paleo-water budget reconstmction of Georgian Bay around 7,500 yBP Ill

Figure 50: Modem water budget for Minnesota, MN 116

List of Tables

Table 1 : Climate parameters of Georgian Bay 27

Table 2: Morphological characteristics for lakes Michigan, Huron, Erie and Georgian Bay 30

Table 3: Georgian Bay drainage basins and surface hydrology 31

Table 4: Pollen zonation in Georgian Bay in relation to vegetation change 48 .'V' , •-. .:(

iL X

Table 5: Georgian Bay core locations 60

Table 6: Runoff gauging stations for Georgian Bay drainage basin 75

Table 7: Mean monthly runoff from the Georgian Bay drainage basin 76

Table 8: Paleoclimate averages for pollen zones of cores LH95-035, 038 & 041 78

Table 9: Sediment volume and total thecamoebian counts for cores LH95-035, 038 «& 041 80

Table 10: Paleoclimate averages for pollen zones of cores 97802-020, 026 & 032 84

Table 11: Sediment volume and total thecamoebian counts for cores 97802-020, 026 & 032.... 87

Table 12: Paleoclimate averages for pollen zones of cores 97802-035 & 036 91

Table 13: Sediment volume and total thecamoebian counts for cores 97802-035 & 036 92

Table 14: Paleoclimate averages for Severn Sound cores (97802-020,97802-026, 97802-032,

97802-035, 97802-036) compared to Porqui Pond and Wye Marsh 94

Table 15: Paleoclimate averages for Tobermory (LH95-041) and Manitoulin Island (LH95-035,

LH95-038) cores compared to Shouldice Lake 94

Table 16: Average paleoclimate values for different pollen zones in cores WB4 and WB6 105

Table 17: Sediment volume and total thecamoebian counts for cores WB4 and WB6 107

11 1.0 Introduction

The postglacial water levels of the Great Lakes have been extensively studied through the last century. In the early 1990's to today, various studies by the Geological Survey of Canada

(GSC)- Atlantic were conducted to investigate the evolution of the Georgian Bay basin through

the Holocene Period (Figure 1). Georgian Bay has experienced various lake phases since the retreat of the Laurentide Ice Sheet as a result of isostatic rebound which led to alternating drainage outlets through Fenelon Falls, North Bay, and Port Huron. Earlier work by Blasco

(1994) on Georgian Bay discovered a lowstand through seismic profiles. This lowstand is referred to as the Late Hough phase. The Late Hough lowstand in Georgian Bay was lower than

Lake Huron (Rea et al., 1994). Palynological data from McCarthy et al. (1997) had dated the

Late Hough lowstand around 7,500 yBP. The 7,500 yBP lowstand in Georgian Bay experienced a lowstand apparently unrelated to geological processes because the water level was lower than

North Bay and the local sills, and there is no known outlet discovered at present (Blasco, 1994;

Blasco et al., 1997; Blasco et al., 1999). Around 7,500 yBP, Georgian Bay would have been a closed basin for a few hundred years, when water levels were controlled by evaporation and precipitation until the Nipissing Rise. This study was to investigate a climate explanation for the lowstand using paleo-indicators.

The main objectives for this thesis were to use palynological data from McCarthy et al.

(1997) for reconstructing the paleoclimate of Georgian Bay. Pollen data enabled the sediments to be dated by comparison with established regional pollen stratigraphies, and transfer functions established quantitative reconstructions of regional paleoclimates to be derived from pollen data

(Dubas, 1997; McCarthy et al., 1997; Tiffin, 1998). Pollen and spores are a common paleo- indicator used to reconstruct paleoenvironments and are found in various depositional

environments (i.e. lakes and bogs). Pollen analyses provide past vegetation records, age

Figure 1: Great Lakes Region and study areas.

13 estimation of sediments and past climate reconstructions (Birks & Birks, 1 980;

McAndrews, 1981). Thecamoebians were used in this study because they are excellent paleolimnological indicators to reconstruct lake levels. Thecamoebian samples were taken according to pollen zones from the GSC cores. Thecamoebian species vary with water quality parameters (water temperature, nutrient levels, and salinity) then their presence or absence in the

sediment record can indicate a change in environmental conditions (McCarthy et al., 1997;

Tiffin, 1998; McLachlan, 1999).

Modem historical climate data from Georgian Bay was derived from Environment

Canada (Cook personal comm., 1998). The climate data was required to correlate with modem

Lake Huron water levels and to establish a modem water budget for Georgian Bay. The basis of the modem water budget was to compare and reconstmct a paleo-water budget for Georgian Bay

7,500 years ago. The present day relationship between water levels in Georgian Bay and climate

is obvious from instmmental records in Tobermory and Midland, therefore variations in paleotemperature and paleoprecipitation can be expected to have influenced prehistoric levels.

The effects of climate on prehistoric lake levels in Georgian Bay have received little study,

although climate is known to have had a noticeable effect on the level of Lake Ontario

(McCarthy and McAndrews, 1988) and of lakes Michigan and Huron (Fraser et al. 1990).

Willoughby Bog in the Niagara Peninsula is a wetland outside the Georgian Bay basin.

Wetlands preserve their evolution in the organic sediments, where decay processes are slow compared to the surface (Wamer & Bunting, 1996). There are various types of wetlands (i.e. bogs, marshes and swamps) that are identified by land that is saturated with water long enough to promote aquatic processes indicated by poorly drained soils, hydrophytic vegetation and

biological activity which are adapted to wet environments (Tamocai et al., 1988). Willoughby

Bog was further investigated because bog formation occurred around 7,000 yBP, similar time period of the lowstand in Georgian Bay (Pengelly, 1990; Sarvis, 1997). Continued

14 paleoecological research on thecamoebians in additional cores taken from Willoughby Bog provides insights into the hydrological evolution of the modem peat margin swamp from the floodplain of the fluvial system connecting to glacial Lake Tonawanda. Lake Tonawanda was a shallow lake that was formed in the Erie basin from the meltwaters of Lake Agassiz. The paleoecological evidence from Georgian Bay and Willoughby Bog may reconstruct the paleoclimate that affected southern Ontario and the water levels of the Great Lakes.

,;• 1.1 Previous Studies on Georgian Bay v,v h r

There have been various studies of Georgian Bay ranging from microfossil analyses (i.e. pollen and thecamoebians) (Jesseau, 1972; McAtee, 1977; Terasmae, 1979; Hoey, 1981;

McAndrews, 1981; Anderson & Lewis, 1992; Fraser et al., 1990; McAndrews, 1994; Dubas,

1997; McCarthy et al., 1997; Burbidge & Schroder-Adams, 1998; Reinhardt et al., 1998;

McLachlan, 1999; Tiffin, 1998; McCarthy & Tiffin, 1999; Sarvis et al., 1999a; Sarvis et al.,

1999b), paleolimnology (Hough, 1962; Anderson & Terasmae, 1966; Larsen, 1985; Larsen,

1987; Bishop, 1990; Lewis et al., 1994; Blasco, 1994; Moore et al., 1994; Rea et al., 1994;

Dobson et al., 1995; Blasco et al., 1997; Blasco et al., 1999), climate (Phillips & McCuIloch,

1972; Steinhauser, 1979; Brown et al., 1980; Edwards & Fritz, 1988) and Quaternary geology

(Hoffman & Richards, 1954; Hoffman et al., 1962; Karrow et al., 1975; Chen, 1981; Chapman &

Putnam, 1984; Karrow, 1986; Karrow, 1989; Bamett et al., 1991; Bamett, 1992).

Some pollen analyses have been performed on sites in and around Georgian Bay (Karrow

et al., 1975; McAtee, 1977; Anderson «& Lewis, 1989; McAndrews, 1994; Dubas, 1997;

McCarthy et al., 1997). McAtee (1977) analyzed fluvial transportation, distribution, and depositionof arboreal pollen in the Georgian Bay drainage basin. He interpreted the paleoclimate around 7,500 yBP as moist and cool during rising lake levels in the basin. Karrow

et al. (1975) used multiple paleo-indicators including arboreal pollen in interpreting the

15 paleoenvironment during the Lake Algonquin lake (1 1,200 to 1 1,000 yBP). Rising lake

levels recycled arboreal pollen from Lake Algonquin sediments (McCarthy et al., 1997). Dubas

(1997) interpreted several cores from Tobermory using palynological evidence and seismic stratigraphy to interpret the Nipissing transgression (approx. 7,000 yBP). As the lake level declined after the Nipissing Rise (5,500 yBP), sediment availability decreased resulting in a diastem in Georgian Bay (Dubas, 1997). Georgian Bay lowstands are represented by erosional

unconformities dated palynologically as oxidized layers (McCarthy et al., 1997). Absence of sediments younger than 6,500 yBP in northern Georgian Bay is related to a sharp decrease in sediment flux with the termination of the Nipissing transgression due to the geology of the basin

(McCarthy et al., 1997).

Presently, a small number of studies have analyzed thecamoebians in the Great Lakes basin (Scott & Medioli, 1983; McCarthy et al., 1997; Tiffin, 1998; McLachlan, 1999; Sarvis et al., 1999a). McCarthy et al., (1997) used thecamoebians to reconstruct sedimentation rates and

paleolimnological parameters in Georgian Bay (Figure 2). The high concentrations of thecamoebians in historic sediments and other sediments deposited since the end of the Nipissing transgression record low rates of sediment accumulation.

High thecamoebian populations rich in Cucurbitella tricuspis were found in recent sediments of Severn Sound. Preceding the Ambrosia Rise, the rise in the C. tricuspis population indicates nutrients entering the bay prior to European settlement suggesting the occupation of paleo-Indians (Hurons) in Severn Sound (Figure 2) (McCarthy et al., 1997; McLachlan, 1999).

Relatively high abundances of Difflugia oblonga and Difflugia bacillifera record hypsithermal conditions (6,500 to 5,000 yBP) (Figure 2). Muds deposited between 8,000 and 6,500 yBP are

dominated by Centropyxis species (Figure 2). This reflects high concentrations of dissolved ions

resulting from high evaporation rates leading to low water levels (McCarthy et al., 1997; Scott et

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6r § I 3 17 al., 1999). Periglacial lacustrine clays contained very sparse assemblages dominated by

Difflugia oblonga (Figure 2).

1.2 Previous studies on Willoughby Bog

In the past decade various studies have reconstructed the paleoenvironment of

Willoughby Bog (Pengelly, 1990; Tinkler et al., 1992; Tinkler, 1994; Pengelly et al., 1997;

Sarvis, 1997; Sarvis et al., 1997). Earlier research concluded that Willoughby Bog was a fluvial channel that drained into Lake Tonawanda 7,500 years ago (Figure 3) (Pengelly, 1990). Sarvis

(1997) affirmed the hypothesized fluvial channel by analyzing various cross-sections of the biogenic basin. Willoughby Bog is part of the Carolinian Forest in the Niagara Peninsula

(OMNR, 1979, 1980; MacDonald, 1980). Palynology of Willoughby Bog is characteristic of the transition from the mixed deciduous forest to a full deciduous forest during the Hypsithermal

Period. The pollen record of Willoughby Bog reflects the evolution from a fluvial system to a peat margin swamp environment.

2.0 Study Areas

2.1 Georgian Bay

Georgiein Bay is an extended arm of Lake Huron (Figure 4) and water levels are a product of inflow from lakes Huron, Michigan, and Superior relative to outflow to lakes Erie and Ontario via the outlet at Port Huron, as well as effective precipitation in the drainage basin. Georgian Bay is located between 44° 30 and 46°N latitude, and 80° and 82°W longitude. Georgian Bay has a

surface area of 14,742 km^ with the long axis oriented in a northwest-southeast direction and its

drainage basin has an area of 65,723 km^ (including major lakes) (Figure 4). Georgian Bay is connected to Lake Huron by shallow narrow channels between Bruce Peninsula and Manitoulin

Island (Figure 4) (McAtee, 1977; Chen, 1981). Eastern Georgian Bay is characterized by a

nt / 19

Figure 4: The Georgian Bay basin and bathymetry (modified from McAtee, 1977),

20 shallow bathymetry controlled by the bedrock of the Canadian Shield (Figure 4) (Chen,

1981). The western section of Georgian Bay becomes deeper towards the eastern shoreline on

the Bruce Peninsula and the Niagara Escarpment. The maximum water depth is approximately

168m (Figure 4) (McAtee, 1977; Chen, 1981). The main drainage basins in northern Georgian

Bay are the Muskoka, Magnetawan, French, Temagemi, Upper Wanapitei, Lower Wanapitei and

Killamey (Figure 5). The southern drainage basins are the Sauble, Nottawasaga and Severn

(Figures). i

Tobermory is located in the Bruce Peninsula at 45° 14'N latitude and 81° 4VW longitude

(Figure 4). The Bruce Peninsula has an area of 3,852 km . The location of core LH95-041 is

offshore north of Tobermory in an ancient fluvial channel (Blasco et al., 1999).

Manitoulin Island is located in the northern section of Lake Huron between 45° 3 1 'N and

46° OO'N latitude and 81° 34'W and 83° 13' longitude (Figure 4). The locations of cores LH95-

035 and LH95-038 are offshore southeast of Squaw Island (45° 49'N lat. and 81° 39'W long.).

Various lake basins are located on Manitoulin Island, including large three lakes called

Kagawong, Mindemoya and Manitou, combining an area of 168.35 km^ (Hoffman et al., 1959).

The three main tributaries that empty into Lake Huron are Blue Jay Creek, Manitou Creek and

Mindemoya River (Hoffman et al., 1959).

Severn Sound is located in the southeastern part of Georgian Bay between 44° 40' and

44° 50'N latitude and 79° 40' and 80° OO'W longitude (Figure 4). The Severn Sound is a shallow embayment consisting of the Nottawasaga and Severn watersheds that cover a combined area of 9,252 km^. Seven tributaries drain into Severn Sound, including the Copeland and Hog creeks, and the Wye, Sturgeon, Severn, Coldwater and North Rivers (Chapman & Putnam, 1984;

McLachlan, 1999). it -

h s.r*.;:, • in-. -,..'.•!!. U^.if..

.' ' ! "L; ;< (if /x;ti ' 21

Figure 5: Georgian Bay drainage basin (modified from McAtee, 1977).

22 2.2 Willoughby Bog

Willoughby Bog is located in the Niagara Peninsula, Regional Municipality of Niagara

near the city of Niagara Falls, Ontario (Figure 6). The Willoughby Bog is designated as a

provincially significant Class 1 wetland, an Area of Natural and Scientific Interest (ANSI) and an Environmentally Sensitive Area (ESA) (Limnoterra, 1998). The Niagara Peninsula

Conservation Authority (NPCA) protects an area of about 1 .5 km^ in Willoughby Bog (Figure 7)

(Poirier et al., 1982). Its dimensions are 0.5km wide and 3km (north to south) (Figure 6)

(Pengelly, 1990). Willoughby Bog is the headwaters for three tributaries named Usshers, Tee,

and Black creeks (Figure 7). There is one tributary that flows towards Willoughby during the

spring melt out and it is called the Young Drain (Sarvis, 1997).

3.0 Georgian Bay Background Information

3.1 Climate of the Great Lakes

Under the Koppen classification the Great Lakes are in a Humid Continental-Cool

Summer Climate (Dfb) region. In general, the region experiences long mild summers where the

warmest monthly mean is below 22°C and four or more months are above 10°C, with no dry

season (Scott, 1992). The Great Lakes basin, other than Lake Superior, is designated as a humid

continental zone due to the interaction of four air masses affecting the climate (Eraser et al.,

1990). The humid continental regions are in latitudinal zones, which are most frequented by polar fronts and are in a collision zone between subtropical and sub-polar air masses. These interactions bring sharp changes to the climate in the Great Lakes. The Great Lakes are unique because different weather patterns can be affecting different areas at similar moments in the basin. Mean January and July climate maps of eastern North America show the temperature and

precipitation contrasts in the region (Figure 8). The Great Lakes region consists of four distinct ':>• (;;;/•'

ypi]; -^I'tf.i ^ 23

Niagara Peninsula

arocn^^Js^M^^jB^y

Figure 6: a: Map of the Niagara Peninsula and local wetlands (Beard personal comm., 2000). b: Closeup of Willoughby Bog and local creeks in the area (Poirier et al., 1982). V

(

J, "•: 24

Scale Icm to 33in

Figure 7: Topographic map of the Willoughby Bog and surrounding area (Scale 1 : 50 000).

Figure 8: Climate maps of eastern North America and the Great Lakes basin. A = mean January Temperature (" C), = B = Mean January Precipitation (mm), C = Mean July temperature (° C), D = Mean July Precipitation (mm), E Mean Annual Temperature (" C) and F = Mean Annual Precipitation (mm) (modified from Steinhauser, 1979).

26 seasons, variations in precipitation types and sources, temperature and humidity differences

are a major source for atmospheric moisture (Brown et al., 1980). The Great Lakes act as huge reservoirs for heat exchange and thus exert a moderating effect on the temperature and precipitation regimes over the region (PhilHps & McCulloch, 1972). The lee side of Lake Huron and Georgian Bay experiences more precipitation, more cloud cover and moderate temperatures

in the basin (Figures 8). Prevailing winds cross Lake Huron making it the most influential lake to affect the climate of southern Ontario.

3.1.1 Climate of Georgian Bay

The mean annual air temperature for Georgian Bay ranges from 5°C to 7 °C with mean daily temperatures ranging from -12°C to -6°C for January and 18°C to 19°C for July (Brown et

al., 1974; McAtee, 1977; Chen, 1981). During summer months (June, July, August and

September) the mean air and water temperatures range between 10 to 20°C and 15 to 20°C,

respectively (Table 1) (Saulesleja, 1986). The water temperature of Georgian Bay increases rapidly during the months of June and July (Appendix A, Figure 1) (Webb, 1972; Saulesleja,

1986). Georgian Bay consists of a distinct water temperature profile during these months including the warm eplimnion layer overlying the colder hypolimnion (4°C) (Webb, 1972). The

eplimnion layer increases in thickness as it cools by convective mixing, and eventually deepens to include the entire bay (Webb, 1972). Evaporation and evapotranspiration occur through the

summer and fall months from increased air and water temperatures and activated vegetational processes. From a mid-summer low, evaporation from Georgian Bay will increase through

September until reaching a peak in December (Webb, 1972).

Table 1: Climate parameters for Georgian Bay (Saulesleja, 1986). '/) 28

Water Budget: Tobermory, Ontario

25 Deficit 20 . Surplus 15 ' 10 ' so s 5

-5 -10

-)Ij.S<5-)^<(^Ozq |P (mm) Month

Figure 9: Water budget of Tobermory, Ontario illustrating the temperature and precipitation rates in the Georgian Bay basin (Steinhauser, 1979).

Peninsula influences the wind and climate (i.e. temperature and precipitation) (Figure 9). Low pressure cells produce adequate amounts of precipitation through the year by the influence of air masses and positioning of the Jetstream. The Muskoka area receives high amounts of precipitation due to the topography of the Canadian Shield (Chen, 1981). Areas with low amounts of precipitation are found in the Nottawasaga watershed in the south and in the north near Sudbury. In winter months snowfall can accumulate for several months to an approximate

depth of 280cm in the Georgian Bay drainage basin (Brown et al., 1968; McAtee, 1977; Chen,

1981; Hoey, 1981). Distribution of snowfall is related to the lake effect on the lee of Georgian

Bay and other Great Lakes. In general, as a cold air mass passes over an ice-free, warm lake in

winter months lake effect snow is produced on the shoreline.

3.2 Vegetation and Soils of Georgian Bay

The vegetation of Georgian Bay is situated within the Great Lakes-St. Lawrence

Deciduous zone characterized by Acer, Fagus, Quercus, and Pinus and is intermixed with boreal

species. This region is also referred to as the transition forest, ranging from the Midwest to the ».;•> • a 29 Maritimes (Sutton & Sutton, 1997). The transition forest lies between the coniferous

woodlands of the boreal forest in the north and the mixed deciduous temperate woodlands in the

south. The warm climate of the transition forest provides longer growing season, higher

evaporation rates and quicker snowmelts (Sutton & Sutton, 1997). Rowe (1977) subdivided the

Georgian Bay drainage basin into six vegetation sections based on climate, topography, soils and

geology (Appendix A, Table 4). These sections are Huron-Ontario, Georgian Bay, Sudbury-

North Bay, Temagemi, Algoma, and Missinaibi-Cabonga (Appendix A, Table 4).

The major soil types in the Bruce Peninsula and Manitoulin Island areas are grey brown

luvisols, melanic brunisols and humo-ferric podzols (Hoffman & Richards, 1954; Hoffman et al.,

1959; Rowe, 1977). Exposed bedrock is generally observed in the Bruce Peninsula (Hoffman &

Richards, 1954). In Severn Sound the dominant soil types are humo-ferric podzols, dystric

brunisols and organics (Hoffman et al., 1962; Rowe, 1977). North of Severn Soimd, outcrops of the Canadian Shield are exposed or consists of a thin soil cover.

3.3 Hydrology of Georgian Bay

Georgian Bay is a body of freshwater, which is part of Lake Huron. Eastern Georgian

Bay is characterized by a shallow bathymetry (0-3 7m) controlled by the bedrock of the Canadian

Shield (Figure 4) (Chen, 1981). The western section of Georgian Bay becomes deeper (73-

146m) towards the eastern shoreline on Bruce Peninsula containing a maximum depth of

approximately 168m (Figure 4) (McAtee, 1977; Chen, 1981). Georgian Bay contains

approximately 870 km^ of freshwater (Table 3), and consists of a higher ratio of drainage area to

surface area in comparison to Lake Huron (Table 2) (Figure 5). The average lake level of the

bay over eighty years of collected data is approximately 176.5m (Environment Canada, 2000)

(Appendix C, Table 10).

30 Georgian Bay is connected to Lake Huron by shallow narrow channels collectively referred to the Main Channel between Bruce Peninsula and Manitoulin Island (Figure 4) ^

(McAtee, 1977; Schertzer et al., 1979; Chen, 1981). The Main Channel is approximately 25 km wide consisting of five parallel channels of varying widths and controlled by underwater bedrock sills. Schertzer et al. (1979) distinguished two aspects that control flow fi-om Georgian Bay to

Lake Huron through the Main Channel. The first is surface flow. The net flow mainly enters

Georgian Bay through the southern channels and outflow occurs through the northern charmels.

The second aspect is the subsurface flow of the Main Channel, containing vertical stratification.

The net subsurface outflow is towards Lake Huron. The Little Current Channel connects the

North Charmel and northern Georgian Bay. Water flows from Georgian Bay to Lake Huron most of the year. For most of the year water flow is dominantly to the North Channel, while the

summer months flow is towards Georgian Bay (Schertzer et al., 1979).

Table 2: Morphological characteristics for lakes Michigan, Huron, Georgian Bay and Erie (Ristic, 1981; 1985 and 1989; Botts & Krushelnicki, 1995). Morphological Parameters

31 The drainage basins in the northern Georgian Bay region include numerous lakes, rivers and tributaries that eventually drain into the bay (Table 3) (Figure 5). The river systems

that occupy these basins flow through the Canadian Shield following bedrock structures (i.e. fault lines) with minimum down grading (Chapman &. Putnam, 1984). Some rivers may flow

through unconsolidated material (i.e. sand, silt and clay) changing their gradient. The South and

Sturgeon rivers empty into Lake Nipissing, which drains through the French River via narrow lakes and three individual channels that empty into Georgian Bay (Chapman & Putnam, 1984).

The Magnetawan River empties into Georgian Bay at Byng Inlet draining lakes Ahmic and

Wahwashkesh. The Muskoka and Moon rivers drain Lake Muskoka into Georgian Bay.

The southern Georgian Bay region is characteristic of the Nottawasaga and Severn Rivers

(Table 3) (Figure 5). The Severn River flows into southeastern Georgian Bay into Severn

Sound. The Severn River drains two lake basins that include Lakes Couchiching and Simcoe

Table 3: Georgian Bay drainage basins and surface hydrology (Chapman & Putnam, 1984). Drainage Basin

32 tributaries, the Pine and Mad rivers that flow over through the Niagara Escarpment

producing steep gradients and bedrock valleys (Chapman & Putnam, 1984).

3.4 Geology of Georgian Bay

Georgian Bay lies between the Ordovician sedimentary rocks (i.e. limestones, shales) in

the southwest and the Precambrian Shield to the east and northeast (Figure 10) (McAtee, 1977;

Chen, 1981). The Grenville Structural Province lies partly to the north and east of eastern

Georgian Bay (McAtee, 1977). The Niagara Escarpment is the most prominent topographic relief

feature in the Bruce Peninsula and Manitoulin Island region (Figure 10). It rises at Rochester,

New York through the Niagara Peninsula in southern Ontario and bends at Hamilton extending

towards the Bruce Peninsula and Manitoulin Island. In Severn Sound, the contact between the

Ordovician and Silurian sedimentary formations and the Precambrian Shield is the obvious

feature controlling the surficial geology of the area (Figure 10). The topographic relief at the

contact zone changes from an undulating landscape underlain by the sedimentary bedrock to the

flat bedrock of the Precambrian Shield.

3.5 Geomorphology and Quaternary Geology of Georgian Bay

The Quaternary geology of Georgian Bay is complex due to the glacial landforms

surrounding the basin and the bottom sediments beneath the water plane. The bottom glacial

sediments along the eastern shoreline of Georgian Bay are characterized by glacial tills,

glaciolacustrine clays (Main Algonquin), and bedrock (Figure 1 1) (Sly & Thomas, 1974;

McAtee, 1977). The Georgian Bay basin consists of an irregular bottom topography due to a

discontinuous sedimentation rate (Anderson & Terasmae, 1966; Sly & Thomas, 1974; McAtee,

1977; Rea et al., 1981; Moore et al., 1994; Rea et al., 1994; Dobson et al., 1995). The

controlling factors are the bedrock geology of the northern igneous/ metamorphic Canadian '^• >''<^: -,)/ . ;j.

:>;• ,\iy ' i!*^'^•- r —

I I I — Scale 1cm to I6)an V^^'^'P^^^^^^flTl SCTSv g 7TTS Ai-

Phanerozoic Legend Cenozoic Esker or area of eskera flTl GbdolacustriiM deposHs Paleozoic 1.*^ >ik aid cliy; minar antt bmi atd Quaternary I Beclrock; aibonMJ citaic Kdimcntry ' Terrace etcarpmcnt Recall rod: CDHcd a avftcc « covered by (fan gla:iolKiianne. gnciomasic. 1 GtocMknlal Mitmih depptfti Ityvofdrtfl lasiahne cr mane (^Mndcned

' pwalwawaKiacMMprtigteaiirivg ; M Organic depotitB j AoreMuffl

I 1 pot. mudc nd iinri aalMitediparili ^ Beach, bar or spit: f Lacustrine depotlti Precambrian ^•cioitcialrine. glaciamvint. ^ I widgrtvelly said Bid gravcJL Gbciofluvlal I ke-contact da oaHi J I p Bedrock; igncouW mctanwphK rod: npo ttcuariic or mnw ndbodi dcpoots gnvcl aid twd. mnar nil, mcludci cnr. I a avfkce or covered by (hin lijn of (kifl kanc. a*d monuoc icc-magmil ddta nd E Terrace escarpment; Pleistocene Symbols nuvlal Fluvial dcpottts i predominMify laidy siH BBB r>fH/» Sand dune or gnwi. and. all aid cliy. drumlin Itcldi area of on bMdoned fk)o4>lani icnvce rannaiu conUat in mirn sand dunes

rrrn GbckOacwtria* dMMalto Elma Till muron-Gcorglan(Huron-Gcon lobe) Bedrock escarpment r^n Bay * • * ' Trend or end moraine crett Hf antty nk to nk mMnx. modcniay

Figure 10: Quaternary geology of the Georgian Bay basin (modified from Bamett et al., 1991).

Legend

Melange, Bedrock, till and Glaciolacustrine clay

Glacial till

Glaciolacustrine clay

Mud

Scale Icm to 10km

Figure 11: Bottom sediments of Georgian Bay (modified from Sly & Thomas, 1974). -4

» • .•..»•«'- 35 Shield and southern Paleozoic sedimentary lithologies and the Main Algonquin sediments.

The Ordovician shales and limestones beneath Georgian Bay were differentially eroded by the

Laurentide Ice Sheet (McAtee, 1977). The result has been various unconformities in the basin.

During lowstands (Hough phases), bathymetric highs containing Main Algonquin (pollen zone 1)

sediments were exposed, but didn't erode due to the compaction and cohesiveness fine grained

clay materials (Dubas, 1997; Tiffin, 1998). As water levels gradually rose from isostatic

rebound (Nipissing Rise, Mattawa), yoimger sediments (pollen zones 2 & 3) were eroded by

wave activity and deposited in bathymetric depressed areas. As the water levels declined after the Nipissing Rise, sediment availability decreased resulting in another unconformity in the

Georgian Bay basin (Dubas, 1997). The recession of the Nipissing Rise caused a major decrease

in sediment supply due to the resistant bedrock lithologies of the Georgian Bay basin. However,

deforestation during European settlement brought an increased supply of unconsolidated

sediments into the basin producing a present day erosional surface (Tiffin, 1998; McLachlan,

1999).

The region around Georgian Bay can be subdivided into six general physiographic regions which include the Niagara Escarpment, Bruce Peninsula, Manitoulin Island, Simcoe

Lowlands, Simcoe Uplands and Georgian Bay Fringe (Chapman & Putnam, 1984). The

Quaternary geology of the southern region of Georgian Bay is complex due to the glacial sediments and landforms from various advances and retreats of the Laurentide Ice Sheet (Figure

10). These glacial landforms produce an undulating relief around southern Georgian Bay.

4.0 Willoughby Bog Background Information

4.1 Climate of Willoughby Bog

Willoughby Bog is located in the Niagara Peninsula and within the Koppen humid temperate-cool summer climate region. Lakes Erie and Ontario and the Niagara and Onondaga f' - ::/r ;:' ijr. » 36 escarpments affect the temperatures in the area. Weather is variable due to movement of air masses and topography creating microclimates. The annual daily temperature ranges between

8.0 to 9.0°C (Shaw, 1994). The cool winters have an annual mean January air temperature of approximately -A.7°C and the warm summers have an annual July temperature of 21.4°C (Shaw,

1994). The topography and wind stress produces extreme temperature gradients in the region.

Annual Precipitation in the region is quite uniform ranging from 807 to 995 mm/yr. (Shaw,

1994). The region receives most of its precipitation in the autumn and winter months. The interaction of convective showers and cold frontal systems during summer months and the passage of frontal systems in December and January attribute to severe weather in the region.

Approximately 84% (750mm) of precipitation falls as rain (Shaw, 1994). Precipitation is lowest during summer months. Climate data has not been recorded specifically for Willoughby Bog therefore a water budget was constructed using climate data from nearby Welland, Ontario

(Figure 12). .!^-^Hii-.'-,l 37 rubrum, Quercus rubra, Quercus alba, Populus balsamifera, Prunus serotina, Ironwood,

Betula and Quercus macroarpa (Rowe, 1977; OMNR, 1979).

The major soil types in the Willoughby Bog area are orthic humic gleysol (Welland),

gleyed gray brown luivsol (Niagara) and terric mesisol (Quarry) soils (Kingston & Presant, 1989;

Sarvis, 1997).

4.3 Hydrology of Willoughby Bog

The drainage basin can be subdivided into Usshers (20.5km^), Tee (46.5km^) and Black

(103 km^) (Figure 6) (Poirier, 1982; Sarvis, 1997). Willoughby Bog is the headwaters for

Usshers, Tee and Black creeks (Figure 6) (Sarvis, 1997). The Onondaga Escarpment affects

surface drainage towards the Niagara River. Low areas in this part of the region are generally

poorly drained and are excellent areas for bog development. Usshers and Black Creeks drain

directly into the Niagara River and into Lake Ontario.

4.4 Geology and Geomorphology of Willoughby Bog

The prominent physiographic and geological features in the Niagara Peninsula are the

Niagara and Onondaga escarpments. The Onondaga Escarpment consists of Bois Blanc

limestones and Bertie dolostones with a relief ranging between 10 to 25m in southern Niagara

Peninsula (Feenstra, 1981; Tinkler, 1994). The Salina Formation (Upper Silurian) consists of soft

dolostones and shales (Feenstra, 1981; Poirier, 1982).

The Niagara Peninsula contains unique geomorphological features. Above the Niagara

Escarpment proglacial lakes deposited varied thickness of clay producing the Haldimand Clay

Plain (3,496 km^) (Poirier, 1982; Chapman & Putnam, 1984; Sarvis, 1997). Moraines that were

formed by the retreat of the Ontario-Erie ice lobe were the Vinemount, Niagara Falls and Fort

Erie moraines (Feenstra, 1981; Chapman & Putnam, 1984). The southern Niagara Region * V 38 contains various wetland environments that include the Wainfleet Bog, Humberstone Bog

and Willoughby Bog conservation areas (Figure 6). These formations have been related to high

Lake Erie waters breaching the Onondaga Escarpment forming shallow lakes (10,500 yBP)

draining into Lake Tonawanda (Pengelly 1990).

5.0 Postglacial History in the Great Lakes (Late Wisconsinan Period)

Southern Ontario has undergone various changes in landscape from the postglacial

formation of assorted proglacial lakes, ice advances and changing drainage outlets in the Great

Lakes basin during the Late Wisconsinan period. The three Late Wisconsinan glacial stadials are

called Nissouri, Port Bruce, and Port Huron stadials (Bamett, 1992). These glacial stadials were

followed by interstadials that included Erie, Mackinaw and Two Creeks interstadials

respectively. For this study we will only concentrate on the different lake phases since the Two

Creek Interstadial.

The Two Creeks Interstade (12,400 yBP) signifies the final retreat of the Laurentide Ice

Sheet fi-om the Great Lakes basin causing major water level fluctuations throughout the system

(Appendix A, Table 5) (Bamett, 1992). Through the glacial history of the upper Great Lakes the

controlling drainage outlet has changed numerous times between Fenelon Falls, North Bay and

Port Huron. Since the Middle Holocene period. Lakes Huron and Michigan have influenced the water levels in the Erie basin. Present streams and rivers of the Erie drainage basin account for

about 13% of the discharge illustrating the importance water inflow from Huron has on Erie

water levels (Spencer, 1907; Tinkler et al., 1992; Pengelly et al., 1997). Lowstands and

highstands are reflected by low or high water levels in the Erie basin. Lowstands were generally

longer in the Erie basin due to the requirement of waters to fill the Huron basin before flowing

through the Port Huron outlet, and the duration of the highstands were shorter (Pengelly et al., !'; :',.-•': .' -iiil . v-r

'' jj-i. -i:-- f\,: ' tV 39 1997). The upper Great Lakes experienced two episodes of meltwater inflow from Lake

Agassiz (1 1,000 to 10,500 yBP and 9,600 to 8,300 yBP) resulting in a cool climate, shortening

the growing season in the surrounding area (Anderson & Lewis, 1992). During the early Lake

Algonquin stage (12,100 to 1 1,500 yBP) and the Kirkfield lowstand (1 1,900 yBP), Lake Huron

flowed into the Erie basin (Lewis et al., 1994; Pengelly et al., 1997). Lake Agassiz outflows produced the Early Lake Algonquin phase (190m asl) that entered the Erie basin increasing water

levels (Coakley & Lewis, 1985; Tinkler & Pengelly, 1995a; Pengelly et al., 1997). This glacial meltwater flowed out of Erie forming the first of three stages of Lake Tonawanda via Lake

Ontario (Figure 3) (Kindle & Taylor, 1913; D'Agostino, 1958; Pengelly et al., 1997). The early

stage of Lake Tonawanda had no prominent beach ridges suggesting a shallow lake environment

(D'Agostino, 1958). The Early Lake Algonquin phase ended in the upper Great Lakes as water was diverted through the Kirkfield outlet producing a low lake levels in the Huron basin

(Appendix A, Table 5) (Karrow et al., 1975; Kaszycki, 1985; Pengelly et al., 1997).

During the Main Algonquin phase (1 1,200 to 1 1,000 yBP) the drainage was through

Fenelon Falls into Lake Simcoe via the Trent River System and into the Champlain Sea (Table 8)

(Figure 13) (Calkin 8c Feenstra, 1985; Karrow, 1989; Bamett, 1992; Lewis et al, 1994). Water inflow from Main Algonquin that occupied the Michigan and Huron basins entered the Erie

basin (1 1,000 to 10,500 yBP) (Larsen, 1987; Rea et al., 1994). As the ice front retreated the La

Cloche saddle on the Little Current channel was exposed resulting from dropped lake levels. The

controlling outlet was the Dalles sill on the French River creating the early stages of proglacial

Lake Hough in the Georgian Bay basin (approximately 10,000 yBP) (Figure 13) (Lewis et al,

1994; Blasco et al., 1999). The water levels in Lake Huron and Georgian Bay remained low

between 1 0, 1 00 to 9,600 yBP forming Early Lake Stanley and Early Lake Hough respectively

(Figure 13) (Sly & Lewis, 1972; Lewis et al., 1994; Blasco et al., 1999). Water drained from

Superior, Huron and Michigan through the North Charmel into Georgian Bay via the Dalles sill ' ( i:.'->i/:j- '„,!

'u' .:v, /':.fi 40

Indian & Grand R. Port Huron Ice Front CM

Fenelon Falls Kirkfield Phase

Main Algonquin Port Huron Post Algonquin Ice Margin to Ottawa - Early Lake — O Stanely/Hough = Early Mattawa s Mid Lake Stanely/Hough (0 5 « p > 5^ I Main Mattawa (0 Late Lake — 00 Stanely/Hough

Q. CQ

O)c « — CO CO

Nipissing Great — •* Lakes

Algoma

3 (0 X (D 5 CM I 1

200 150 100 50

Main Outlets Lake Levels (metres ASL)

Figure 13: Postglacial lake levels of the Great Lakes and major drainage outets (modified from

Uwis et al., 1994 and Blasco, 1994). <)i^

n -.

g^sr^q bl9it)hi."H 2l'6^

'> i

Pi o>

"1r .j3 < ^ * ninM

-0 CQ 0)

(0 CO>

Q OS oor o?> oos

(J8A 89l}9m) 219 41 to North Bay, down the Ottawa Valley and into the Champlain Sea (Lewis et al., 1994;

Blasco et al., 1999). Waters from the upper Great Lakes entered the Erie basin, reactivating old spillways and forming Lake Tonawanda by flooding western New York and extending into the

southern Niagara Peninsula creating Lake Wainfleet (Anderson, 1989; Tinkler et al., 1992). This middle stage of Lake Tonawanda had the same surface area, but the depth decreased due to sediment deposition (i.e. deltas) from Lake Erie waters (D'Agostino, 1958). Lake levels rose from isostatic rebound after the Main Algonquin stage in the upper Great Lakes. The shorelines of Lake Tonawanda during Main Algonquin and Nipissing Rise are at the same elevations

(180m) (Kindle & Taylor, 1913; D'Agostino; 1958; Pengelly et al., 1997). A lowstand called the Chippewa phase occurred in the Lake Michigan basin between -10, 000 and 5,500 years ago,

from the time the North Bay outlet was uncovered to the Nipissing Rise (Hansel et al., 1985).

Hansel et al. (1985) noted that the Chippewa-Nipissing transition corresponds approximately in time with the end of the Hypsithermal of the Holocene climatic history (Terasmae, 1968), when the warmer and drier conditions of the early Holocene were replaced by cooler and moister conditions in the northern Midwest after 6,000 years ago.

The second inflow (9,600 to 8,300 yBP) fi-om Lake Agassiz caused rising water levels expanding the Mattawa phase in the Huron-Michigan basins. The Early Mattawa highstand

(155m) took place in the Huron basin around 9,500 yBP and the Main Mattawa highstand

(160m) around 8,500 yBP) (Figure 15) (Rea et al., 1994; Lewis et al., 1994; Pengelly et al.,

1997). Succession from spruce to a pine dominated forest and low 5*^0 ratios in the Michigan

basin illustrates the cool climate cheinge from the presence of Lake Mattawa (Coleman et al.,

1990; Anderson & Lewis, 1992). By 9,300 yBP the water levels dropped from Early Mattawa to

Mid Lake Stanley/ Hough lowstand, meltwater and runoff flowing easterly towards North Bay was confined fluvial channels of the La Cloche (North Channel) and Dean saddles (north of

Bruce Peninsula) (Figure 13) (Blasco et al., 1999). South of Manitoulin Island three fluvial f--' iJn>V'

(,! , ;: ^w^« -_! ii-^\!' ;i; i 42 channels, referred to as Fitzwilliam, Lucas and Flowerpot, were eroded through the

bedrock of the Dean Saddle (Lewis & Anderson, 1989; Blasco, 1994; Blasco et al., 1997; Blasco et al., 1999). In the Erie basin evidence from peat blocks that washed onshore indicate a lower

Lake Erie corresponding to the Mid Stanley/ Hough lowstand in the Huron basin (Pengelly et al.,

1997).

The Main Mattawa highstand occupied the Huron basin between 9,000 to 8,100 yBP due

to renewed inflow from Lake Agassiz (Lewis et al., 1994; Blasco et al., 1999). Lake Erie lake levels rose around 8,500 yBP as a result of outflow from the Huron basin during Main Mattawa highstand (Pengelly et al., 1997). By 8,100 yBP reduced meltwater from Lake Agassiz lowered water levels in the Huron and Georgian Bay basins activating the fluvial channels across the

Dean Saddle. This lowstand established the Late Stanley and Hough lake phases in Huron and

Georgian Bay respectively (8,200 to 7,300 yBP) (Figures 13 & 14) (Lewis et al., 1994). Around

8,000 yBP, waters flowed through North Bay outlet controlling the drainage in the upper Great

Lakes (8,000 yBP) (Figure 13) (Lewis et al., 1994). Stratigraphic evidence, palynology and

radiocarbon dates have illustrated that lake levels were lower than the level of the controlling sill in the Huron and Georgian Bay basins at 7,500 yBP (Figure 15) (Sly & Sandilands. 1988; Rea et

al., 1994; Moore et al., 1994; Dobson et al., 1995; Blasco et al., 1997; Janusas et al., 1998;

Blasco et al., 1999; McCarthy & Tiffin, 1999). The Late Lake Stanley lowstand occurred around

8,000 yBP before the Nipissing Rise (Figures 13 «& 14) (Lewis et al., 1994). The lowest lake level occurred between 7,800 and 7,400 yBP on the basis of the presence of coarse grained

sediments (Figure 15) (Rea et al. 1994). Between 8,100 to 7,400 yBP light 50 '^ values occurred during the Late Lake Stanley lowstand and the deposition of coarse material. The last time

glacial light isotopic waters entered the Huron basin was 7,600 yBP (Rea et al., 1994). The Late

Lake Stanley low water unconformity matches the coarse grained material (light blue reflector of ' ;.J . ,,5 «:, ijti/UMJf) ;.'( 43

Late Hough/ Stanley mt^ 44 Rea et al., 1994) confirming a lake level lowstand in the Huron basin during the Mattawa stage (Figure 15) (Rea et al., 1994; Lewis et al., 1994). The lowering of the water levels caused

Late Lakes Stanley and Late Lake Hough to gradually become closed basins in the Huron and

Georgian Bay basins, respectively (Figures 15 «& 16). The geologic processes can not explain this lowstand in postglacial history and is the major focus of this study. Climate change may

control lake levels through time by precipitation and temperature fluctuations (Larsen et al.,

1985). Isostatic rebound continued to raise the North Bay outlet, gradually v/ater levels began to rise around 7,200 yBP in the Huron and Georgian Bay basins continuing through to the Nipissing

Rise (5,500 yBP) (Figure 16) (Lewis et al., 1994; Blasco et al., 1999).

In the Erie basin, Lake Tonawanda existed in the Niagara Region and Western New York prior to 7,500 yBP (Figure 3) (D'Agostino, 1958; Pengelly, 1990; Pengelly et al., 1997). High water levels from Erie breached the Onondaga escarpment and flooded low lying areas to the

north creating Lake Wainfleet (Flint & Lolcama, 1985; Pengelly, 1990; Tinkler et al., 1992).

Water levels rose in Lake Erie around 7,000 yBP due to the climate becoming wetter and warmer

(Pengelly et al., 1997). Colman et al. (1994) discuss the lowering of waters in the Michigan

Basin around 7,000 yBP from ostracode evidence and high total dissolved solids in the basin.

Water from the upper Great Lakes gradually flowed into the Erie basin increasing lake levels, but lowering lake levels in Lakes Michigan and Huron. North Bay remained the drainage outlet during the Nipissing Rise as Georgian Bay water levels rose, until 5,500 yBP when the main outlet was changed to Port Huron outlet and has remained to the present day (Figure 17). There

are two Nipissing phases, Nipissing I (180m) (5,500-4,700 yBP) and Nipissing II (178.5m)

(4,500-3,700 yBP) (Figures 16 «fe 17) (Pengelly et al., 1997). Lake Erie waters elevated by the

Nipissing Rise II flooded low depressional areas within the Niagara Region and Western New

York reforming lakes Wainfleet and Tonawanda (Pengelly, 1990). Between 3,500 and 2,900

Figure 16: Nipissing transgression around 7,000 yBP (modified from Larsen, 1987).

Nipissing Rise

46 yBP the Lyell/ Johnson sill was eroded by Niagara Falls that drained Lake Tonawanda and

Lake Wainfleet (Figure 3) (Pengelly, 1990; McCarthy et al., 1994; Tinkler et al., 1994; Pengelly et al., 1997). Pengelly (1990) provided three radiocarbon dates of Willoughby Bog of 7,670 +/-

240 yBP at 290 cm, 7,470 +/- 180 yBP at 225cm and 6,900 +/- 100 yBP at 195cm. Around

6,700 yBP the source of water from the Niagara River was cut-off creating a low energy enviromnent to trap fine grained sediments (Sarvis, 1997). These small lakes were gradually cut-

off as Lake Erie water levels declined once the Lyell/ Johnson sill was breached and initiating bog formation in the low relief depressions. Bog formation took place in the region approximately 3,700 yBP as the Nipissing water levels receded (Donaldson, 1987; Nagy, 1992;

Douglas, 1997). After the Nipissing II rise there was another lake level rise in the Michigan

Basin called the Algoma stage (3,500 yBP) (Larsen, 1985). Between 3,400 to 3,000 yBP water

levels in the Lake Erie basin were falling due to the Lyell/ Johnson sill eroding, but lakes Huron,

Michigan and Ontario were rising (Pengelly, 1990; Tinkler et al., 1992; McCarthy (1986);

Pengelly et al., 1997). Other less significant highstands in the Erie basin occurred from inflow

by the upper Great Lakes and a warmer/ wetter climate (Pengelly et al., 1997).

5.1 Glacio-Isostasy in the Great Lakes

The major geological process that affect lake levels in the Great Lakes basin is isostasy, commonly referred to as isostatic rebound. When the ice sheet begins to melt, the crust will rebound differentially, rising by the greatest amount and rate where the ice was thickest (Gray,

1996). The thickness of the Laurentide Ice Sheet depressed the Great Lakes basin. As the ice sheet retreated, the land has slowly rebounded affecting the drainage of the lakes (Larsen, 1985,

1987, 1994; Lewis & Anderson, 1989; Lewis et al. 1994). The gradient of a given terrace and the uplift rate, which increases exponentially toward the former ice centers, shows a relation between altitudes of former shorelines and time and distance from the center of ice loading .•!!': -: "i \umy-3i ,

i:}Hi-r> Tifll t. 47 (Larsen, 1987). In the Michigan and Huron basins, the southern part of the lakes have a relatively small rate of uplift (0.05 and 0.07 m/ century, respectively) (Larsen, 1987). Beach ridges and abandoned shorelines are the main geomorphological features to study in estimating the differential uplift in the Great Lakes basin. Relating lake levels to dated beach ridges establishes a geological history of events (Larsen, 1985, 1987, and 1994). Beach ridges indicate constant rates of uplift that increase northward across the region (Larsen, 1994). As the Great

Lakes basin rebounds, abandoned shorelines in the northeast will rise and in the southwest older shorelines will be submerged under water. Major lake phases in the Great Lakes basin resulted from isostatic rebound diverting water flow to controlling drainage outlets. For example, the

North Bay outlet controlled lake levels for the upper Great Lakes from 7,400 to 5,000 yBP, then

North Bay rose diverting water flow to the Port Huron outlet in the lower Huron basin. Through the Holocene period, isostatic rebound has been responsible for the majority of lake phases and outlet charmel changes. The 7,500 yBP lowstand in Georgian Bay is proposed to be related to climate change instead of isostatic rebound because the water level was lower than the elevation of the lowest known outlet (which was North Bay at the time).

6.0 Palynology

Pollen and spores have been used in Quaternary sediments to reconstruct past climate and vegetation conditions. Pollen and spores are diverse and most can be identified to the genus level. Palynologists use modem plant species and ecological conditions to reconstruct

paleoenvironments (i.e. the 'modem analogue' approach) (MacDonald, 1996). Fossilized pollen and spores are used to determine biostratigraphic sequences, correlate geographically and interpret paleoecological conditions (Jarzen & Nichols, 1996). Vegetation is responsive to

climate change and adapts to the surrounding environment (Table 1). Present day forest stands in Ontario can be related to physical and climatic parameters in the Georgian Bay region and in 1 (. 48 southern Ontario. Modem climatic conditions are reflected by the appearance of arboreal pollen adaptive to relatively warm winters and moist soils (e.g. Acer, Tsuga and Fagus) at the expense of trees capable of withstanding cooler temperatures and dry soils (e.g. Pinus and

Quercus) begirming around 7,200 yBP (McCarthy & Tiffin, 1999). In addition, pollen records can be zoned by comparison with established stratigraphies, and thus be used to correlate and date sediments (McAndrews, 1994).

Table 4: Pollen zonation in Georgian Bay providing an estimated age of the sediments in relation to vegetation change in southern Ontario (McAndrews, 1994). Pollen Zone L , ']-h-,Or ' !';(

''",v,h(^u',.vi'

i .;-'f;;

I.''',

A /i A.?., I i 49 Palynological reconstructions could be misinterpreted because different species produce different amounts of pollen. For example, one pine tree can produce 6 million pollen grains and their abundance will overshadow the other species (Birks & Birks, 1980). This does

not mean a pine forest dominated the landscape. Therefore abundance values must be first corrected in order to properly interpret the paleoenvironment. The paleorecord interprets the regional vegetation that can lead to a paleoclimate reconstruction.

Other sources of error associated with palynological interpretation are pollen dispersal mechanisms, variations in sedimentation and redistribution by burrowing biological organisms.

For example, wind can disperse pollen over great distances and various depositional envirormients (i.e. lakes, oceans, wetlands, and rivers). The bi-saccate shape of pine and spruce pollen allows grains to be transported to the middle of the ocean. Pollen size, density, shape and settling velocity determines where pollen grains are deposited in lake and wetland basins. Pollen sedimentation in lake basins could be redistributed either horizontally or vertically in the bottom sediments by wave activity, biological and/or geological factors (Birks & Birks, 1980). Pollen grains are generally abundant in shallower water levels. The movement of the littoral zone deposits the pollen rich sediment to the lake's center equalizing distribution throughout the basin

(Birks & Birks, 1980). Burrowing animals in lake and wetland environments may mix the sediment causing smearing and misinterpretations in the paleorecord. Pollen grains are poorly preserved in aerated soils in comparison to lake sediments and peaty soils. Pollen redeposition dictates that younger pollen grains may be incorporated into an older stratigraphic layer producing misinterpretations of the paleorecord. Biological agents can rework the sediment near the surface producing an inaccurate reconstruction.

50 6.1 Palynology and Wetland Environments

Wetlands were the first terrestrial depositional environments examined by Quaternary palynologists (MacDonald, 1996). Input of pollen is mainly through wind and gravity transport by overhanging vegetation, but pollen could be transported by streams depending on wetland environment. A spherical grain will descend vertically whereas a vesiculate grain will move horizontally (Birks 8c Birks, 1980). The fossilized pollen records of wetlands are used to reconstruct regional paleoenvironments, but they tend to be highly biased by local vegetation.

Pollen is preserved in wetlands when rate of the accumulation of organics exceeds the rate of decay of palynomorphs. As organic material accumulates, capillary action raises the water table producing an anaerobic environment for pollen preservation (MacDonald, 1996).

6.2 Palynology and Lake Environments

Lake sediments contain autochonuous organic material produced by aquatic plants and animals, and terrestrial material transported by gravity, wind and water. In lakes, pollen generally settles to the bottom of the basin and are preserved by anoxic conditions of bottom

waters or buried rapidly by sediments (MacDonald, 1996). Pollen deposition in lakes is affected by grain shape. Wave activity, currents and bioturbation effect lake sediments. Sediment focusing occurs when high rates of accumulation occur in deep zones rather than in shallower

areas. The bathymetry of a lake basin is very important for sediment and pollen deposition.

Sediment focusing can impact the pollen accumulation rates in the basin (MacDonald, 1 996), because the deep areas of the lake basin act as a "sink" for both pollen and sediment deposition.

6.3 Transfer functions

Transfer fiinctions provide quantitative estimates of past climate on the relationship

between modem climate and pollen deposition (Arigo et al., 1986; MacDonald, 1996). Modem ))>' [i: :»• e

51 vegetation analogues and geographical regions are required with the transfer functions for paleoenvironment reconstruction (Maycock, 1963; Bartlein & Whitlock, 1993). The mathematical regression equations establish the relationships between specified climate variables to pollen predictors (Bartlein & Whitlock, 1993; MacDonald, 1996). The application of the equations to the fossil pollen grains provides an estimated paleoclimatic reconstruction in southern Ontario (Eq. 1, 2 & 3). Modem climate analogues are applied to the paleoclimatic information in order to reconstruct the paleoenvironment.

Equation 1: Regression equation for January Mean Temperature ("C) (Bartlein & Whitlock, 1993) Calibration Region: 40-55°N, 75-87°W

Pollen Sum = Abies + Acer + Alnus + Betula + Carya + Fagus + Fraxinus + Gramineae + Picea + Pinus + Quercus + Tsuga + Ulmus

January = - - Temp. 5.348 + 0.564 Acer"^ 0.58 1 Alnus^^ 0.256 Betula"^ + 1 .042 Carya°^^ - 1 .282 Picea°^ 1.740 Pinus°^^ + 1.979 Quercus°^^ + 0.776 Tsuga^^ - 0.646 Ulmus°^ R^= 0.949; F=236.92; Pr= 0.000

Equation 2: Regression equation for July Mean Temperature (*C) (Bartlein & Whitlock, 1993) Calibration Region: 40-55 °N, 75-87 °W

Pollen Sum = Abies + Acer + Alnus + Betula + Carya + Fagus + Fraxinus + Gramineae + Picea + Pinus + Quercus + Tsuga + Ulmus

July Temp. = 20.767 - 0.253 Alnus°^ - 0.160 Betula"^ + 0.372 Carya"^^ + 0.170 Fraxinus"^ - 0.294 Picea^^

- 0.81 1 Pinus°^^ + 1.097 Quercus°^^ R^= 0.903; F=154.92; Pr= 0.000

Equation 3: Regression equation for Annual Precipitation (cm) (Bartlein & Whitlock, 1993) Calibration Region: 40-55 °N, 75-87 °W

Pollen Sum = Abies + Acer + Alnus + Betula + Carya + Fagus + Fraxinus + Gramineae + Picea + Pinus + Quercus + Tsuga + Ulmus

Annual Precip. = 65.552 + 1.624 Abies + 0.708 Betula + 0.589 Carya + 0.298 Gramineae + 0.132 Picea + 0.402 Quercus + 0.865 Tsuga - 3.693 Ulmus"^ R^= 0.680; F= 12.33; Pr= 0.000

Although there are general problems with the transfer function method, they can be useful tools as long as taphonomic factors are taken into account and modem analogues exist for the fossil pollen assemblages (Guiot, 1990; Bartlein & Whitlock, 1993). The assumptions for the transfer fianction method are modem vegetation from pollen data is assumed to be in equilibrium with the modem climate on a temporal and spatial scale. A problem in recognizing climate ui:'-

' .' ' . i 52 change through the paleorecord is interpreting the fossiHzed pollen and the vegetation response to climate change (Bartlein & Whitlock, 1993). Environmental and depositional differences may cause problems in applying transfer functions to different areas. Previous vegetation and climate may have an impact on present or past vegetation landscapes. Lastly, human activity has affected present climate and vegetation through deforestation and the increase in greenhouse gases (Guiot, 1990; MacDonald, 1996).

'\ ] 6.4 Holocene Climate of Southern Ontario

The Holocene Climate of southern Ontario can be subdivided into four stages which include the postglacial period of a cold and drier climate that ended around 7,800 yBP, a warmer and drier climate between 7,800 to 6,000 yBP, the Hypsithermal period of warmer and moist conditions (6,000 to 700 yBP) and a present day cool temperate climate (Figure 1 8)

(McAndrews, 1981; Edwards & Fritz, 1988; Edwards & Buhay, 1994; Edwards et al., 1996).

Between 12,800 and 10,000 yBP the mean July temperature was approximately 9°C with a rapid

warming at the Pleistocene/ Holocene boundary to 19°C (McAndrews, 1994). Increased meltwater inflow from Lake Agassiz (10,000 to 8,000 yBP) created cold localized climates over the proglacial Great Lakes delaying seasonal warming, resulting in a reduced growing season that effected the surrounding vegetation (Lewis & Anderson, 1992).

The 50'* isotope record derived from ostracodes for southern Ontario is characterized by low values during the late glacial and Early Holocene rising to a maximum around the Main

Hypsithermal (5,000 yBP) before declining to values of present day sometime after 3000 yBP

(Figure 18) (Edwards & Fritz, 1988; Edwards & Buhay, 1994; Edwards et al., 1996). The expansion of cold glacial Lakes Agassiz and Objiway-Barlow (9,600 to 8,300 yBP) of the '(J /J' 53

6"0 (ppm) ARHr/o)

-16 -14 -12 -10 -20 -10

modem

2000

main hypsithermal 4000 warm/ moist

6000 early hypsithermal warm/ dry

Late Hough/ Stanley 8000

postglacial cold/diy 10000

12000

Figure 18: Paleoclimate curve for southern Ontario (modified from Edwards et al., 1996).

1200 X — .^. T 55 Mattawa phase would have cooled the overlying air masses suppressing the summer climate around southern Ontario (Figure 19) (Anderson & Lewis, 1992).

During the Early Holocene there was a good correlation between temperature and humidity until the Middle Holocene when Arctic air diminishes and is replaced by moist Atlantic

air mass (Figure 19) (Edwards & Fritz, 1988; Edwards & Buhay, 1994; Edwards et al., 1996). In the Early Holocene a difference between rising temperature and humidity around 7,500 to 6,000 yBP was caused by an increase of warm, dry Pacific air mass (Figures 18 & 19) (Edwards et al.,

1996). The Holocene climate after the retreat of the Laurentide Ice Sheet became gradually warmer around 7,500 yBP, peaking during the Hypsithermal period approximately 5,500 yBP

(McAndrews, 1994). Humic layers in dunes along the Erie shoreline indicate wetter climates and high lake levels during the following periods 3,400 yBP +/-80, 3,070 yBP+/-120, 2,170

yBP+/-80, 1,350 yBP+/-90, 820 yBP+/-130 and 430 yBP+/-70 (Pengelly et al., 1997). McCarthy and McAndrews (1988) land evidence of a rapid water level rise in Lake Ontario, which they attributed to climatic causes, approximately between 3,000 and 2,000 yBP.

Present day cool/ temperate climate in southern Ontario is affected by the convergence of four air masses that includes the following: cool/ dry Arctic air mass, warm/ moist tropical air from Gulf of Mexico, cool/ moist air from the north Pacific and mild/ dry air from west Pacific

(Eraser et al., 1990).

7.0 Thecamoebians

Thecamoebians, testae amoebae or arcellaceans (protozoa) are benthic freshwater imicellular animals, which secrete a single silica, organic, or calcareous rigid shell (autogenous

test) or build it from their surrounding environments (xenogenous test) (Figure 20) (Medioli &

Scott, 1983; 1988; Scott & Medioli, 1983; Scott et al., 1991; McCarthy et al., 1995; Scott et al.,

1999). Because of their small size (45-300 ^m) they are abundant in most freshwater sediments. !.:<•! . »1 . i: 56

Arcella vulgaris Centropyxis aculeata Centropyxis aculeata " "aculeata" "discoides

Centropyxis constricta Centropyxis constricta Pontigulasia compressa " "spinosa "constricta"

Difflugia protaeiformis Difflugia protaeiformis Cucurbitella tricuspis " " "amphoralis "claviformis

Difflugia bidens Difflugia urceolata Difflugia oblonga " "urceolata" "spinosa

Difflugia oblonga Difflugia oblonga " " "oblonga "tenuis

Figure 20: Common Thecamoebian assemblages in Georgian Bay and Willoughby Bog. All SEM photographs are from Kumar and Dalby (1998).

57 Statistically significant numbers of thecamoebians can be obtained using small volumes of sediment (-1-10 cc) (Decloitre, 1953; Medioli and Scott, 1983). Thecamoebians are good paleoecological indicators because they have diverse and abundant populations, durable in the

fossil record and are easy to collect and separate from sediments (Warner, 1990; Scott et al.

1999). Because they are short-lived protists with high reproductive rates, they are generally more sensitive to changing environmental conditions than the surrounding vegetation (McCarthy et al., 1995), and being benthic, they are generally more sensitive to local envirormiental

conditions (i.e. climate, water levels, pollution, dissolved solids).

Autogenous tests are seldom fossilized due to organic content. The majority of fossilized thecamoebian assemblages are xenogenous. They build their tests by agglutinating foreign

particles (i.e. quartz, diatom frustules, siliceous sponge spicules and glass fragments) with an autogenous cement from the surrounding environment (Figure 20) (Medioli & Scott, 1983; Scott

et al., 1999). The nature of xenosomes is exceedingly variable, but is associated with sediment availability in the environment (Medioli & Scott, 1983; Medioli et al., 1987; Scott et al., 1999).

The test shape and presence of an aperture (Medioli & Scott, 1983) identify thecamoebian taxa.

Their tests resist decay in certain environments (i.e. low pH conditions, various water temperatures, nutrient availability) enabling a fossilized record (Warner, 1990; Ellison, 1995).

Thecamoebians are ubiquitous and live in a wide spread of environments including lakes, bogs, rivers, ponds, and moist environments such as mosses and wet soils (Medioli & Scott,

1988; Warner, 1990). Although their modem biogeographic distribution and environmental tolerances are not very well known, studies in North American lakes indicate that thecamoebian

distribution is controlled by environmental factors such as climate, salinity, and water and

sediment pollution (Scott and Medioli, 1983; Patterson et al., 1985; Honig and Scott, 1987;

Collins et al., 1990; Burbidge and Schroder-Adams, 1998; McCarthy et al., 1995; Scott et al.,

1999). Thecamoebians have been used in estimating the remediation rate of the sediments in a ,(•:•: ^'

'. ,: 'i< \ ;. J,* :'.'

•: • H' •..! i, 'f i

./ ii.. !,,?:- 'i*i , -.''N/f. lacustrine environment (Scott et al., 1999). They are near the bottom of the food chain and tend to reflect a change in the environment over time. There is a higher diversity of thecamoebians in lake sediments than in wetland environments (Warner, 1990). Xenogenic

thecamoebians tend to dominate lake environments due to high sediment load (Scott et al.,

1999), in comparison to a wetland environment where dominant species secrete their own tests

due to different sediment supply. ' •

Morphological response of thecamoebians to environmental conditions has recently been

exploited by various workers (e.g. Reinhardt et al., 1998) who have linked various "strains" with different environmental tolerances. In environmental assessments thecamoebians were used as

proxies for arsenic and mercury contamination in Ontario lakes (Patterson et al., 1996; Scott et

al., 1999). Abundances of Difflugia protaeiformis "claviformis" in these northern lake environments have been associated with high levels of pollutants (e.g. Hg, As, Cd, Cr, Cu, Pb)

(Scott et al., 1999). The Centropyxid species can be related to cold temperatures, low salinities and nutrient levels, oligotrophic conditions and Hg and As contamination (Decloitre, 1953;

Decloitre, 1956; Scott & Medioli, 1980b; Schonbom, 1984; Patterson et al., 1985; Honig &

Scott, 1987; Scott et al., 1999). They do not reflect a concentration of Hg compared to higher organisms which tend to bioaccumulate Hg in their tissues (Scott et al., 1999). In coastal environments, thecamoebians are applied with foraminifera for sea level reconstruction.

Presence of thecamoebians indicate a freshwater environment and a low sea level (Scott et al.,

1999). In Europe thecamoebian research has interpreted moisture conditions in peatlands

(Tolonen, 1966; Aaby, 1976; Aaby & Tauber, 1975; Barber, 1981; Beyens, 1985). Arcella artocrea and Heliopera shagni occur in wetland environments and high organic sediments

(McCarthy et al., 1997; Scott et al., 1999). w,-

y. .

'> '.i. '

l:<'\

• r. '; S / t: fKi

- )•> .'>V»I .!V<''vS l« ., .-I . i;, Jo 59 7. / Problems of Thecamoebians as Paleoecological Indicators

The major problem with thecamoebians is the taxonomy, complexity and variability of the test (Medioli & Scott, 1983). Their taxonomy remains somewhat problematic due to the variable nature of the test. This variability has resulted in the identification of thousands of morphotypes (Medioli & Scott, 1983). The organic component of the test matrix confirms the validity of many species (Ogden & Ellison, 1988), but environmental conditions also appear to affect test morphology (Medioli & Scott, 1983; McCarthy, 1984). Thecamoebians have been subdivided into thousands of phenotypes causing difficulty in interpreting the paleoecology of the fossil record (Medioli & Scott, 1988). Taxonomists working with living organisms use soft diagnostic features not recognized in fossilized thecamoebians (Ogden & Ellison, 1988).

The effectiveness of thecamoebians as paleo-indicators requires fiirther ecological research. Thecamoebians as the only paleoenvironment proxy are insignificant due to the lack of information on assemblage ecological conditions. Thecamoebian analysis requires correlation

with a well known proxy (i.e. palynology, diatoms) in order to reconstruct paleoclimatology, paleohydrology and paleoecology environments. Arboreal pollen and spores provide an extensive paleoecological background in reconstructing the paleoclimate. The vegetation and climate relationship allows arboreal pollen to be ecological foundations for thecamoebians as

interpretive paleo-indicators (McCarthy et al. 1997). Most thecamoebian studies have focused on taxonomy, and less on ecological preferences. This has caused a major problem in paleoenvironmental reconstructions. Limited research on living thecamoebians may be the direct cause for paleoecological misinterpretations. Other articles (Medioli & Scott, 1988; Warner,

1 990) have inferred that thecamoebians are not abundant in lake environments. This study demonstrates that thecamoebians thrive in Georgian Bay sediments and illustrate an abundance and diversity in assemblage populations and environmental conditions. Most of the *? 60 thecamoebian assemblages were analyzed from lake and wetland sediments. Further

research should examine fluvial sediments compared to the Willoughby Bog data.

8.0 Methodology

8.1 Field Methods

8.1.1 Georgian Bay

In 1995, the Bedford Institute of Oceanography- Atlantic division in Halifax, Nova Scotia investigated Tobermory and Manitoulin Island on cruise number LH95-800 in northwestern

Georgian Bay (Table 5). Core 041 was taken off of the shoreline of Tobermory in the ancient channel between the Bruce Peninsula and Manitoulin Island (Figures 21). The Tobermory

airphoto was taken in 1994 illustrating present conditions in the area (Scale 1: 50,000) (Figure

22). Cores 038 and 035 are located near Manitoulin Island (Figure 21). The cores were described and tested for bulk density, magnetic acoustic velocity and susceptibility by technicians at the GSC-Atlantic, where they are curated.

Table 5: Georgian Bay cores locations at Tobermory, Manitoulin Island and Severn Sound (* indicates that analysis was performed in the referenced study). Core ,'l>

.JMfi! 61

TOBERMORY M '•

: *.

LH 95-035

• LH 95-038 Depths in feet

Manitouiin

land (Q

f^'wi>' 7.

'',•''1'"^

',/ ' • - '|il^P'

/ ' X 'H (•ii, ^f ,^i.

..>;? 4 / ,; -.ii''. 5 • ;^ \

' ^ '!- '- t^-">;^'

<,, --,»r •^. ^-.-^;^. , ^s

;n> t~. -60 Georgian ^ ~B

LH 95-041 HURON '"^:i r^ Tobermory" -^f 8ruc,>3 Peninslll^

Figure 21: Core Locations from Tobermory and Manitouiin Island.

Figure 22: Tobermory airphoto showing core LH95-041 (Scale 1:50 000) (Natural Resources, 1995).

In 1997, the Bedford Institute of Oceanography-Atlantic team investigated Severn Sound on cruise number LH97-802 in southeastern Georgian Bay in collaboration with the Ontario

Ministry of Natural Resources at Wye Marsh Conservation Centre and Midland Remedial Action

Plan (RAP) (Table 5). All cores were obtained in the Severn Sound embayment (Figure 23). The

Severn Sound airphoto was taken in 1994 illustrating present conditions in the area (Scale 1

50,000) (Figure 24). Cores 020, 026 and 036 were obtained using a Lehigh Gravity Core and cores 032 and 035 were obtained using a Benthos Gravity Core.

8.1.2 Willoughby Bog

In September 1998, a soil transect was constructed to define the biogenic basin of

Willoughby Bog using a soil probe (5m in length). The task was to re-soil probe transect-A

(Sarvis, 1997) north of Sauer Rd. since old markers were lost or torn down in passing years. The

Willoughby Bog airphoto was taken in 1 994 illustrating the biogenic basin at present conditions

in the area (Scale 1 : 20,000) (Figure 25). The area of the biogenic basin is approximately 1 .026 km^ (Figure 25). The transect started 180m east from the forest edge, then 130m north to initiate soil probing (Figure 26). The new transect (referred as transect-S) followed a similar path, but observed slightly different depth measurements (Figure 27). The maximum soil depth was

4.95m observed at 360m east from the forest edge (Figure 27). Two cores were taken along the soil transect-S north of Sauer Rd. using a modified Livingstone sampler (Wright, 1967). Cores

WB4 and WB5 were taken 360m and 300m, respectively east from the forest edge (Figure 25).

In October 1998, two additional cores were taken along transect-S. Cores WB6 and WB7 were taken 390m and 405m, respectively east from the forest edge (Figure 25). -».- > iT

•II .1 64

Figure 23: Core locations from Severn Sound and bathymetry.

Figure 24: Severn Sound airphoto showing core locations (Scale 1 : 50 000).

BPI y,]'-. ^' .MB-'-'' K *^1BB|'

- "li: '•' "iSl ;ft?'' ^- • J Netherby Road

vi'.^r

Figure 25: Airphoto of Willoughby Bog. Red outline indicates the biogenic basin (Scale 1: 20 000).

67 N

Legend

Biogenic Basin

Vegetation

Transects Survey Lines

Scale 1 cm to 120m Core Sites

Figure 26: Soil transects and core locations at Willoughby Bog (Sarvis, 1997).

Willoughby Bog Cross-section along S - S' Transect s S' ~3o CO 90 3o iJo 9 iS So aC 3o~~

I I L_ a GkB-lSB

Glaciolacustrine clay boundary

J5- cUb-USb fi-

n ' 68 8.2 Laboratory Methods

8.2.1 Stratigraphy and Thermal Analysis (LOI)

8.2.1.1 Georgian Bay

Loss-on-ignition (LOI) analysis followed the procedure of Dean (1974). A volume of sediment was obtained from Severn Sound cores and placed in a pre-weighed crucible and

heated at a temperature of 1 1 0°C for 24 hours to determine the percentage of moisture. The

amount of organic carbon ignited is determined after one hour from the difference between the

550°C weight and dry weight. The weight difference between 550°C and 1000°C determines the amount of carbon dioxide evolved from carbonate minerals (Dean, 1974). Calcium carbonates is calculated by the amount of carbon dioxide lost between 550°C-1000°C (Dean, 1974). Loss-on- ignition diagrams were plotted using CANPLOT, a FORTRAN computer program (Campbell &

McAndrews, 1992).

Radiocarbon dating is a reliable method in providing a numerical and calibrated age for an organic substance. The calibrated ages were received from Howard Melville of Brock

University 's-Jaan Terasmae Radiocarbon Lab in the Earth Sciences Department.

8.2.1.2 Willoughby Bog

The Willoughby cores were described according to colour, texture, decomposition and

plant material (i.e. wood). The clay/ boundary layer was observed in all cores. The clay/ peat boundary was observed at 3.61m of core WB4, 1.32m of core WB5, 2.96m of core WB6 and

1.54m of core WB7. The composition for all of the cores was described using loss-on-ignition

(Dean, 1974). yfl

P'^}i.Xi\i^'^^ iwh 69 8.2.2 Pollen Analysis and Transfer Functions

8.2.2.1 Georgian Bay

Transfer functions were produced from the Georgian Bay palynological data (Dubas,

1997; McCarthy et al., 1997). The paleoclimate information was obtained using versions of the

CANPLOT software called CANJAN, CANJULY and CANRAIN (Campbell & McAndrews,

1992; McAndrews personal comm., 2000). The geographical region (g) used for this study was according to Bartlein and Whitlock (1993).

Transfer functions were conducted for three independent sites in the Georgian Bay basin to compare paleoclimate data (Figures 28- 30). These locations are called Wye Marsh, Porqui

Pond and Shouldice Lake (McAndrews personal comm., 2000) (Figure 4). Wye Marsh (44° 44N lat. and 79° 51W long.) and Porqui Pond (44° 17N lat. and 79° 77W long.) are located in the

Severn Sound area and Shouldice Lake (44° 5N lat. and 81° 42W long.) is found in the Bruce

Peninsula.

8.2.2.2 Willoughby Bog

Cores WB4 and WB6 were prepared and analyzed for pollen. The pollen analysis procedure begins by obtaining 1.0ml volume of sediment at desired depths in the core.

Approximately lOmL of Calgon is added until clay particles are dissipated. The sample is then sieved using a lOwm Nitex mesh. The sediment was washed with 10% HCl to dissolve any

carbonates in the sediment and the Lycopodium clavatum tablet. Approximately 1 OmL of

hydrofluoric acid (HP) is then added to the samples and placed in a hot water bath for one hour.

The HF eliminates any diatoms or silica elastics from the sample. After the HF procedure the final steps include adding lOmL of glacial acetic acid to the samples. Glacial acetic acid removes water from the sample prior to initiating the acetylosis process. This process breaks

down the organic material in the sediment. It is accomplished by initiating an exothermic VM :(!.,

• a '..i ' .{

,^: ^<' :! ; !•.(, 70 h 71 ^ ta

"" tso 3 -S

S3 .2? It e s u X>

•- .s T3 S Uca cu * .- >-.

Oh -'--

O§1o

r c

"a "o £ ex O C3 1.1

^~ ^^^

iii:^ 03 ^el o 00 ^-^-^ .r

1) «*«t* ,o

il is VvV-' E 2 B -^ \ A^--^--\ ^ 60:5 ca (J CO \ b \%

It Ji,-^

(un) ifldsa '

73 reaction through the combination of one part sulfuric acid and nine parts acetic anhydride.

This procedure was necessary with the Willoughby Bog sediments because the high organic content interferes with palynomorph analysis.

Pollen was identified using McAndrews et al. (1973), Kapp (1969), Hyde & Adams,

(1959) and Bassett et al., (1978). Pollen diagrams were produced using CANPLOT (Campbell

& McAndrews, 1992). Pollen diagrams were produced for each core. Transfer functions were produced from the same software as the Georgian Bay samples. A minimum of 100 pollen grains was identified from each sample (Appendix D, Tables 3-8).

8.2.3 Thecamoebian Analysis

8.2.3.1 Georgian Bay

Cores 035, 038 and 041 from northwestern Georgian Bay and 020, 026, 032, 035, 036 from Severn Sound were sampled and analyzed for thecamoebians. Depending on the age of the sediment and sedimentological rates the volume of sediment varied at certain depths for this analysis. Thecamoebian samples were taken at appropriate depths in the distinguished pollen zones (Appendix B). The volume of sediment was measured and sieved using a 45 wm mesh screen under a warm consistent spray of water. Samples were then stored in a container for later examination. Thecamoebians were identified according to Kumar and Dalby (1998). For each sediment sample, a minimum of 300 specimens was identified for statistical purposes.

Thecamoebian diagrams were produced using the CANPLOT (Campbell & McAndrews, 1992).

8.2.3.2 Willoughby Bog

Willoughby Bog cores WB4 and WB6 were sampled and analyzed for thecamoebians.

The samples were taken near the clay/ peat boundary. The same thecamoebian procedures used

74 in the Georgian Bay samples were applied. Thecamoebian samples were taken at appropriate depths in the distinguished pollen zones (Appendix D).

8.3 Water Budget Modeling

The modem fluctuating water levels of the Great Lakes have been strongly affected by climate. If present climate change has lowered lake levels than similar results could have occurred in the past when there was a warmer climate. The hydrological cycle describes the movement of water through a system, powered by solar energy and the ability of water to change phases (Scott, 1992). The water supply on the land surface depends upon several factors in the hydrological cycle, including rates of precipitation, evaporation, stream flow and subsurface flow. Other parameters that are not considered in the water budget are human activities. This includes deforestation, urbanization, irrigation, hydroelectric, flood control, land drainage and abstraction of groundwater and surface water for domestic and/or industrial use (Ward &

Robinson, 1992).

The modem water budget for Georgian Bay is subdivided into seven components which includes precipitation, evaporation, runoff, outflow/ inflow from the Main Channel and North

Channel and groundwater (Equation 4) (Appendix C). A modem water budget was modeled for

Georgian Bay using climate records from Tobermory and Midland, Ontario from Environment

Canada. The earliest records of air temperature and precipitation data were from Tobermory

(1918-1998). Both air temperature (1974-1998) and precipitation (1948-1998) records for

Midland were incomplete (Cook, personal comm., 1998). Spreadsheet software was used to compute the potential evaporation and the water balance for each year (Dexter & Avery, 1991).

Georgian Bay receives input from precipitation, river mnoff, the North Channel and Lake

Huron. The above equation simplifies the natural hydrologic cycle by eliminating factors

associated with water storage in the soil and groundwater. In a lacustrine environment the .• . .

Equation 4: R + P + G - E + Qn + Qd + Qs = ALA / At

Where: R = runoff Qn = net outflow from Georgian Bay to North Channel (negative: flow to E = evaporation North Channel) (estimated volume for 80 yrs. is 3.12 x 10 m') P = precipitation Qd = deep net exchange (depth > 5m) at the Main Channel (negative: G = groundwater inflow flow to Lake Huron) A L = lake level change Qs = surface net exchange (depth <5m) at the Main Channel (negative:

A t = time flow to main Lake Huron) A = water surface area of Qd + Qs = total net exchange at the Main Channel (estimated volume for Georgian Bay 80 yrs. is 1.81 x lO'^m')

precipitation rate is defined as water falling on land or the lake surface. At present it is estimated, that about half of the water of the lake surface from lakes Huron and Michigan is lost

through evaporation (Appendix A, Table 1) (Hough, 1962; Phillips & McCulloch, 1972; Eraser

et al., 1990). The evaporation rate is proportional to the difference in temperature between air and water (Webb, 1972). Runoff allows the drainage of water from the land that may enter the lake basin or be intercepted by vegetation through evapotranspiration back to the atmosphere. In this case, we have ignored evapotranspiration and assumed total runoff enters the Georgian Bay basin. For the Georgian Bay basin seven gauging stations and three ungauged areas were used in estimating the amount of runoff entering the bay over seventy years covering 70% of the area

(Table 6) (Appendix C, Table 13) (Schertzer et al., 1979). Most runoff occurs during the spring

and fall months (April, May and Dec.) (Table 7).

Table 6: Gauge stations for Georgian Bay drainage basin. Ungauged stations are the southern part of Killamey (A), eastern part of Manitoulin Island (B) and the Georgian Bay shoreline in the Bruce Peninsula (C) (Schertzer et al., 1979). Gauge Station i: M'

I/O: 76 Table 7: Mean monthly runoff from the Georgian Bay drainage basin. Measured in cubic meters per second.

(French River (1), Magnetawan River (2), Muskoica River (3), Nottawasaga River (4), Severn River (5), Sydenham River (6), Wanapitei River (7) and ungauged stations (8)) (Schertzer et a!., 1979; Environment Canada, 1990). iJ-i'i :!1;.'"f i iJ > t

I 1 77

°- I ' I I ^

8 §

1=1 XJ^^

«>

3 o c

in oCO in

c CS

O i o ; eS SI

July paleo-temperature of 17.97°C (Table 8). The individual average January paleo- temperatures for palynological zones la and Ip are -1 5.48°C and -1 5.76°C, respectively. The average July paleo-temperatures for pollen zones la and Ip are 18.22°C and 17.72°C,

respectively (Table 8). The July temperatures in this zone range between 16.8 to 18.8°C. The

average annual paleo-precipitation rate for core LH95-035 is 681.7nim (Table 8).

9.1.1.2 Thecamoebians

Samples were barren of thecamoebians or consisted of very low abundances of Difflugia oblonga in core LH95-035 (Figure 31) (Appendix B, Table 2 & Figure 1). For a statistical valid

thecamoebian analysis large amounts of sediment were required from older sediments (Table 9).

Table 8:Paleoclimate averages for the different palynological zones of cores LH95-035, 038 and 041 .>:•"' ; V^

'• ' ',J:, ' i-i-' i'f-

'4' V^;\- . 79

"§883

§ §

ii^ •;£* 80

coldest palynological zone (range between -13.2 to -16.2°C) (Table 8). The average paleo-

precipitation rates for core LH95-038 is 681 .94mm (Table 8).

9.1.2.2 Thecamoebians

The thecamoebian diagram for core LH95-038 is characterized by the presence of three species, Centropyxis constricta, Centropyxis aculeata and Difftugia oblonga (Figure 32)

(Appendix B, Table 1 & Figure 1). The upper section of the core (pollen zone 2b) is dominated by Difflugia oblonga whereas the deeper section of the core is abundant in Centropyxis constricta and Centropyxis aculeata assemblages, and Difflugia oblonga is noticeably absent

(pollen zone la) (Figure 32). The abundance of Difflugia oblonga correlates with the increasing trend in mean January paleo-temperature in pollen zone 2b. Thecamoebians in the lower section of core LH95-038 are absent or in low abundances because of the small sample of core

sediments available for this analysis (Table 9).

Table 9: Sediment volume and total thecamoebian counts for cores LH95-035, LH95-038 and LH95-041. LH95-035 (Manitoulin Island) ill .-.4'. 1. i'_ r'fsl 1." >} 81

°l

^A>jf^4B

*. o *• 82 9.1.3.2 Thecamoebians

The older sediments (pollen zone 2a) are dominated by Centropyxis constricta and

Centropyxis aculeata and low Difflugia oblonga (Figure 33) (Appendix B, Table 1 & Figure 1).

The younger sediments at (92cm) reflect similar distribution as the older sediments of the

palynological zone 2a (Figure 33). Between 300 to 100cm (2a/ 2b boundary) there is an increasing trend in mean January and July paleo-temperatures. Samples from the older sediments required a greater volume of sediment because thecamoebian assemblages were fewer

or either absent (Table 9). The transfer ftmction performed on this core produced a January temperature curve emphasizing colder temperatures than present in pollen zone 2a (Figure 33).

The January curve is reflective of the arboreal pollen, which tolerated the colder envirormients.

9.1.4 97802-020 (Severn Sound)

9.1.4.1 Transferfunctions t

The transfer functions illustrate a pattern between January and July paleo-temperatures

(Figure 34). There is a declining trend (pollen zone 3a) in mean January paleo-temperatures between depths 120cm and 40cm (Figure 34). Therefore the older sediments (pollen zone 3a) are representative of a warmer time period (Table 10). There was a corresponding decline in

annual paleo-precipitation during the same period (Table 1 0). The average for the January and

July paleo-temperature of core 97802-020 is -8.07°C and 19.48°C (Table 10). There is a rise in

January and July curves (1-1 0cm) (Figure 34).

9.1.4.2 Thecamoebians

The thecamoebian analysis on core 97802-020 ranged between depths 40 and 95cm

characterizing the 3d / 3a palynological boundary (Table 1 1) (Figure 34). This portion of core ;•.- \'. >,

...!h.

(':::f}' u'*^ 'r'f-. '4

;V'

- \

-_ ..»: J 84 97802-020 is dominated by Difflugia oblonga assemblages. The species Difflugia oblonga appears in the younger sediments of pollen zone 3d (Figiire 34) (Appendix B, Table 2 & Fig. 2).

Table 10:

s s

Si-I

I I

A,^,

""^ x

o I o 00

o en %- e >u C/5 \

5 I c i

(uio)qjdaa- S S S

(u»)ilt(laa]g -***»—'-« -n... 2

9. 1. 5. Thecamoebians

The thecamoebian analysis on core 97802-026 ranges between 36 to 93cm consisting of

pollen zone 2a and partially pollen zone 3a (Table 1 1) (Figure 35). The older sediments (pollen

zone 2a) show an abundance of Centropyxis constricta and Centropyxis aculeata thecamoebian

populations (Figure 35). In the younger sediments (pollen zone 2b) the population shifts to a

dominance of Z). oblonga "oblonga", D. oblonga "spinosa" and D. oblonga "tenius"

assemblages (Figure 35) (Appendix B, Table 3 & Figure 2).

9.1.5.3 Loss-on-ignition

Silicates dominate the composition of core 97802-026 (Figure 37b) (Appendix B, Table

5). Organics are present in small abundances, but consistent throughout the core length.

Approximately 90cm organic content rises and silicates slightly decrease in the bottom of the

core. The calibrated radiocarbon date was 9,890 +/- 200 yBP (BGS 2120) at a depth of 90-

100cm of core 97802-026 (Figure 35). A second radiocarbon date for core 97802-026 at a depth of 50-60cm was 9,900 +/- 115 yBP (BGS-2191) (Figure 37b).

Radiocarbon dates were obtained for core 97802-026 (Figure 35). The older calibrated

age of 9,900 +/- 1 15 yBP (BGS-2191) (depth 50cm, pollen zone 2b) was above the slightly

younger age of 9,890 +/- 200 yBP (BGS 2120) (depth 90cm, pollen zone 2a) of core 97802-026

(Figure 35). This indicates that the 2a/ 2b boundary is around 9,900 yBP instead of the proposed

7,500 yBP by the palynological data. The ages are separated by a hundred years suggesting that

the sedimentation rate was continuous during the Main Algonquin stage. However, another

potential source of error for the calibrated dates is related to reworking or smearing of the older

sediments into the younger material causing contamination. This appears the logical proposal

based on the paleoecological evidence, however the paleo-indicators appear to be unaffected by

the reworking and contamination.

87 Table 11: Sediment volume and total thecamoebian counts for cores 97802-020, 97802-026 and 97802-032. 97802-020 (Severn Sound) :.:-, . -I -u' -J V '...'. ' -.A

"•-! §

"l-=

0) . O "5 £

o da o

^^X'^, ^e. ^<9 /«.% u 89

LOI: 97802-020 LOI: 97802-026 (Severn Sound) (Severn Sound)

* iP\^ r cs> t t Q

C'dtlc 9,900 +/-1I5 yBP

% 60 Percentage

% 60 Scale Icm to 20cin Percenuge

Scale Icm to lOcm

LOI: 97802-032 (Severn Sound)

33 1 LOI: 97802-035 (Severn Sound)

o^^ c* t & 20- % 60 Percentage

Scale Icm to Scm

LOI: 97802-036 (Severn Sound)

% 60 Petcentage

Scale Icm to lOcm

Scale Icm to 10cm

Figure 37: Loss on ignition for the Severn Sound cores. I 3--* p-^ o r^2

< J)

bo»

T3C o CO

> CO in CO cp CVJo 00

CO c .55 o i El (0 I

91 an increasing trend for mean January and mean July paleo-temperatures at a depth between 30

cm and 60cm (pollen zone 2a) (Table 12). There is a corresponding increasing trend in the

annual paleo-precipitation (Table 12).

9. J. 7.2 Thecamoebians

The thecamoebian analysis was conducted between depths 10 and 72cm of core 97802-

035 consisting within palynological zones 3a, 2a and la (Figure 38). The older sediments of

pollen zone la have a sparse population of Difflugia oblonga changing to dominant assemblages

of Centropyxis constricta and Centropyxis aculeata (pollen zone 2a) (Figure 38) (Appendix B,

Table 3 & Figure 3). In the younger sediments (pollen zone 3a) Difflugia oblonga are again

dominant. Colder climates show a dominance in Centropyxis species, whereas warmer time periods have an abundant Difflugia assemblage population.

9.1.7.3 Loss-on-ignition

Silicates dominate the composition of core 97802-035 (Figure 37d) (Appendix B, Table

5). The organic content is a consistent abundance throughout the core length. Carbonates

remain low through the upper portion of the core (10-60cm), but begin to rise between 60 and

70cm (Figure 37d).

Table 12: Paleoclimate averages for the different palynological zones of cores 97802-035 and 97802-036.

92 9.1.8 97802-036 (Severn Sound)

9.1.8.1 Transferfunctions

Transfer functions indicate cold January paleo-temperatures in pollen zone 2b; the curve

rises in pollen zone 3a and then declines in pollen zone 4 (Figure 39). At 40cm to 20cm depth there is an increasing trend in the mean January and mean July paleo-temperatures and a

corresponding slight increase in the annual paleo-precipitation (Table 12).

9.1.8.2 Thecamoebians

The thecamoebian analysis conducted for core 97802-036 was between 10cm and 45cm

in the palynological zones 3a and 2b (Figure 39). The younger sediments in pollen zone 3 at

10cm are dominated by Difflugia oblonga assemblage. The older sediments in pollen zone 2b is

characterized by the dominance of Centropyxis constricta and Centropyxis aculeata species replacing Difflugia oblonga (Table 13) (Figure 39) (Appendix B, Table 4 & Figure 3).

Table 13: Sediment volume and total thecamoebian counts for cores 97802-035 and 97802-036. 97802-035 rmu. -hI >

c ON g g < «^ o

3 3I— c

•Its

(ujo) mdea T

isostu auefljo

94 localities. The 2a/ 2b boundary illustrates a trend of cold January paleo-temperatures and low annual paleo-precipitation. The lacustrine environments in Severn Sound observed this trend, however Shouldice Lake detects a higher annual paleo-precipitation than the other localities

(Tables 14 «& 15).

Table 14: Paleoclimate averages for Severn Sound (97802-020, 97802-026, 97802-032, 97802-035, 97802-036) cores were combined and compared to Porqui Pond and Wye Marsh.

95 surplus occurs during the fall and winter months because evapotranspiration rates decline

outside the growing season and water is stored as snow cover during winter months. Other

factors controlling the water level of Georgian Bay is the inflow/ outflow from Lake Huron and the present climate of the region. Constructing the water budget for Georgian Bay illustrates the

interaction between hydrological and climatical processes affecting the bay over the last century.

The modern water budget provides necessary information in reconstructing the Middle Holocene paleolimnological water balances.

Modern Georgian Bay Water Budget (1918-1998)

80 25 Surplus 70 20 E E. 60

111 M5» a. -D 50 c 8^ la 10 g 40 sf 1. 30 E « o •- 20 I >d> -5" i 10

-10 Jan Feb Mar Apr May Jun Jul Aug Sept Oct Nov Dec

Figure 40: Georgian Bay water budget from the average climate data of Tobermory and Midland, Ontario over an eighty year time period (Cook personal comm., 1998).

' 9.2.1 Historical Climate Recordsfor Georgian Bay

Historical climate records for Tobermory, Ontario in the Bruce Peninsula and Midland,

Ontario in Simcoe County were obtained from Environment Canada in 1 998 (Cook personal comm., 1998). The records consisted of detailed information on temperatures and precipitation values for the last century. However, there are some data gaps in the historical documents.

Climate records from Georgian Bay provided a fundamental basis in establishing a link to

paleoclimates. The Tobermory records extend from early 1900's to the present day provides a

climatic relationship to Huron lake levels (Figure 41) (Appendix C, Tables 4-6) (Cook, personal

comm., 1998). The Midland data set contained many incomplete records, but they did have a M.

'/• I -fiusi: • J ." :

i^ 'I'-^i :

>ii"'! ' '' i .; U >'^^l U : . x''^

iv'v'- vrtb ' -j^r; .u "J!.i a, >- ;/''[ "•i"^-5("^•5 rr

.' ' .*('';'»: :':v\~:^,' >] ^r. •'.,i /;r.;-. , '^''.(•ii 22 j^®*****^*^© (O 6ep) ejnjDjedujei eBojeAv

97 continuous record from 1974 to present day (Figure 42) (Appendix C, Tables 7-9). The main

assumption that climate affected lake levels was proposed by early work from Fraser et al.

(1990) done on Lake Michigan. Tobermory climate records were compared to low and high

water levels through the 1930's and 1970's, respectively (Figure 43) (Appendix C, Tables 1 1 &

12). The realization that temperature and precipitation rates in the last century have played a role

in fluctuating lake levels is important for paleoenvironmental reconstruction and future

predictions of water levels. }

The historical water level data from Lakes Huron and Michigan within the past century

illustrates fluctuating lake levels caused by the effects of precipitation and evaporation. During

cool / humid summers, cloud cover limits infiltration of solar radiation increasing precipitation

and gradually lake levels. The cooler temperatures may attribute to fluvial flooding and increase runoff into the lacustrine basin. Dry summers on the other hand have higher water losses

through evaporation, lowering the lake level. This climatic response is observed in the lake level

history of lakes Huron and Michigan (Fraser et al., 1990). The total winter precipitation from

Tobermory and Midland were compared to Lake Huron water levels (Figure 44). Lake levels appear to respond to high accumulations of snowfall and spring runoff. In Lake Huron evaporation surpasses precipitation amounts between the months April and August. This is related to increased lake and atmospheric temperatures, incoming solar radiation and free ice cover. Climate historical records from Tobermory and Midland, Ontario were used to construct

a Georgian Bay water budget to illustrate the relationships of climate and hydrology. For this

study, we are estimating that the southern climate of Georgian Bay represents the entire basin.

This assumption may not be valid and could be verified by incorporating additional climate

stations into the Georgian Bay water budget. u.'

' '." fU.- .'">-J' - -.:i.'

ti. .• )<:<'/ .) 98

177 RIO -flOO S. 176.5 1 80 176 § 60 I ^ 175.5 1 20

175 J f J'^ 9\

176.6 S 1 3

177.2.

40-

176.9

176.5

ON9^ ^ 9t^Ch^O\ OiChd\

I I

176.1

Figure 43: Climate data compared to the Georgian Bay water levels. A - Tobermory climate data compared to 1929-1939 lowstand, B - Tobermory winter precipitation vs. 1929-1939 water levels, C - Tobermory climate data compared to 1969-1979 highstand, D - Tobermory winter precipitation vs. 1969-1979 water levels and E - Midland winter precipitation vs. 1969-1979 water levels. r

Tobermory Jan-Mar Precipitation vs. Lake level

S 300 -^^^1 c o a

•G 200

100

100 9.2.2 Problems with the Georgian Bay Water Budget

The major problem in constructing a realistic water budget for Georgian Bay over

seventy years are the various parameters or interactions of water in the landscape and the large

size of this drainage basin. There are numerous hydrological pathways and interactions to

consider for Georgian Bay. For example, inflow and outflow rates are generally distinguished by the amount of water entering or exiting the lake basin. These influx rates are more complex in

Georgian Bay due to the interaction between the bay and the Main and North Channels. On an annual basis, three times as much water flows to the North Channel from Georgian Bay than

Georgian Bay receives from North Charmel (Schertzer et al., 1979). Therefore for 9 or 10

months water flows out of the bay and for 2 or 3 months water flows into the bay (1 ,236 m /s).

At the Main Charmel on an annual basis, the net exchange is towards Lake Huron (7,170 m Is)

(Schertzer et al., 1979). A major drawback is estimating the flow between Georgian Bay, Main

Channel and North Charmel over a long time period. For this study, the inflow/ outflow ratio at these locations are based on the 1974 study due to the lack of historical discharge information

(Schertzer et al., 1979). Taking the discharge values at Main Channel and North Channel, volumes were estimated over a eighty year period (1.81 x lO'^'m^ and 3.12 x 10 m , respectively) (Equation 4) (Schertzer et al., 1979). Additional factors affecting Georgian Bay water levels are ice cover, wind and winter precipitation. The amount of winter precipitation reflects a recharge period from the spring runoff to the lake basin. Ice cover depends on the heat

storage of the lake, lake temperature, bathymetry and wind conditions (Webb, 1972). Freezing

occurs during calm cold periods with little vertical mixing allowing ice cover to persist through

the winter months depending on wind conditions (Webb, 1 972). Other significant factors in the

Georgian Bay basin are the geological and vegetation regions of the bay. Georgian Bay lies

within several different forest regions resulting in different rates of evapotranspiration and is

characterized by two dominant geological regions, the Ordovician sedimentary rocks and glacial '; ' :.. i

Oi . t' -;!..' ^; ' '.

'>i ,U!'K.. . ..;.!., !ii;

ft •)'-

^>:'. fh'

'*VM'

H -,,:

.»!,-) 'M ':!,•':.,; r ; '?

^ .;- iV'U'O!''' -M

"iisl, ^"i :k-:.' Kt-f" ';;• -ffn 101 deposits in the south, and in the north igneous and metamorphic rocks of the Canadian Shield.

A third of the drainage basins flow over the Canadian Shield contributing little sediment to the bay, in respect to the southern drainage basins that flow through various glacial landforms and drift thickness' contributing unconsolidated sediments and pollutants into Georgian Bay. The river basins receive water through precipitation, runoff and groundwater. Therefore, each drainage basin should be regarded as an individual system (Ward & Robinson, 1992). For a precise water budget, a hydrological model should be done to each individual drainage basin and on a yearly basis in order to evaluate the entire catchment area of Georgian Bay. High evaporation rates and low winter precipitation are commonly associated with low water levels

(Figure 40). The historical water level records of Lakes Huron and Michigan has provided general information towards fluctuating lake levels through several decades, but Georgian Bay is shallower in comparison to the upper Great Lakes. The bathymetry of Lakes Huron and

Michigan is important because a decrease of a meter in lake level may not be as drastic in deeper lake environments in comparison to Georgian Bay's eastern shoreline where a small change in lower water levels, would increase beach environment exposure. Direct lake level measurements

for Georgian Bay is required to comprehend the effects of climate change.

9.3 Willoughby Bog

9.3.1 WB4 (Channel Environment)

9.3.1.1 Palynology

The pollen diagram for core WB4 is characterized by Pinus, Gramineae, Quercus, Acer,

Ulmus, Fagus, Nyssa, Alnus, Sambucus and Cornus (Figure 45) (Appendix D, Tables 3-5). Core

WB4 can be subdivided into six palynological zones that include 4 (6- 10cm), 3d (10- 160cm), 3c

(160-240cm), 3b (240-265cm), 3a (265-320cm) and 2b (320-370cm) (Figure 45). Pollen zone 4

(150 yBP to present) is characterized by the Ambrosia rise (Figure 45). Pollen zone 3 is

102

1.^

103 represented by the presence of Acer, Fagus and Tsuga. Pollen zones 3b and 3c identify a rise in Sambucus and Cornus, and a decline in Nyssa, Carya, Alnus and Ulmus (Figure 45). Pollen zone 2b (10000 yBP) is characterized by the dominance oiPinus and Gramineae (Figure 45).

Transfer functions from the palynological data produced a fluctuating curve throughout core WB4 (Figure 45). Increasing trend in mean January paleo-temperatures from pollen zone 2 to pollen zone 3d (Figure 45). Followed by colder mean paleo-temperatures in modem sediments. The overall average January and July paleo-temperature for core WB4 is -7.26°C and

20.53°C, respectively (Table 16). The annual paleo-precipitation follows a similar trend to the mean paleo-temperatures (Figure 45). The average paleo-precipitation is 780.69mm for core

WB4 (Table 16).

9.3.1.2 Thecamoebians

The thecamoebian analysis of core WB4 was between depths 248cm and 370cm (Figure

45). This corresponds between pollen zones 2 and pollen zone 3b (Figure 45). Colder climate conditions prevalent in pollen zone 2 show a dominance of Centropyxis aculeata assemblages

and near absence of Difflugia oblonga (Figure 45) (Appendix C, Table 1 & Figure 1). They also

show their dominance throughout the core (Table 1 7).

9.3.1.3 Loss-on-ignition

Core site WB4 was 360m from the forest edge at a total length of 3.74m. Organics (i.e. peat) are abundant the upper portion (12- 120cm) of core WB4 with low percentages in silicates and carbonates (Figure 46a) (Appendix D, Table 9). From 120cm to 300cm core WB4 consists of a mixture of organics and silicates (Figure 46a). The remainder of core WB4 is abundant in <: hU • 1

.:'/'» I I

.' ••' M.,:i 1 .' I Ti'i

^ ,' nr l.hUi-''^': '.'I ' '.A' 104

Willoughby Bog Core 4 (WB4) LOI and Stratigraphy

Description J' .o" .^* cf litter layer (dark humic) ^ ^ dark peat horizon (wood at 20cm)

peat (liglit brown), fibrous

peat (light brown), fibrous

l@J wood (11 5-13 1cm) I^^H peat / clay (light brown) ^^H wood present at lS6cm and 181cm

peat / clay (dark brown)

gray clay horizon

peat / clay sediment (dark brown peat) wood present at 241cm and 286cm

dark gray clay with less peat I O^date 9,015+/- 100 yBP dark gray clay I T |— I I — — [ % 60 60 100cm Percentage Scale 1cm to 50cm

Willoughby Bog Core 6 (WB6) LOI and Stratigraphy

Description

dark brown peat (humic)

fibrous peat

dark brown peat (fibrous)

very fibrous (roots) wood (115 to 133cm)

peat/ clay sediment

wood present at 241cm and 269cm

O^date 9,400+/- 100 yBP gray clay

0/, 60 60 100cm Percentage Scale 1cm to 50cm

Figure 46: Stratigraphy description and LOI for cores WB4 (a) and WB6 (b).

105 silicates (i.e. clay) with little amounts of organics. The calibrated radiocarbon date of gyttja in core WB4 at a depth of 345-357cm was 9,015 +/- 100 yBP (BGS-2210) (Figure 46a).

Table 16: Average paleoclimate values for the different palynological zones in cores WB4 and WB6.

106

t-^rrHi^ OQ

o P^-W a

^ O

O CD 60 OQ 2 > 0)

o -ao CD

Ji: O)

o u CQ l§

.2 . E (0 tC o >- '^"^ CO li-O b §^ c (0 •s ^ • - s

(U o 6 ° E 1^1 (0 I .2 o u s 0) * c « .2 a! o -^ • o ""^ o> 8 S CM « o iSo u is J= -a a

c -

'^ §

.s (ua) q)d>a =.2 ISBT3 'Oi

sC 107 855mm and the coldest January (-8.0°C) and coolest July (19.9°C) average paleo- temperatures (Table 16). The average January and July paleo-temperatures for core WB6 are

6.3°C and 20.73°C, respectively (Table 20).

9.3.2.2 Thecamoebians

The thecamoebian analysis was conducted between depths 230cm and 330cm representing the pollen zones 2 and 3b (Figure 47). There was one sample depth close to the surface at 79cm representing pollen zone 3d (Figure 47). The thecamoebian data reflects the abundance of Arcella artocrea in the peaty portion (79cm) of the core. This section of core WB6 is very humic and well saturated. Centropyxis constricta and Centropyxis aculeata are dominate

in the lower depths of the core (Figure 47) (Table 17) (Appendix D, Table 2 & Figure 1).

Table 17: Sediment volume and total thecamoebian counts for cores WB4 and WB6. WB4 ..r!(

, ': /\! '-fO-- -.UJ-ji'-'. ^;u•1;/^

>" ffr':.''i'X' if'v- M^^^ , ; ! i

-. •(! 'ii,;v • J'ar] . ('ij! 1 : {;;... ''-j^i) /; t' 108 10.0 Discussion

Paleoenvironmental reconstruction of the Great Lakes during the Holocene period is complicated. Various lake levels have occupied the Great Lakes basin as a result of isostatic rebound and climate change. Around 7,500 yBP, a lowstand generally referred as the Late Lake

Hough phase produced a closed lake within the Georgian Bay basin (Blasco, 1994). The most

convincing evidence for this lowstand is the seismic record identifying an apparent basin-wide

erosional unconformity (Figure 15) (Rea et al., 1994; Dobson et al, 1994; Moore et al., 1994;

Blasco et al., 1997; McCarthy et al., 1997). The age of the sediments were estimated palynologically around 7,500 yBP, near the pollen zone 2a/ 2b boundary, prior to the main

Hypsithermal Period (Figure 48) (pollen zones 3a-3c; Dubas, 1997; McCarthy et al., 1997). An additional characteristic of the 2a/ 2b boundary is relatively high abundance of Cyperaceae

(sedges) and other aquatic pollen in cores LH95-041 and 97802-026 (Dubas, 1997; McCarthy et

al., 1997). This palynological signal records a shoreline or marshy environment near Tobermory and Severn Sound are consistent with low water levels in the bay (McAndrews, 1994; McCarthy

et al., 1997; Tiffin, 1999). A radiocarbon date of 7,940 +/- 1 10 yBP in floodplain sediments prior to increasing water levels and marl lake environment Wye Marsh (depth of 360-364cm) also records low water levels in Georgian Bay (Figure 28) (Chittenden, 1990).

Recognition of fluvial channels in the area of Fathom Five National Marine Park provided additional evidence of waters draining through the channels from Late Lake Stanley

(Huron) to Late Lake Hough (Georgian Bay) around 7,500 yBP (Blasco et al., 1997; Blasco et

al., 1999). The distinction of these fluvial channels indicates water levels were lower than present day in Lake Huron and Georgian Bay (Blasco, 1994; Blasco et al, 1997; Blasco et al.,

1999). The paleoenvironmental reconstruction from this evidence suggests Lake Huron waters drained through these fluvial channels over the Niagara Escarpment producing knick points

(waterfalls) into Georgian Bay via the North Bay outlet around 8,000 yBP. - ti-. ':/; . .I,'! : , .'; ^^r^' ;i'i r-i''--

._.,• A^:)/- 1.

•t.<'v

"i- ,. " ..I

^Jl.(r;i ;! "li:)' '"'; -I'f

;f.:!:i'' )1 ';o i'i' . I

"j V i- •.M'jvl

:-; ,»

'J ' J .

'i ,

I Ice Front u — Kirkfield Phase

Main Algonquin a. v> Post Al gonquin Early Lake Stanely/Hough Early lyiattawa I Mid Lake (B Stanely/Hough 8 CM p CO Main Mattawa

C. •on tricuspis 90 \

o -

09 -

m I M

CD a^ 5>

in 110 The discharge of water between Lakes Stanley and Hough is unknown for this time

period. The net exchange of flow that currently exists between the Main and North Channels

and the bay did not occur during the Late Lake Hough/ Late Lake Stanley lowstand (7,500 yBP)

implying water from the upper Great Lakes drained through the bay towards the North Bay

outlet (Figure 49). The North Bay outlet controlled a consistent flow from Lake Agassiz that

drained via the Ottawa River through the lowstand (Figure 48). Runoff from northern rivers in

Georgian Bay may have been diverted from the bay due to the isostatic depressed land and drained directly to North Bay resulting in low water levels in the bay. Another general trend

associated with lowstands is fluvial deposition and aggrading streams and/or delta formation

depending on lithology (McCarthy & McAndrews, 1988). This trend may have occurred in the

Severn Sound area (Figure 49). The problem with identifying the 7,500 yBP lowstand is in misinterpreting landforms and the stratigraphic record from the erosive processes of rising lake

levels. Lower postglacial lake phases formed beach ridges that were later drowned by present day lake levels due to isostatic rebound (Larsen, 1987). However, the short-lived 7,500 yBP

lowstand probably did not form distinctive lacustrine shoreline features due to the bedrock

geology of the area and the subsequent Nipissing transgression.

This 7,500 yBP lowstand was originally attributed to isostatic rebound and the diversion

of meltwater to controlling outlets, but no potential outlet has been found. This lowstand

postdates the fransfer of drainage to the Ottawa River via the North Bay outlet that resulted in the

Late Lake Hough lowstand, and no geological explanation has been found for this lowstand

which is lower than the elevation of the North Bay outlet (Blasco et al., 1999). Waters from

Georgian Bay continued to drain through the North Bay outlet until 5,000 yBP due to isostatic

rebound, uplifting North Bay and diverting flow through the lower Port Huron outlet in southern

Lake Huron (Figure 48) (Lewis et al., 1994). Lake Erie also experienced low water levels at this

time, thus a climatic explanation appears plausible. .}< '..ur^ !4 7V 0-

\U r-^ .It ly'ib^/ co' ' :'<'>'i :>%';>! ..^..

''r).-'i., Ill

Legend

H Difflugia oblonga Low Winter Precipitation Q Centropyxis constricta

^fl Centmpyxis aculeata

Mattawa/ Ottawa Rivers

Legend

(j Difflugia oblonga High Winter Arctic Centropyxis constricta ^ JPrecipitation Air Mass (cold/ dry) ^g Centropyxis aculeata

Figure 49: Paleo- water budget reconstruction of Georgian Bay around 7,500 yBP (a) and present water budget conditions (b) (modified from Blasco, 1994; Edwards et al., 1996 and

McCarthy et a!., 1997). Illustrating Georgian Bay sediments and diverse thecamoebian assemblages. Lake highstands would recycle and redeposit the sediment in bathymetric lows. Lowstands would expose the sediments to erosional agents producing unconformities or an incomplete paleorecord.

112 Palynological and quantitative paleoclimate evidence from Georgian Bay suggests a dry/ cold climate similar to modem Minneapolis, MN existed around 7,500 yBP (Dubas, 1997;

McCarthy et al., 1997; Tiffin, 1999; Sarvis et al., 1999a). Data from Willoughby Bog in Niagara

Falls suggest that relatively dry conditions affected all of southern Ontario at this time, not just the Georgian Bay basin. The 2a/ 2b pollen zone boundary (-7,500 yBP) recovered in the

Tobermory core LH95-041 and Severn Sound core 97802-026 (Figure 49) records a distinct climate change. This boundary is characterized by a change in the palynological record from

Pinus banksiana/ resinosa (pollen zone 2a) to Pinus strobus (pollen zone 2b) dominated assemblages (McAndrews, 1994). Both Pinus species represent different modem climatic zones-

Pinus banksiana/ resinosa characterizes the Boreal Forest (cold/ dry), as is found in modem

Minneapolis Minnesota, and Pinus strobus is abundant in the Mixed Deciduous Forest (moist/ warm), more similar to conditions found in the Georgian Bay basin today.

Earlier palynological information from Georgian Bay interprets a different paleoclimate and paleolimnology environment. McAtee (1977) interprets a warm/ moist climate with rising lake levels around 7,500 yBP in Georgian Bay. Recent paleoecological evidence suggests that

Georgian Bay was lowered by climatic mechanisms (Dubas, 1997; McCarthy et al., 1997) from the intmsion of the dry Pacific air mass (Figure 18) (Edwards et al., 1996). The paleoclimatic trend from the Georgian Bay pollen diagrams illustrates cold January paleo-temperatures in the

older pollen zones (1 and 2a) with low amounts of paleo-precipitation (Tables 14 & 15). Paleo- precipitation values from the transfer ftmctions must be viewed with caution because various

factors (i.e. soil moisture and vegetation) control the precipitation response.

Palynological records from smaller lacustrine environments surrounding Georgian Bay

(Porqui Pond, Wye Marsh and Shouldice Lake) record a dry/ cold climate in palynological zone

2a (Figures 28 - 30). From pollen zone 2b to today, in contrast, the January paleo-temperatures are close to modem values (McAndrews personal comm., 2000). Annual paleo-precipitation rn

^''1 .:

:.:ii{iOi.

.'(:;?;•.!;

-y:n- ii\ '

'.'.' • -a-. V ! ':'^iytr..irr,^\ii-.

'-•! - ;«'/

:'*"'* i::.<'-7» 1^ I'

"'' .:^-^:: ''' lir,: •n . . -ji„.i ":;i ,» t>. .J U C ''-T-

• ',! '. '.'.,; ''"j* '.: l"V-' L- .f;< s ; 113 values have been close to modem values from palynological zones 3a to present (Tables 14 &

15). ^ -

Paleolimnological evidence from thecamoebians were used to interpret the paleohydrology of the Georgian Bay basin (Figure 2) (McCarthy et al., 1997). Thecamoebians were studied in the Georgian Bay sediments because of their sensitivity to environmental change.

Thecamoebian assemblages changed in response to climate and water level changes in the

Georgian Bay basin (Figure 48). The older sediments (pollen zone 1) recovered primarily from the Tobermory region (cores LH95-041, LH95-038 and LH95-035) contained low thecamoebian abundances, recording the harsh environment (cold water and air temperatures lowering

thecamoebian productivity) and/or the high sediment influx from the ice margin (Table 9). High sediment influx affects the food web by lowering primary productivity (decreasing " photosynthesis), therefore less phytoplankton and less overall abundance of thecamoebians. The existence of this very sparse Difflugia o6/o«ga-dominated fauna during the late glacial/early postglacial is consistent with the very low thecamoebian Difflugia o6/o«ga-dominated

populations in modem Arctic lakes (Collins et al., 1990). The replacement of this fauna by a sparse Centropyxis aculeata- and Centropyxis con5/r/cto-dominated fauna about 9,000 years ago

(pollen zone 2a) is not consistent with the wanning climatic trend (McCarthy et al., 1995), but appears instead to record an increase in dissolved ions in Georgian Bay. High dissolved ions could represent high groundwater discharges and/ or high evaporative rates. Sparse thecamoebian faunas dominated by Centropyxis aculeata characterize brackish water environments such as estuaries and salt marshes around the world (Medioli & Scott, 1988). Cold water from glacial discharges would keep thermal environmental conditions cold, thus in favor of Centropyxis aculeata assemblages. When glacial discharges diminishes and precipitation is the dominate water source Difflugia oblonga emerges as the controlling species. :' ^^'' -/.fU;v'>v! . , )i. !:> Vdff'-

i|:> 'r? '/ci r->i-5>.'

' ''.*h,i

<• :;!

:;.<- -iL-fa I

/r ^!i-? ((,.< c'-^/jsj);^^ ';'• '-"^i: • '>',-

•»j '<; 'f. .-. «,. , ^, n ..-a' , ':>:i'. \>\vA..! 'V! c^<<>>< 114 Palynological zone 2a is represented in cores LH95-041, 97802-035, 97802-026 and

97802-036 and sediments are rich in Centropyxis aculeata and Centropyxis constricta. The presence of the Centropyxid species represents shallow water levels, together with an abundance

of aquatic pollen (McCarthy et al., 1997). The cold waters and highly evaporative conditions

recorded by the abundant Centropyxid species (Medioli & Scott, 1988; McCarthy et al., 1997)

are supported by the pollen record and by the paleoclimatic reconstruction of Edwards et al.

(1996). The Centropyxid species flourish through the Mattawa Lowstand due to the influx of meltwater and high dissolved solids characteristic of glacial runoff conditions from Lake

Agassiz. The transfer fiinctions from these cores illustrate the cold paleo-temperatures associated with the abundant Centropyxid population.

The Centropyxis constricta- and Centropyxis aculeata-dormmXQd populations were replaced by an abundant and diverse Dijflugia oblonga - dominated assemblage during pollen zone 2b. In sediments assigned to pollen zone 2b in cores LH95-038, LH95-041 and 97802-026, total thecamoebian populations increase dramatically, and thecamoebian assemblages are dominated by Difflugia oblonga and other Difflugiid species. This suggests a extraordinary environmental change in Georgian Bay. Difflugia oblonga represents a change towards warmer and freshwater conditions, as water levels gradually increased in the bay (7,500 to 7,000 yBP)

(McAndrews, 1994; Scott et al., 1999). Difflugia oblonga can also represent warmer water conditions and normal precipitation because there was no more glacial runoff during this time period. This fauna persisted throughout Georgian Bay until human landclearing for agriculture caused eutrophication of Severn Sound, allowing Cucurbitella tricuspis to become very abundant

(McCarthy and Tiffin, 1999; McLachlan, 1999). The increase in thecamoebian diversity and the establishment of the more abundant and diverse Difflugia oblonga- dominated faima approximately 7,000 years ago appears to have been related to sharp increase in mean January temperature recorded by the pollen record. >>iJ

tiU-rjiJ.' ."! '.'^•,» VN' >'!'i >il 'h'\'iy

;.!Uii:( i: <\w\A^^ iy.

•J

bv.. 115 In some areas, e.g. in Severn Sound cores 97802-035 and 97802-036 (Figures 45 &

46), pollen zone 2b has been eroded from the paleorecord, and the contact between pollen zones

2a and 3a is clearly unconformable. High thecamoebian populations in sediments rich in Acer,

Ulmus and Tsuga (pollen zone 3a) in cores 97802-020, 97802-032, 97802-036, 97802-035 and

97802-026 are dominated by Difflugia oblonga and other Difflugiid species (Figures 34- 36, 38,

39) (McCarthy et al., 1997). The abundant and diverse thecamoebian assemblages characterize

the hypisthermal (McCarthy et al., 1 997, Scott et al., 1 999).

The historical Huron-Michigan water levels and climate records from Tobermory and

Midland suggest a paleoclimatic/paleohydrological explanation for the 7,500 yBP lowstand: a climate-induced change in the water budget (Figures 41 & 42). Fraser et al. (1990) provided information between the relationships of climate and Michigan lake levels in the late Holocene, showing that climate has influenced the lake levels in the Great Lakes. Climate-induced low lake levels are characterized by less runoff and increased evaporation rates from the lake surface

(Larsen, 1985). The historical climatic records from Tobermory and Midland provide evidence of the similar trends between lake levels and precipitation, specifically winter precipitation

(Figure 44). Lake levels generally increase after a duration of high winter snowfall for

consecutive seasons, whereas they fall following several consecutive dry winters, as they have during the late 1990's. Sufficient annual supply of precipitation (approximately 900mm) is important for a healthy modem Georgian Bay water budget.

Quantitative estimates, made by running transfer functions on the palynological data, suggest that Minneapolis, Minnesota is a good modem analogue for Georgian Bay during pollen zone 2a (10,000 to 7,500 yBP), with abundant Pinus banksiana/ resinosa and Quercus characterizing dry/ cold winters and hot/dry summers (Appendix A, Table 6). The present day mean annual precipitation in Minneapolis, Minnesota is approximately 750 mm compared to

Tobermory, that with an approximate annual rainfall of 900 mm (Appendix A, Table 6) •:i

»:•.! ';>!-. f-i-j-,'

L'.:t

^j ' r'

i ji'i! U

> V

' . .1 , ,

•' :.<<,'i'". .Jtov-'-.-xf •\q'. 116 (Steinhauser, 1979). Minneapolis, MN experiences a surplus period for seven months between November and May. This is different from modem Georgian Bay, which has a surplus from September to May meaning that Minneapolis, Minnesota is significantly drier than the

Georgian Bay basin. Cold, dry winters and warmer, highly evaporative summers produce a small recharge period and low total winter precipitation (Figure 50) (Steinhauser, 1979). The emphasis on the winter months is important for the water budget because total snow accumulation reflects the amount of runoff during the spring melt-out causing potential low water levels in Georgian

Bay (Figure 50) (Appendix C, Table 2).

Water Budget: Minneapolis, IVIinnesota

to 0) " Q- re -J u. 5 < 5 Months

Figure 50: Water budget of Minneapolis MN. Mean temperature and precipitation values (Steinhauser, 1979).

The main assumption of this climate-induced lowstand hypothesis is that the

Minneapolis, Minnesota-like climate persisted over several hundred years. Bengtsson and Malm

(1997) constructed water budgets for open and closed basins for Sweden lakes highlighting the importance of precipitation and evaporation, especially in closed lake basins. High evaporation rates decrease the lake surface area and lower water levels in the basin.

Applying the lower snowfall characteristic of modem Minneapolis, Minnesota to the size of Georgian Bay over a few hundred years would decrease the surface area of the bay (Appendix

C, Figure 1). Modem Georgian Bay and Minneapolis, MN climates were applied to the specified hi 117 time interval and Georgian Bay drainage basin area in order to compare potential evaporation and precipitation rates (Appendix C, Table 2). The evaporation/ precipitation ratio forecasts the condition of the water levels in a lake basin. The evaporation/ precipitation ratio for Minneapolis,

Minnesota is relatively equal illustrating highly evaporative summers and low precipitation rates compared to the present Georgian Bay values over the same time period (Appendix C, Table 2)

(Figure 50).

Other factors to be considered for the Late Lake Hough lowstand are ice cover, inflow and outflow and runoff. During a warm/ dry paleoclimate, ice cover would be minimum allowing moderate evaporation rates to extend into the winter months. The amount of inflow/

outflow from the North Channel and Main Channel is theoretical, but the main assumption is a consistent discharge rate that created the Late Lake Hough closed basin in a dry climate. Runoff would be a minimal factor under these paleoenvironmental conditions (Bengtsson & Malm,

1997). All these factors would have a negative impact on the water levels in the bay, and suggest

that it is plausible that the 7,500 yBP lowstand was caused by climatic conditions.

Lake Erie water levels responded to changes in the upper Great Lakes only when drainage was through the Port Huron outlet. Otherwise, Lake Erie water levels were independent of those in Georgian Bay. The paleoenvironmental reconstruction of Willoughby Bog in the

Niagara Peninsula provided an age of peat bog initiation (paludification) around 7,400 yBP, recording the onset of more moist conditions. Pollen-derived paleoclimatic reconstructions from Willoughby Bog in the Niagara Peninsula also recorded drier conditions until around 7,400

years ago (Sarvis, 1997; Pengelly et al., 1997; Sarvis et al., in prep), suggesting that cold/ dry conditions existed throughout southern Ontario around 7,500 yBP.

An older calibrated radiocarbon age of 9,400 +/- 100 yBP (BGS 22 11) was obtained from a core taken from a deeper channel in Willoughby Bog, suggesting that the channel drained

into Lake Tonawanda prior to the Main Algonquin stage (D'Agostino, 1958). Erie waters were xu

•?Ltor, 1.1.' ; 1 [;,:.

'..) -7 -;»/!"'_» J)>'. 1 X)it,}?.f. .A'i'}\ 118 diverted producing tlie shallow tributary that connected through modem Usshers Creek to

Lake Tonawanda (9,400 yBP) (Figure 3). These radiocarbon dates indicate two different stages of peat accumulation in Willoughby Bog, following Early Lake Hough (9,400 yBP) and Late

Lake Hough (7,400 yBP). A later stage of Lake Tonawanda occurred during the Nipissing Rise resulting in a shallow lake basin from the affects of rising Erie waters and sediment deposition.

The thecamoebian assemblages rich in Centropyxis constricta and Centropyxis aculeata are present in the grey massive glacial clay of Willoughby before the transgression to bog environment, recording cold, possibly slightly "saline" conditions. The saline conditions may represent glacial discharge or groundwater flow and high evaporation rates. As in Georgian Bay, this assemblage rich in Centropyxid species in the Georgian Bay sediments characterizes pollen

zone 2a (10,000 to 7, 500 yBP).

The Willoughby Bog channel was used a second time as Lake Tonawanda re-emerged

prior to 7,000 yBP (Pengelly, 1990; Pengelly et al., 1997, Sarvis, 1997; Sarvis et al., 1997). As

Lake Tonawanda was slowly tilted towards the Niagara Region, eastward flowing river gradients were reduced, drainage towards the Niagara River slowed, and Lake Tonawanda flooded slowly backing up the Welland River, Lyons Creek and Usshers Creek (Tinkler, 1994). The palynological record reveals the slow shift of a river to a river lake to a peat bog (Pengelly, 1990;

Tinkler, 1994; Sarvis, 1997; Sarvis et al., in prep). The falling water levels record a dry climate, while the onset of peat accumulation records the onset of more moist/ less evaporative conditions. The apparently dry climate may instead represent the reduced size of the Willoughby

watershed (i.e. very little drainage towards Willoughby Bog).

A second inflow from Lake Agassiz prior to 7,000 yBP was proposed to have formed

Lakes Wainfleet and Tonawanda as rising waters entered through the abandoned channels in the

Lowbanks area (Tinkler et al., 1992; Pengelly et al., 1997). The existence of low lake levels in both the upper Great Lakes and in the Lake Erie basin around 7,500 yBP, however, is not l"! !

.'tji, " r

• i"i ,. r J , , I "!;..(•

•i. ?:.:- 119 consistent with the rise in water levels in the Erie basin being a response to drainage transfer

from the Huron basin.

11.0 Conclusions

This study investigated the 7,500 yBP lowstand of Georgian Bay. Also known as the

Late Lake Hough phase, this lowstand was identified from seismic profiles and could not be

explained through geological processes (i.e. isostatic rebound and drainage outlets) (Blasco,

1994). Relatively abundant sedges and aquatic pollen found at the 2a/ 2b boundary from near

Tobermory and in Severn Sound (McCarthy et al., 1997), and the floodplain environment at Wye

Marsh (Chittenden, 1990) support the existence of lower water levels in the Georgian Bay basin

around 7,500 yBP. The North Bay outlet, whose elevation is higher than this lowstand, is thought to have controlled the drainage for the upper Great Lakes during this time period, and outlets lower than North Bay have not been identified. This thesis suggests an alternative paleoclimatic hypothesis for the 7,500 yBP lowstand, using paleoecological evidence to constrain a paleo- water budget for the Georgian Bay basin for the Late Lake Hough phase of Georgian Bay.

Pollen analysis allowed for reconstruction of the paleoclimate for Georgian Bay. The

7,500 yBP lowstand of Georgian Bay corresponds to the 2a/2b palynological boundary. This boundary is characterized by the changeover from Pinus banksiana/ resinosa to Pinus strobus dominated assemblages (McAndrews, 1994), recording the replacement of a paleoenvironment comparable to modem Minneapolis, Minnesota by one more similar to modem conditions around Georgian Bay. The main difference between the two modem analogues is the substantially colder and drier winters (and much lower snow accumulation), and the somewhat lower total annual precipitation in Minneapolis, Minnesota compared to meteorological stations in Tobermory and Midland. 0!

'}d ':?(! b!i!

iB;:;n :vi

{' .1M

;'-• U^ Linoitv'TKoliif! i;'-..rn^;'>:'t'-;tn ->,b

.u; r jU, J-,..rr

•'.:.;. i I,',' >"i (I '''::''qi" lij L'i-'n 120 Pollen diagrams from Porqui Pond and Wye Marsh near Midland, and Shouldice

Lake, near Tobermory, and from the Willoughby Bog in the Niagara Peninsula record a similar

cold/ dry period around 7,500 yBP, suggesting that drier conditions existed throughout southern

Ontario at this time. This is consistent with the hypothesis that the dry Pacific air mass entered

the Great Lakes basin, lowering the precipitation rates at this time (Edwards et al., 1996).

Thecamoebians were analyzed and compared with the pollen data for the Georgian Bay.

These protists record conditions within the lake, rather than the upland vegetation as the pollen record does, but paleolimnological interpretations based on thecamoebian assemblages are

limited by the relatively small amount of paleoecological research on these organisms. More

research is needed with modem thecamoebian distribution in both climate and habitat zones in

both hemispheres to use in interpreting "pa/eo" conditions (i.e. dry conditions/ hot climate, wet

conditions/ cold climate). The thecamoebian succession in Georgian Bay records can be

subdivided into four zones; the first zone consists of similar conditions found in modem Arctic

lakes containing a sparse Difflugia oblonga assemblage population (10,500 to 9,000 yBP).

Pollen zone 2a (9,000 to 7,500 yBP) sediments are characterized by a Centropyxis-dommdlQd assemblage, identifying brackish conditions and increased dissolved solids during the Mattawa

lowstand. This is replaced just above the pollen zone 2a/2b boundary by an abundant and diverse

Difflugia o6/o«ga-dominated assemblage, signifying a dramatic change in environmental conditions in the bay. This assemblage persisted from -7,500 to 750 yBP, until the sedimentation

and nutrient influx increased as a result of European, and possibly Native American, agricultural

activity especially in the Severn Sound area (McLachlan, 1999; Sarvis et al., 1999b). A

thecamoebian assemblage dominated by Cucurbitella tricuspis in these sediments characterizes

the eutrophication of Sevem Sound

The constmction of a modem water budget for Georgian Bay was important in order to

reconstmct a paleo-water budget during the Late Lake Hough lowstand. The historical climate i^ Mtl^'^

li ' ' /.'-'! Hi. J l! - iifiiU V:.s):;a 121 records from Tobermory and Midland were incomplete, but the relationship between

Georgian Bay's climate and water levels over a seventy-year period is clear. The more important factors that control Georgian Bay lake levels are precipitation, evaporation and inflow/ outflow rate at the Main and North Channels. I examined the precipitation more closely and discovered that winter precipitation had a correlation with lake level changes. Consecutive moist winters would result in an increase in lake levels, and vice versa for dry winter periods. The precipitation/ evaporation ratio of the paleo-water budget for Georgian Bay appears to be another controlling factor on lake levels. Another important factor related to Georgian Bay's water budget is the outflow/ inflow ratio from Lake Huron and North Channel. Limited flow data was achieved for the Main and Little Current chaimels resulting in an estimation for a seventy-year period. However, it is important to note that waters from Georgian Bay flow towards Lake

Huron for most of the year. There are many variables that control the movement of water through the Georgian Bay drainage basin. For example, rivers, inland lakes and man-made dams that transports and stores water over a specified time period confrolling the water levels of the bay. Other factors include the geology, vegetation and soil type in the drainage basin. For this study it was difficult to consider all of these factors for the size of the Georgian Bay catchment area. Future research should be aimed at smaller lake and wetland basins around Georgian Bay.

Additional climate stations should be incorporated into the Georgian Bay water budget.

The Tobermory and Midland climate data reflects southern Georgian Bay conditions and was

used to estimate the entire water budget for the bay. Northern climate stations (i.e. Sudbury) should be analyzed and compared to lake level changes from Georgian Bay. Water budgets should be calculated on a yearly basis for individual stations in order to comprehend the movement of water in a smaller drainage basin and/ or Georgian Bay. This study illustrates the importance of hydrology and climate research in the basin due to the drastic lowering of the water level in recent years. •'; .I'j.-.ii.:. .X'- . tjKjll; Ijr. -, /i;

sri i

..J^.- U'. ;...-i' j;,,. > l^ .iflCV'-V' ' .iO&hi

•• '!•''< •*.' J'' ' f>'ir ^-;;'-!Si .^, . l,-j-iur;yi:...'i iii-R i> 'V

3i>*«M *,f!;. !TW;6d jyEfiit,.' -lir.ait' :? fl' U>:.; v (o

i.i.

'V K!S>-V 122 The palynological and thecamoebian analyses at the pollen zone 2a/2b boundary

suggest that brackish and highly evaporative conditions were responsible for the unexplained

early Holocene lowstand in Georgian Bay. Climate thus appears to have had a substantial impact

on prehistoric water levels in the Georgian Bay basin. I conclude that low winter precipitation

and high evaporative rates over several centuries had a dramatic impact on Georgian Bay,

resulting in the Late Lake Hough lowstand.

The paleoecological research on Willoughby Bog has established the hydrological status

of the wetland, and a paleoclimatic reconstruction for Niagara Peninsula during the early

Holocene. The paleoclimatic record from Willoughby Bog reflects cold/ dry conditions during pollen zone 2a, similar to the conditions the Georgian Bay basin. The radiocarbon dates from

Willoughby Bog indicate low water levels in the Erie basin (Early Lake Hough) until peat

accumulation was initiated around 9,400 yBP. Prior to 9,400 yBP, Willoughby Bog was a

fluvial channel that drained into Lake Tonawanda (10,500 yBP) (D'Agostino, 1958). The

Willoughby Bog experienced a second process of peat accumulation around 7,400 yBP (Late

Lake Hough) (Pengelly, 1990, Pengelly et al., 1997; Sarvis, 1997) that can be related to

paludification following a dry/ cold paleoclimate in the Niagara Peninsula.

The paleoecological evidence from Georgian Bay and Willoughby Bog suggests that all

of southern Ontario was affected by a dry/ cool climate around 7,500 yBP. Holocene

paleoenvironment reconstruction in the Great Lakes is complex and controversial, but is a key to

the future climate predictions of the basin. Their emphasis on studying the present climate and

hydrology of the Great Lakes is focused from the increasing concerns of global warming and

lake levels. The assumption of a colder/ drier climate similar to Mirmeapolis, MN would have a

dramatic impact on the status of shallow lake basins in the Great Lakes region and dependent

industries along their shorelines. The modem Great Lakes are a vast finite water resource that

requires steady monitoring and conservation management methods for future generations. iijd i.yiiVj'ovi.

Hi vro.rn'.j"; Utr. ,:, T<: ;>*ii '.:;!•»< .j i!;;i?'

;-3'^^?. '!' (lOl'^V, r;VtS.I jil;;*)

.r''':::;!!^ V.^ -"I.- 123 References Cited

Aaby, B., 1976, Cyclic climatic variations in climate over the past 5500 years reflected in raised

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i: .''W.'-::-.:^ r .':>(. ;i l • 134

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n I

^. .,;/; Ui' • • '-M ,f\v»*V

nr»u-.!'"'.."i3y. V:> v^-iniii.tvi

'•;')•.»,'. ,/JJ':-' Appendix A

Water Budget, Climate, Vegetation, Lake Phases and Modern Analogue of Georgian Bay

136

Table 1: Water budgets for lakes Huron and Michigan ( • f-; 137

Georgian Bay: Monthly Water Temperatures 120 <0C Bbet.OC&SC abet. 5C & 10C abet. IOC & 15q abet. 15C & 20C abet. 20C & 25C abet. 25C & 30C 100

Jan Feb Mar Apr May Jun Jul Aug Sept Oct Nov Dec Months

Figure 1: Monthly water temperatures for Georgian Bay (modified from Saulesleja, 1986)

Georgian Bay: Monthly Air Temperatures 120 I<-20C Hbet. -20C&-10C Qbet. -10C&0C

Ibet. 0C&10C abet. 10C&20C Bbet. 20C&30C 100

Jan Feb Mar Apr May Jun Jul Aug Sept Oct Nov Dec Months

Figure 2: Monthly air temperatures for Georgian Bay (modified from Saulesleja, 1986).

Georgian Bay: IMonthly Precipitation Rates 35

30 ^

I 20 S a. 15

a. 10

138

Table 4: Vegetation zones of Georgian Bay (Note: Bl. = black, Wh. = white and East. = eastern) (Rowe, 1977). Zones .(' , t: .'.-'i' .'iVy

Wisconsinan (Isos.= isostatic) (Calkin & Feenstra, 1985; Karrow, 1989; Bamett, 1992; Lewis et al., 1994; Pengelly etal., 1997). Name '.'..: ojt>/

_« I- Appendix B

Thecamoebian and Loss-on-ignition data from Georgian Bay

141 Table 1: Thecamoebian data from cores LH95-041 and LH95-038.

142

LH41: Pollen zone 2b LH41: Pollen zone 2a (92-93cm)

143

97802-020: Pollen

144

978024)35: Pollen zone 3a (10-12cni) 18- .17 14- 10-

145 Table 4: Thecamoebian data from cores 97802-035 and 97802-036.

146 Nomenclature of Thecamoebians

The following thecamoebian assemblages were identified using Ogden and Hedley (1980); Kumar and Dalby

(1998).

Arcella vulgaris (Ehrenberg, 1830). (Plate 1, fig 1).

Arcella artocrea (Leidy, 1 876).

Centropyxis aculeata (Ehrenberg, 1 832) (Plate l, fig. 2 C. aculeata "aculeata", fig. 3 C. aculeata "discoides").

Centropyxis constricta (Ehrenberg, 1 843) (Plate l, fig. 4 C. conslricta "spinosa", fig. 5 C. comtricta "constricta").

Centropyxis platystoma (Penard, 1 890).

Cucurbitella tricuspis (Carter, 1856) Medioli, Scott and Abbott, (1987) (Plate 1, fig. 9).

Difflugia bidens (Penard, 1902) (Plate l, fig. 10)

Difflugia oblonga (Ehrenberg, 1 832) (Plate 1 fig. 12 D. obhnga "spinosa", fig 13. D. oblonga "oblonga", fig 14. D. oblonga "tenius").

Difflugia protaeiformis (Lamarck, 1 8 1 6) (Plate 1 , fig. 7 D. prolaeiformis "amphoralis", fig. 8 D. prolaeiformis "claviformis").

Difflugia urceolata (Carter, 1864) (Plate 1, fig. II D. wceolata'^urceolata")

Heleopera sphangi (Leidy, 1 874).

Pontigulasia compressa (Carter, 1864) Medioli and Scott, 1983 (Plate 1, fig. 6).

References

Carter, H.J., 1856, Notes on the freshwater Infusoria of the island of Bombay, no. 1 Organization, Annals and Magazine of Natural History

Carter, H.J., 1864, On fi-eshwater Rhizopoda of England and India, Annals and Magazine of Natural History, series 3, 13: 18-39

Ehrenberg, C.G., 1830, Organisation, systematik und geographisches verhaltnis der Infiisionsthierchen, Druckerei der Konigliche Akademie der Wissenchaften, Berlin, 1-108

Ehrenberg, C.G., 1832, Uber die Entwicklung und Lebensdauer der Infusionsthiere, nebst femeren Beitragen zu einer Vergieichung ihrer organischen Systeme, Konigliche Akademie der Wissenchaften zu Berlin Physikalische Abhandlungen, 1831, 1-154

Ehrenberg, C.G., 1843, Vertbreitung und Einfluss des mikroskopischen Lebens in Sud-und Nord Amerika, Konigliche Akademie der Wissenchaften zu Berlin Physikalische Abhandlungen, 1841, 291-446

Lamarck, J.B., 1816, Historic naturelle des animaux sans vertebres, Verdiere, Paris, 2: 1-568

Leidy, J., 1874, Notice of some new fresh-water rhizopods. Proceedings of the Academy of Natural Sciences of Philadelphia, series 3: 77-79

Leidy, J., 1879, Freshwater Rhizopods of North America in Vol. 12 United States Geological Survey of the Territories, Washington, pp. 324

Medioli, F.S., Scott, D.B., & Abbott, B.H., 1987, A case study of protozoan interclonal variability: taxonomic implications. Journal of Foraminiferal Research, 17: 28-47

Medioli, F.S. & Scott, D.B., 1983, Holocene Arcellacea (Thecamoebians) from eastern Canada, Cushman Foundation For Foraminiferal Research Special Publication 21

Penard, E., 1890, Etudes sur les Rhizopodes d'eau douce, Memoires de la Societe de Physique et d' Historic

Naturelle de Geneve, 3 1 (2), 1-230

Penard, E., 1902, Faune Rhizopodique du Bassin du Leman, Henry Kundig (Geneve), 1-134 rOT?;;. . ^i

•

^ - ' .( >v.-. Appendix C

Georgian Bay Climate, Water Budgets, Lake levels and Runoff Data

148 Table 1: Georgian Bay Climate data from Midland and Tobermory climate stations. Precipitation (P), Potential evaporation (PE) compared to lake level changes (LL) and subdivided into decades (Cook personal comm., 1998; Environment Canada, 2000).

149 Table 2: Paleo-water budget comparing present Georgian Bay climate to the 7,500 yBP lowstand using present Minnesota climate (Steinhauser, 1979; Cook, personal comm., 1998). Georgian Bay data was averaged from Tobermory and Midland climate records, All values are measured in cubic metres per second.

150

Minneapolis, Minnesota Climate

Jan Feb Mar Apr May Jun Jul Aug Sept Oct Nov Dec

w a.

Legend

Georgian Bay 500 yrs

-• Georgian Bay 700 yrs -50 Jan Feb Mar Apr May Jun Jul Aug Sept Oct Nov Dec

Figure 1: Paleo-water budget for Georgian Bay around 7,500 yBP using Minneapolis, Minnesota modem climate. The water budget illustrates the high evaporative summers and low winter precipitation interpreted to lower water levels in the Georgian Bay basin over a few hundred years. r 151 Table 4: 1/ 152

153 5c

154 5h

155 6e ? I 156 7b Al.' m

8b r{) tst

Table 10: Huron-Michigan lake levels (m) from 1918 to 1998 (tables 10a to lOh) (Environment Canada, 2000). 10a

159

lOf iti' ! 160 lie

161 llh

162

12e Coi Appendix D

Thecamoebian, Palynological, Loss-on-ignition and Field data from Willoughby Bog

164 Table 1: Thecamoebian data from Willoughby Bog core WB4. Willoughby Bog (WB4) |

165

WB4: Pollen zone 3a

166

Table 3 con't

167

Table 5 con't

168

Table 7 con't ' />rt 169

Table 9: Loss-on-ignition data ft ^OlH