A Thesis Entitled Chronology and Sedimentology of the Imlay

Home , Clear Lake (Michigan), Lake Saginaw, Lake Souris, Lake Whittlesey, Marshall Sandstone, Michigan Basin

A Thesis

entitled

Chronology and Sedimentology of the Imlay Channel, Lapeer County, Michigan

Jonathan Neil Luczak

Submitted to the Graduate Faculty as partial fulfillment of the requirements for the

Master of Science Degree in

Geology

______Dr. Timothy Fisher, Committee Chair

______Dr. David Krantz, Committee Member

______Dr. Donald Stierman, Committee Member

______Dr. Cyndee Gruden, Interim Dean College of Graduate Studies

The University of Toledo

December 2018

This document is copyrighted material. Under copyright law, no parts of this document may be reproduced without the expressed permission of the author. An Abstract of

Chronology and Sedimentology of the Imlay Channel, Lapeer County, Michigan

Jonathan Neil Luczak

Submitted to the Graduate Faculty as partial fulfillment of the requirements for the Master of Science Degree in Geology

The University of Toledo

December 2018

The Imlay channel on the thumb of Michigan was one of two outlets for the

Maumee stages (i.e. I–III or IV) of ancestral Lake Erie (ALE). The history and chronology of sediment within the channel is poorly understood, having been studied only sporadically over the last 100 years. Previous studies interpreted the floor of the channel to be the aerially exposed topographic divide near the intersection of Clear Lake and Van Dyke roads. However, geotechnical borings and water-well records along Clear

Lake Road suggest that the bedrock sill that controlled drainage through the channel may be buried by more than 12 m of sediment that post-dates channel formation.

A single radiocarbon age, i.e. ~16.7 calibrated ka BP, from ~6 km north of the

Imlay channel was interpreted in most of the literature to mark the cessation of channel activity; an age from organic material found within colluvium replicates this age in this study. Radiocarbon and OSL ages from fill within the Imlay channel from this study reveal that much of the fill was deposited by ~14.9 ka, with the uppermost fans appearing to have stabilized sometime in the early Holocene.

A better understanding of the study area was gained through the mapping of surficial deposits in the Burnside 7.5’ USGS quadrangle, the creation of a bedrock topography and drift thickness map, and the use of LiDAR data for identifying previously unrecognized moraines and kettle lakes. A plot of rebound curves for highest Maumee, lowest Maumee, and Lake Whittlesey strandlines, as well as a geologic cross-section along Clear Lake road in Lapeer County, Michigan, indicate that waters from all stages of

Lake Maumee could have drained over the bedrock sill of the Imlay channel, although other ALE stages (e.g. Lake Whittlesey and lower) could not.

This thesis is dedicated to my wife, Nikki, who was supportive of me throughout this process.

iii

Acknowledgements

I would like to thank everyone who helped me with this project, starting with my thesis committee: Dr. Timothy Fisher, Dr. David Krantz, and Dr. Donald Stierman. Your advice and time are very much appreciated. Also, thanks to Dr. Kenneth Lepper who processed the OSL data and offered advice on the interpretation of the results.

Next, I would like to thank everyone else who has helped me with the project, whether in the field or as a sounding board for ideas: J.P. Finnegan, John Dilworth, Eric

Armstrong, Tom Valachovics, Mitchell Dziekan, John Yellich, Everett Root, Dr. Harry

Jol, Dr. Henry Loope, Dr. Thomas Lowell, Dr. James Martin-Hayden, Dr. James Harrell,

Dr. Richard Becker, Dr. Carl Savage, Dr. Richard Freund, Dr. Thomas Modal, and soon- to-be Dr. Andrew Rankin. Thanks to my family who were supportive of this endeavor.

I would like to extend a special thanks to Brian Samsen. I couldn’t have asked for a better undergraduate researcher/field assistant for this project.

Thanks to all the property owners in Lapeer County who gave me permission to access their land and to the various organizations and municipal agencies who granted access as well.

Funding for the project was provided by: The U.S. Geological Survey’s EDMAP

Program through grant number G17AC00132 to Tim Fisher; the Geological Society of

America’s Graduate Research Grant; and the Ohio Geological Survey’s License Plate

Fund.

Table of Contents

Abstract ...... iii

Acknowledgements ...... iv

Table of Contents ...... v

List of Tables ...... vii

List of Figures ...... viii

1 Introduction ...... 1

1.1 Introduction ...... 1

1.2 Regional Deglaciation ...... 4

1.3 Cyclic Lake Level Model ...... 8

1.4 Episodic Lake Level Model ...... 16

1.5 Study Area ...... 17

1.6 Spillways ...... 22

1.7 Summary ...... 22

1.8 Objectives and Hypothesis ...... 23

2 Methods ...... 25

2.1 Introduction ...... 25

2.2 LiDAR Digital Elevation Model ...... 25

2.3 Sediment Collection ...... 26

2.4 Core Logging and Sediment Description ...... 26

2.5 Naming Scheme for Field Sites and Samples ...... 26

2.6 Radiocarbon Dating ...... 27

2.7 Optically Stimulated Luminescence Dating ...... 28

2.8 Particle Size Analysis ...... 28

2.9 Bathymetry Map of Clear Lake ...... 29

2.10 Surficial Geologic Map ...... 30

2.11 Geologic Cross-Sections ...... 30

2.12 Ground Penetrating Radar...... 31

2.13 Bedrock Topography and Drift Thickness Map ...... 31

2.14 Electrical Resistivity ...... 32

2.15 Horizontal to Vertical Spectral Ratio Survey ...... 33

2.16 Rebound Data...... 33

3 Results and Interpretations ...... 36

3.1 Introduction ...... 36

3.2 Geomorphology ...... 36

3.3 Imlay Channel and Clear Lake Sedimentary Sequences ...... 43

3.4 Deanville Mountain, Weaver Drain, Soil Texture ...... 56

3.5 Bedrock ...... 74

3.6 Geophysics ...... 77

3.7 Lake Maumee Water Levels and the Imlay Channel ...... 81

3.8 Surficial Geologic Map ...... 99

3.9 Summary ...... 101

4 Discussion ...... 102

4.1 Introduction ...... 102

4.2 Deglaciation ...... 102

4.3 Channel Fill and Chronology ...... 106

4.4 Water Drainage Through the Imlay Channel ...... 109

4.5 Rebound Curves ...... 110

5 Summary, Conclusions, and Future Work ...... 112

5.1 Summary ...... 112

5.2 Conclusions ...... 113

5.3 Future Work ...... 114

References ...... 115

Appendix ...... 124

A OSL Lab Results ...... 124

B Particle Size Analysis ...... 126

C Ground Penetrating Radar ...... 127

D Vertical Electrical Soundings ...... 129

E Tromino Stations ...... 132

F Field Locations...... 135

List of Tables

3.1 OSL Ages (This Study) ...... 50

3.2 Radiocarbon Ages (This Study) ...... 52

3.3 Bedrock and Drift Thickness ...... 77

vii

A.1 OSL Data: From Lab ...... 124

A.2 OSL: Elemental Analysis Data ...... 125

B.1 Results of Particle Size Analysis ...... 126

E.1 Site Location Data...... 135

List of Figures

1.1 Extent of Laurentide Ice Sheet ...... 2

1.2 Comparison of Cyclic and Episodic Models ...... 2

1.3 Ice Margins before ~16.6 ka ...... 3

1.4 Late Pleistocene Glacial Activity...... 5

1.5 Digital Elevation Model of Huron-Erie Moraines ...... 6

1.6 Ancestral Lake Erie Stages ...... 8

1.7 Digital Elevation Model of the Study Area ...... 11

1.8 Ice Advance Models ...... 20

1.9 Conceptual Model of Buried Spillway Valley ...... 23

3.1 Geomorphology of Northeastern Lapeer County ...... 37

3.2 Figure Reference Map of Northeastern Lapeer County ...... 38

3.3 Alluvial Fans ...... 40

3.4 Geomorphology of the Interlobate Area ...... 42

3.5 Whittlesey Spits ...... 44

3.6 Sedimentary Facies ...... 45

viii

3.7 Core Log: CORICC ...... 47

3.8 Core Log: CORICB ...... 49

3.9 Core Log: CORICA ...... 51

3.10 Stratigraphy: OSLJTEA ...... 53

3.11 Bathymetry Map of Clear Lake ...... 55

3.12 Core Log: CORCL1B ...... 57

3.13 Digital Elevation Model of Deanville Mountain and Weaver Drain ...... 59

3.14 Bechtal Sand and Gravel Pit (BSG) ...... 60

3.15 Till at BSG ...... 61

3.16 Stratigraphy: OSLBSGA ...... 62

3.17 Stratigraphy: OSLBSGB ...... 64

3.18 Core Log: CORSL2A ...... 67

3.19 Particle Size Analysis Diagram ...... 68

3.20 Stratigraphy: HASWDB ...... 71

3.21 Stratigraphy: HASWDE...... 72

3.22 Bedrock Topography and Drift Thickness Maps ...... 76

3.23 Vertical Electrical Sounding and Tromino Stations ...... 79

3.24 Digital Elevation Model of Maumee Strandline ...... 83

3.25 3D Model of Maumee Strandline ...... 84

3.26 Ground Penetrating Radar of Maumee Strandline ...... 85

3.27 Location of Water-Wells for Geologic Cross-Sections ...... 87

3.28 Clear Lake Road Cross-Section (Water-Well and Geotechnical Data) ...... 88

3.29 Clear Lake Road Cross-Section (Interpretation) ...... 90

3.30 Longitudinal Cross-Section of the Imlay Channel...... 92

3.31 Rebound Curve of Three ALE Stages ...... 96

3.32 ALE Water Planes Adjusted for Glacial Isostatic Rebound ...... 98

4.1 Deglaciation of the Saginaw Lobe ...... 105

C.1 Ground Penetrating Radar at Goodland Cemetery ...... 127

D.1 Processed Vertical Electrical Sounding Data ...... 129

E.1 Processed Tromino Data ...... 132

Chapter 1

Introduction

1.1 Introduction

Glaciers from the Laurentide Ice Sheet (LIS) covered much of northern North

America during the Wisconsinan glacial period, which lasted from ~75–11 ka (Mickelson et al., 1983; Barnett, 1992). The Saginaw and Huron-Erie lobes of the LIS reached their southernmost extent in Ohio and Indiana about 23 ka during the last glacial maximum

(LGM) (Figure 1.1). As the ice retreated from southern Indiana and Ohio following the

LGM, the ice made several readvances during cold periods, i.e. stadials (Figure 1.2).

During the Port Bruce Stadial, ice from the Huron-Erie Lobe extended into Indiana before retreating to the northeast once again. When the glacier began retreating downslope, meltwater became dammed between a moraine in the south and the LIS in the northeast, forming ancestral Lake Erie (ALE) (Figures 1.3 and 1.4). Likewise, water from the ALE basin would take advantage of topographic lows uncovered by the retreating ice sheet. Lake Maumee was the first highest proglacial lake to form in the ALE basin following the Port Bruce Stadial (Figures 1.2 and 1.3). Two outlets for Lake Maumee were proposed: the Imlay channel on the thumb of Michigan and another at Ft. Wayne,

Indiana (Figure 1.3). The Ft. Wayne outlet (FWO) routed water from Lake Maumee directly to the Gulf of Mexico via the Wabash-Mississippi river system. Large volumes of sediment-laden water flowing directly to the Gulf of Mexico would likely induce hyperpycnal flows into the deep ocean basin and would show up in marine isotopic

Figure 1.1. Maximum extent and lobes of the Laurentide Ice Sheet in the Great Lakes region of North America at the Nissourian (last glacial) maximum (<26–23 ka) and during the Port Bruce Maximum (~18 ka) (Modified from Kehew, 2012).

Figure 1.2. Comparison of the cyclic (black line) and episodic (grey line) models of lake levels in the ALE basin (Fisher et al., 2015).

Figure 1.3. Approximate ice margins of the Laurentide Ice Sheet just before ~16,600 cal. years BP (solid line), when activity in the Imlay outlet was estimated to have ceased (Burgis, 1970). records there (Mulder et al., 2003; Roche et al., 2007). Conversely, water draining from

Lake Maumee via the Imlay channel would have possibly been buffered by early Lake

Saginaw, and certainly by glacial Lake Chicago (Figure 1.3). Plant litter collected in the

1960s by Burgis (1970) at the Weaver Drain site, described as a glacial outwash channel northeast of the Imlay channel, is interpreted to represent deglaciation in the area, as well as the end of the Lake Maumee stage of ALE, and thus cessation of channel activity

(Burgis, 1970).

The Imlay channel has been described in the literature for over 100 years, but little work has been done on it in that time. An early study (Leverett and Taylor, 1915) identified recessional moraines on the thumb of Michigan which suggested an oscillatory margin of the Saginaw and Huron-Erie lobes and reported fluctuating lake levels within the ALE basin were the result of those oscillations (Figures 1.2 and 1.4).

1.2 Regional Deglaciation

At the LGM, during the Nissourian Stadial, the LIS advanced as far as southern

Ohio and central Indiana (Figure 1.1), depositing the Hartwell Moraine near Cincinnati,

Ohio around 19,590 ± 35 14C years BP (~23.6 calibrated ka BP), which represents the average radiocarbon age of four sites (Lowell et al., 1999). As the climate warmed during the Erie Interstadial (Figure 1.4), the LIS retreated to the north and splintered into several lobes (Figure 1.1), resulting in the formation of several proglacial lakes within the Great

Lakes basins. Subsequent advances of the Huron-Erie Lobe (Figure 1.1) during the Port

Bruce Stadial (Figure 1.4) overrode any lacustrine deposits in the ALE basin and resulted in the deposition of the Union City-Powell Moraine in central Ohio and Indiana at its southernmost position (Figure 1.5) (Mickelson et al., 1983; Barnett, 1992). A radiocarbon age for the Union City-Powell Moraine comes from wood found within till and suggests a maximum age of 14,780 ± 192 14C years BP (~18.0 calibrated ka BP) (OWU-83) (Ogden and Hay, 1965). However, the stratigraphic significance of the sample is questionable

(Michelson et al., 1983) with Ogden and Hay (1965) reporting that the till in which the wood was found may have been a slump block located to the north of the Union City-

Powell Moraine near Liberty, Ohio. As the ice margin retreated from the Union City-

Figure 1.4. Late Pleistocene glacial activity in the southern Great Lakes region (Fisher et al., 2015).

Powell Moraine, several recessional moraines formed during the Huron-Erie Lobe’s northward retreat (Mickelson et al., 1983). From oldest to youngest these recessional moraines include the Mississinawa, Salamonie, the Wabash, the Ft. Wayne, and the

Defiance (Figure 1.5) (Goldthwaite et al., 1965). A minimum radiocarbon age for the deglaciation of the Wabash Moraine comes from two sites within Ohio. One from the

Quillen site near Lodi, Ohio consists of a basal layer of organic material overlying Late

Woodfordian (i.e. Hiram) till. The organics yield an age of 14,500 ± 150 14C years BP

(~17.7 calibrated ka BP) (GS-402) (Crowl, 1980). The other age for the Wabash Moraine is 14,290 ± 130 14C years BP (~17.4 calibrated ka BP) (ISGS-72) (Coleman, 1973), and comes from wood found in a basal layer of organic-rich

Figure 1.5. Digital elevation model showing recessional moraines of the LIS’s Huron-Erie Lobe following the Port Bruce advance. Yellow hexagon indicates location of Ubly outlet.

silt overlying Late Woodfordian (i.e. Hiram) till near Shiloh, Ohio. One radiocarbon age for the Fort Wayne Moraine was reported by Shane and Anderson (1993) to be 14,680 ±

310 14C years BP (~17.9 calibrated ka BP) (IGS-1679). However, Dyke (2003) retired the age, suggesting that the material that was dated contained an excess of old carbon. Two ages for basal organics from the Ft. Wayne Moraine come from a study by Glover et al.

(2011) in northwestern Ohio and are 13,880 ± 70 14C years BP (~16.9 calibrated ka BP)

(Beta-190864) and 13,690 ± 50 14C (~16.8 calibrated ka BP) (Beta-194054). However,

Glover et al.’s (2011) ages for the Ft. Wayne Moraine suggest that deglaciation of the area occurred ~1,000 years later than expected, considering the ages associated with the older Wabash Moraine, i.e. ~17.5 calibrated ka BP (Colman, 1973; Crowl, 1980).

Continued northward retreat of the Huron-Erie Lobe resulted in closed basins beginning to fill with meltwater from the LIS, forming a series of proglacial lakes in the ALE basin

(Figures 1.5 and 1.6).

Until recently, moraines associated with the Saginaw Lobe were sparsely studied and had little chronologic data associated with them. The Port Huron Moraine is geographically correlated with both the Saginaw and Michigan lobes of the LIS. An age for the Michigan Lobe’s Port Huron Moraine in northwestern Michigan is 12,960 ± 250

14C years BP (~15.3 calibrated ka BP) (TX-651) (Blewett et al., 1993). The Sturgis

Moraine (Figures 1.3 and 1.5) in northern Indiana and southern Michigan is also associated with the Saginaw Lobe. Radiocarbon dates on material collected from kettle lakes on the Sturgis Moraine resulted in a minimum age of 13,750 ± 80 14C years BP

(~16.6 calibrated ka BP) (OS-111986) (Horton, 2015) for the moraine. More recently,

Figure 1.6. Stages of ALE and associated outlets according to most recent literature (Larson and Schaetzl, 2001).

Dziekan (2017) used OSL on basal lake deposits on the Sturgis Moraine which resulted in new minimum age of 20.4 ± 1.4 ka for the moraine.

1.3 Cyclic Lake Level Model

Leverett and Taylor (1915) proposed that a series of transgressive and regressive sequences within the ALE basin was controlled by the oscillatory retreat of ice from the

Ft. Wayne Moraine. Observations that lead to Leverett and Taylor’s (1915) conclusion of fluctuating water levels within the ALE basin include a “washed-over appearance” of several strandlines within the basin due to their faint and discontinuous nature.

Lake Maumee was a proglacial lake that covered portions of northeastern Indiana, northwestern Ohio, and southeastern Michigan (Figures 1.3 and 1.6). Lake Maumee was the first proglacial lake to form in the ALE basin following the retreat of the Huron-Erie

Lobe from the Ft. Wayne Moraine (Figures 1.2, 1.3, 1.4, 1.5, 1.6) (Leverett and Taylor,

1915). Fisher et al. (2105) used optically stimulated luminescence (OSL) to date shoreline sediment deposited in Lake Maumee in northwestern Ohio, which yielded an age of ~16.8 ± 1.0 ka. Most of the literature divides Lake Maumee into three (Leverett and Taylor, 1915; Hough, 1963; Burgis, 1970; Eschman and Karrow, 1985; Larsen and

Schaetzl, 2001) or four levels (Fullerton, 1980; Calkin and Feenstra, 1985). Following the cyclic model, water levels for Maumee are generally described as follows, in decreasing age: Maumee I (244 masl), Maumee II (232 masl), Maumee III (238 masl), and Maumee

IV (232 masl) (Figure 1.2).

There is little doubt that the FWO near Ft. Wayne, Indiana (Figures 1.3, 1.5, 1.6) was responsible for draining one or several stages of Lake Maumee (Leverett and Taylor,

1915; Burgis, 1970; Bleuer and Moore, 1971). Morphologic studies (Leverett and Taylor,

1915), as well as modern digital elevation models (DEMs) show that the Ft. Wayne,

Wabash, and Union City-Powell moraines were cross-cut by the Wabash River system which drained Lake Maumee (Figures 1.3 and 1.6). Furthermore, stratigraphic work at the FWO by Bleuer and Moore (1971) showed that the sill of the outlet was at least 3 m lower than the surface elevations used by earlier workers (Leverett and Taylor, 1915) and was capable of draining every stage of Lake Maumee.

The Imlay channel, located on the thumb of Michigan in Lapeer County (Figures

1.3, 1.6, 1.7), is also reported to be one of the outlets that drained one or more levels of

Lake Maumee when ice had retreated far enough northward (Leverett and Taylor, 1915;

Burgis, 1970; Eschman and Karrow, 1985). While little work has been done on the Imlay channel in the past 100 years, the channel is interpreted to have been a conduit between the ALE and Lake Michigan basins (Leverett and Taylor, 1915) when the Saginaw and

Huron-Erie lobes retreated far enough northward to uncover it (Figure 1.3). Water draining from the Imlay channel flowed from east to west, towards the Lake Michigan basin, by following the ice margin of the Saginaw Lobe before entering the Grand River

Valley (GRV) directly near Maple Rapids, Michigan (Figures 1.3, 1.5, 1.6, 1.7) (Leverett and Taylor, 1915), or possibly by first entering the Saginaw lowlands that formed between Maple Rapids, Michigan and the retreating ice margin of the Saginaw Lobe, forming “early Lake Saginaw” before entering the GRV (Figures 1.3 and 1.5) (Leverett and Taylor, 1915; Hough, 1958; Eschman and Karrow, 1985). More than half a century after Leverett and Taylor’s (1915) work on the Imlay channel, Burgis (1970) collected plant litter from the Weaver Drain ~6 km from the Imlay channel (Figures 1.3, 1.5, 1.7) for radiocarbon dating. The plant material, collected from interbedded layers of silt, sand, and clay overlying poorly sorted sand and till, yields an age of 13,770 ± 210 14C years BP

(~16.7 calibrated ka BP) (I-4899) and is reported to be a minimum age that marks the end of Maumee III and abandonment of the Imlay channel (Burgis, 1970; Buckley and Willis,

1972; Fullerton, 1980; Calkin and Feenstra, 1985). However, it is unclear why material collected from the Weaver Drain was used to determine cessation of activity within the

Imlay channel since it lies ~6 km to the northeast of the channel. The Weaver Drain has been described in studies as being located in a glacial outwash channel (Buckley and

Figure 1.7. Map and digital elevation model illustrating the geomorphology of the study area. Weaver Drain site (red star), Imlay channel divide (green hexagon), Clear Lake (blue hexagon), and the Burnside 7.5’ USGS quadrangle, Kaeregnondi Water Authority Pipeline (green line).

Willis, 1972) and also as being a part of the Maumee lake plain (Eschman and Karrow,

1985).

The next lake stage in the ALE basin was Lake Arkona (Leverett and Taylor,

1915). Three stages of Lake Arkona are proposed (I: 216 masl; II: 213 masl; III: 212 masl) (Calkin and Feenstra, 1985). Several researchers (Hough, 1958; Eschman and

Karrow, 1985) suggest that the three Arkona lake levels were established due to rapid down-cutting of the sill at the head of the GRV near Maple Rapids in the Saginaw lowlands (Figures 1.3 and 1.5), where water from Lake Arkona would have drained through the GRV towards the Lake Michigan basin. For example, rapid down-cutting of the sill would have resulted in catastrophic flooding of Arkona I waters down the GRV until water was lowered to the sill where it established Arkona II, with the process repeating for one last time to establish water levels to those of Arkona III. Calkin and

Feenstra (1985) suggested that while down-cutting of a sill at the head of the GRV may have contributed to three stages of Lake Arkona, another factor may have been isostatic adjustment following deglaciation in the area, which would have tilted the ALE basin to the south. Kehew (1991) described boulder lag deposits and fluvial bar deposits comprised of cobble-sized clasts within the GRV as evidence for catastrophic flooding in the area, possibly due to the rapid down-cutting of a sill near Maple Rapids. A minimum age for Lake Arkona is 13,600 ± 500 14C years BP (~16.4 calibrated ka BP) (W-33)

(Suess, 1954) and comes from a lagoon deposit near Cleveland, Ohio (Hough, 1958).

Goldthwait (1958) reports that the dated wood was found underlying ~3 m of sandy silt and overlying till at an elevation of ~210 masl and interprets the deposit as recording

“water rise at the end of Lake Arkona.” Dreimanis (1977) interprets the deposit as Lake

Arkona-Lake Whittlesey transition beach (Dreimanis, 1977), while Fullerton (1980) suggests the deposit represents the end of Arkona III. However, recent work by Dilworth

(2017) in northern Ohio and southern Michigan suggests that Lake Whittlesey did not overtop littoral sediments associated with Lake Arkona. The study by Dilworth (2017) concluded that Arkona’s washed appearance was, in part, due to its low relief, i.e. 1.5–2 m, and discontinuous nature. A preliminary minimum OSL age for a deposit associated with Lake Arkona in western Ohio is 12.4 ± 1.2 ka (Dilworth, 2017).

The idea of a lowstand, i.e. the Ypsilanti lowstand, in the ALE basin between the

Arkona and Whittlesey stages was first introduced by Hough (1958). Evidence from

Hough (1958) supporting a lowstand in the ALE basin were crossbedded (nearshore) deposits overlain by fine grained (deep water) deposits found in a gravel pit near

Cleveland, Ohio from projected subsurface elevations of ~206 – 200 masl. The crossbedded nearshore deposits were interpreted to be representative of a lowstand within the ALE basin and the fine-grained sediments were representative of deeper water following a transgressive sequence within the basin. Kunkle (1963) also proposed that a lowstand existed in the ALE basin during the Mackinaw Interstadial (Figure 1.4) known as Lake Ypsilanti. Evidence for Lake Ypsilanti is based on crosscutting relationships between the ancestral Huron River Valley (HRV) (Figure 1.5) in southeastern Michigan and glacial, fluvial, and lacustrine deposits there. According to

Kunkle (1963), streams in the ancestral HRV deposited deltaic sediments into water associated with Lake Maumee. As lake levels fell in the ALE basin, presumably due to the ice margins retreating far enough to the north to establish an easterly drainage route, the HRV incised the deltaic deposits associated with Lake Maumee, and cut into till

deposited during the Nissourian Stadial (Figure 1.4). Lacustrine sand and gravel overlying early Woodfordian till in the incised ancestral HRV were reported to have been deposited by transgressive lake sequences associated with lakes Whittlesey and/or

Warren during the advance of the LIS during the Port Huron Stadial (Figure 1.4)

(Kunkle, 1963; Calkin and Feenstra, 1985). Kunkle (1963) estimated that the water levels of Lake Ypsilanti ranged in elevation between ~115 –166 m, i.e. 45 – 95 m below

Arkona III (210 m), based on profiles he created for the ancestral HRV that extrapolated the gradient into the ALE basin. However, Campbell (2011) suggested the ancestral

HRV may have formed due to the presence of a tunnel channel that developed on the ice margin of the Saginaw Lobe. Rhythmites interpreted as varves (Anderson, 2013) found in northwestern Ohio are associated with the Maumee through Warren stages of ALE.

However, Anderson (2013) explained that the varve sequence showed no signs of ravinement that might suggest a transgressive sequence in the ALE basin between the times when lakes Maumee and Warren were established in northwestern Ohio.

The advancing ice margin during the Port Huron Stadial (Figure 1.4) resulted in a transgressive lake sequence in the ALE basin that resulted in the formation of Lake

Whittlesey, and the deposition of the Port Huron Moraine, on the thumb of Michigan

(Figure 1.5). During the transgressive event, it has been assumed by previous workers that Lake Whittlesey overtopped sediments deposited by Lake Arkona, and possibly Lake

Ypsilanti, before draining into the Saginaw lowlands via the Ubly outlet and down the

GRV to the Lake Michigan basin (Figures 1.5 and 1.6) (Eschman and Karrow, 1985). The elevation of Lake Whittlesey is reported as ~225 masl near Birmingham, Michigan and rises to ~246 masl near the head of the Ubly channel (Figure 1.5) (Eschman and Karrow,

1985). Maximum ages for the deposition of Lake Whittlesey are 12,920 ± 400 14C years

BP (~15.36 calibrated ka BP) (W-430) and comes from wood in peat below gravels associated with the lake near Parkerstown, Ohio (Rubin and Alexander, 1958). However,

W-430 has also been associated with the Lake Ypsilanti lowstand (Fullerton, 1980).

Another maximum age for Lake Whittlesey is 12,800 ± 250 14C years BP (~15.0 calibrated ka BP) (Y-240), which comes from wood fragments found within Lake

Whittlesey sediments near Bellevue, Ohio (Barensden et al., 1957; Calkin, 1970).

Another possible maximum age for the formation of Lake Whittlesey is 12,900 ± 200 14C years BP (~13.46 calibrated ka BP) (I-3175) from organics collected in Elyria, Ohio.

However, Calkin (1970) describes the stratigraphic horizon of the material dated in I-

3175 simply as “Wood in Whittlesey beach.” A minimum age for Lake Whittlesey of

12,370 ± 220 14C (15.09 calibrated ka BP) (I-3665) comes from either organic material in lake silts (Calkin, 1970), or from wood found in peat overlain by alluvium (Buckley,

1976) recovered near North Collins, Ohio. Calkin (1970) also indicates that the date for I-

3665 may be a minimum age for the recession of the Gowanda Moraine in western New

York state. A log excavated from a depth of 10–12 m from gravels associated with the

Lake Arkona-Lake Whittlesey transition in southern Ontario yielded a radiocarbon age of

12,660 ± 400 14C years BP (~14.94 calibrated ka BP) (S-31) (Dreimanis, 1966), although

Fullerton (1980) attributes this age to Lake Ypsilanti. Recently, deposits associated with

Lake Whittlesey in southeastern Michigan were sampled for OSL and yielded an age of

15.7 ± 1.5 ka (Dilworth, 2017).

1.4 Episodic Lake Level Model

While many researchers (Forsyth, 1959; Burgis, 1970; Calkin and Feenstra, 1985;

Monaghan and Hansel, 1990) support the cyclic model of lake levels within the ALE basin proposed by Leverett and Taylor (1915), several workers suggested an alternative model. A study of ALE shorelines in northeastern Ohio found no evidence of Leverett and Taylor’s (1915) fluctuations (Totten, 1985), and instead concluded that a gradual, episodic decrease in ALE levels occurred (Figure 1.2). Totten (1985) reasoned that clay material deposited over lacustrine sediments in the ALE basin might actually be reworked Late Woodfordian till. Totten (1985) also argued that a slow rise in lake level within the ALE basin would likely result in the destruction of shoreline features deposited during earlier lake stages due to erosion by wave action.

Most recently, a study in northwestern Ohio (Fisher et al., 2015) has also challenged the generally accepted convention of the cyclic model for ALE and, like

Totten (1985), suggests an episodic drop in lake levels over time (Figure. 1.2). The study used stratigraphic (ground penetrating radar [GPR], electrical resistivity, hand auguring), chronologic (OSL) and morphological (examination of DEMs) evidence to reassess the

“washed-over” Maumee II shorelines that were most recently mapped by Forsyth (1959) prior to the 2015 study. Fisher et al. (2015) found the washed Maumee II shorelines were most likely littoral deposits that formed as waves in ALE reworked diamicton. In the study (Fisher et al., 2015) OSL ages for sediment deposited as Lake Maumee shorelines in northwestern Ohio were reported to be about 16.8 ± 1.0 ka. In agreement with Totten

(1985), the Fisher et al. (2015) asserts that the transgressive sequence that resulted in the

“washed-over” appearance of Maumee II shorelines would have to occur in a relatively short period of time in order to preserve such strandlines.

1.5 Study Area

While the existence of the Imlay channel on the thumb of Michigan has been known for well over 100 years, only two groups of workers have done fieldwork in the area. Leverett and Taylor (1915) were the first workers to recognize the “northern” drainage divide to the west of Deanville Mountain, reporting the elevation of the divide to be ~244 masl, and suggesting that it may have controlled water flowing through the

Imlay channel. Two other drainage divides with similar elevations were also reported by

Leverett and Taylor (1915) near Imlay City (244 masl) and Almont (244 masl), Michigan

(Figure 1.7). However, Burgis (1970) recognized that Leverett and Taylor’s ‘northern’ divide, which will from now on be referred to as the topographic divide (Figures 1.7 and

1.8), stands at an elevation of ~250 masl.

Leverett and Taylor (1915) also reported that oscillations of the LIS’s margins were the cause of lake level changes in ALE and resulted in the deposition of several moraines near the Imlay channel (Figures 1.7 and 1.8). Leverett and Taylor (1915) suggested that the Deanville, Goodland, Imlay, Yale (Figures 1.7 and 1.8) and Otisville moraines (Figure 1.8) represent a minor readvance of the Saginaw and/or Huron-Erie lobes in the area. According to Leverett and Taylor (1915), as the ice sheet retreated north of the Imlay area, a low outlet was uncovered that allowed Maumee I (244 masl) to drop to Maumee II (232 masl) levels. A readvance of the ice margin sometime after

Maumee II levels were established, pinched off or buried the unknown lower outlet causing the water to increase to Maumee III (238 m) levels.

Burgis (1970) tested Leverett and Taylor’s hypothesis concerning the readvancing ice margins in the area by examining the till lithology near the Imlay channel of what are presumed to be glacial deposits. Burgis (1970) reasoned that the sandstone-rich till was locally derived from the underlying Mississippian Marshall Sandstone and Coldwater

Shale bedrock formations, while the limestone-rich till was likely transported from the

Huron basin. Burgis (1970) found that the sandstone-rich till was absent west of the

Imlay channel, in the interlobate area that formed between the Saginaw and Huron-Erie lobes as the ice retreated across the thumb of Michigan. Burgis concluded that the limestone-rich till was deposited by an earlier advance of and that the sandstone-rich till represented minor, local readvances of ice that halted east of the Imlay channel. Burgis’ study found two distinctive till clast lithologies, one limestone/crystalline-rich (mean composition: ~52% limestone; ~17% sandstone and shale; ~31% crystalline) and the other sandstone/shale-rich (mean composition: ~32% limestone; ~52% sandstone and shale; ~16% crystalline).

Three conceptual models (Figure 1.8) are shown for the ice margin at the Imlay channel during the retreat of the LIS across Michigan. In all models the Imlay channel becomes free of ice after ice retreats from its position at time 1 (Figure 1.8). The first conceptual model (Figure 1.8C) is an interpretation of Burgis’ (1970) for ice readvance, where deglaciation occurs with ice margins at time 1 (T1; oldest) pulling back to T2, somewhere to the north of the study area; ice margin readvances occur in both the

Saginaw and Huron-Erie lobes to T3 over Deanville Mountain, depositing sandstone-rich

till on top of limestone-rich Deanville Mountain kame deposits, as well as the Imlay, and

Otisville moraines; at T4 the Goodland Moraine is deposited; and at T5 the Deanville and

Yale moraines are deposited. A second model of ice retreat near the Imlay channel shows no readvance of either the Saginaw or Huron-Erie lobes (Figure 1.8D) and begins at T1

(oldest) following deglaciation of the interlobate area; at T2 the sandstone-rich Otisville and Imlay moraines are deposited along with sandstone-rich diamicton over Deanville

Mountain; at T3 the Goodland Moraine is deposited; at T4 the Yale and Deanville moraines are deposited; at T5 the ice margin begins to pull back into the ALE basin. The third model of ice retreat near the Imlay channel study area (Figure 1.8E) is a readvance- model similar to the one depicted in Figure 1.8A, except that only the Huron Lobe readvances; at T1 the interlobate moraine has been deglaciated; at T2, the ice margin is marked by the Otisville Moraine near Lake Silverwood; at T3 the ice pulls back just north of the study area; at T4 ice from the Huron Lobe readvances into the ALE basin, depositing the Imlay Moraine; at T5 the ice margin retreats and deposits the Goodland

Moraine; at T6 the Deanville and Yale moraines are deposited; following T6 the ice margin retreats northeast into the ALE basin, uncovering lower outlets on the thumb of

Michigan.

Unfortunately, Burgis’ study (1970) did little to elucidate water flow through the

Imlay channel. Burgis’ (1970) conclusions agreed with those of Leverett and Taylor

(1915) that the retreating ice margin uncovered an outlet, now buried, somewhere to the north of the Imlay channel that allowed Maumee I to drain to Maumee II levels. She speculated that the location of the buried outlet is possibly observed today as a series of smaller channels

Figure 1.8. (See page 21 for caption)

Figure 1.8. Regional map of the Great Lakes (A; pg. 20) and a diagram (B; pg. 20) showing moraines (Deanville, Flint, Goodland, Imlay, Otisville, Yale), channels (North Branch, Elm Creek, Lum, Weaver), Silverwood lake plain, Deanville Mountain (solid orange line), Imlay channel (blue arrow) Flint River, Imlay channel divide (green hexagon). A conceptual model (explained in text) of ice dynamics in the area is (page 21) and depicts C) readvance of the Saginaw and Huron-Erie Lobes, after Burgis (1970); D) No readvance in the study area; E) Readvance of the Huron-Erie Lobe only.

in the study area (i.e. the North Branch, the Elm Creek and Weaver Drain channels)

(Figure 1.8).

1.6 Spillways

Megafloods associated with proglacial lakes in the upper Midwest have occurred since the Late Pleistocene and have been well documented (Kehew and Boettger, 1986;

Fraser and Bleuer, 1988; Kehew, 1992; Fisher, 2003; Kozlowski et al., 2005; Kehew et al., 2009; Curry et al., 2014). As the LIS retreated downslope, water from the melting ice filled the topographic depression. Dammed by the ice, the lakes took advantage of topographic lows uncovered by the LIS as it retreated, draining from these areas. The spillways of these lakes have common features including, streamlined residual hills, deeply cut (25–150 m) channels that are relatively uniform in width, and boulder lags

(Kehew et al., 2009). However, these geomorphic features can be absent if the floor of the spillway was buried by sediment, e.g. alluvium, following channel activity. At the former site of glacial Lake Souris, which existed in North Dakota, Kehew and Boettger

(1986) reported that the floor of the spillway associated with it was buried by tens of meters of Holocene sediment (Figure 1.9). Similarly, the floor of the FWO (Figure 1.5) was reported to have been buried by more than 3 meters of sediment after the channel was active (Bleuer and Moore, 1972).

1.7 Summary

The Imlay channel in Lapeer County, Michigan is one of two outlets for the Lake

Maumee stages of ALE, with the second found near Ft. Wayne, Indiana. However, it is

Figure 1.9. Conceptual model of the buried spillway of glacial Lake Souris in North Dakota. Modified from Kehew and Boettger (1986). still unclear which stages of Lake Maumee were accommodated by the channel near

Imlay. It is also unclear whether other stages of ALE, e.g. Lake Whittlesey, might have been able to drain through the outlet at Imlay. Only one radiocarbon age (Burgis, 1970) exists in the area that is associated with the draining of Lake Maumee through the Imlay channel, and a thorough understanding of the stratigraphy and chronology of the channel itself is currently lacking.

1.8 Objectives and Hypothesis

The purpose of this study is to determine which stages of ALE were able to drain through the Imlay channel. Thus, a logical question would be: was the floor of the Imlay channel buried by alluvial sediment after the channel was abandoned?

The hypothesis to be tested in this study is that water from any stage of Lake

Maumee in the ALE basin was able to drain via the Imlay channel. In order to test this hypothesis three objectives will need to be met:

1) Determining the chronology and stratigraphy of sediment within the Imlay

channel.

2) Detailed mapping of the surficial geology proximal to the channel.

3) Developing a better understanding of the bedrock geology in the area.

Chapter 2

Methods

2.1 Introduction

A total of fourteen methods were used to meet the objectives of this project.

These methods include the generation of a LiDAR-derived digital elevation model

(DEM), sediment collection, logging of cores and sediment, radiocarbon dating, optically stimulated luminescence (OSL) dating, particle size analysis, mapping the bathymetry of a lake, mapping surficial geology, geologic cross-sections, ground penetrating radar

(GPR), bedrock and drift thickness maps, electrical resistivity, and calculating rebound curves in the ancestral Lake Erie (ALE) basin using tools in ArcGIS.

2.2 LiDAR Digital Elevation Model

Unpublished LASer (LAS) files of Lapeer County were acquired from an airborne

LiDAR survey conducted by the Federal Emergency Management Agency in 2016. The raw LAS files were converted into a DEM in ArcMap using the methodology of Teft

(2015). The LAS files have a projection of North American Datum (NAD) 1983 State

Plane Michigan South (international feet). Vertical datum is NAD 1983 2011. Cell sizes for the DEM created from the LAS files are 3 ft2.

2.3 Sediment Collection

Sediment was collected using vibracore and hydraulically-assisted Livingstone coring systems in bogs and from water-based platforms, respectively. Sediment was also collected using hand augers, hand-dug pits, and at outcrop exposures within the study area.

2.4 Core Logging and Sediment Description

Sediment cores were examined in the Glacial Lakes and Sediment Science

(GLASS) lab at the University of Toledo. Cores were photographed shortly after opening using the methodology of Horton (2016). The photos were color-corrected and mosaicked using Adobe Photoshop. Cores were examined for sedimentary structures and examined under a microscope, and descriptions of the cores were recorded. When not in use, cores were shrink-wrapped and stored in a cooler at 4o C.

2.5 Naming Scheme for Field Sites and Samples

Field sites and samples were named as follows. The first three letters indicate the primary purpose of the site: Coring (COR), ground penetrating radar (GPR), hand auger/shovel/outcrop observation (HAS), optically stimulated luminescence (OSL), vertical electrical sounding (VES), and Tromino (TRM). The next set of letters correspond to location ID: Weaver Drain (WD), Imlay channel (IC), scour lake (SL),

Clear Lake (CL), Brooks Road (BROOKS), Brown City Road (BRWNCTY), Cade Road

(CADE) Clear Lake Road (CLKRD), Linck Road (LINCK), Marsh Road (MARSH),

Peck Road (PECK), Stine Road (STINE), Bechtal Sand and Gravel (BSG), J.T. Express

(JTE). The pit, station, sample, or core taken at the location is given by a number: 1, 2, 3, etc. The final letter is the site ID within the location, e.g. a transect of geophysical data across the Imlay channel: A, B, C, D, E. Note, if there is only one pit, core, or station no number is given between the location ID and the site ID. As an example TRMIC12A indicates that Tromino data was collected within the Imlay channel and it is the 12th station taken across transect A. Likewise, HASWDA indicates that a hand auger was performed within the Weaver Drain at site A; since it is the only hole at that site there is no number between the location ID and site ID. Radiocarbon samples were named according to the core they were taken in followed by a number which corresponds to the sample, e.g. CORSL2A-1, CORSL2A-2, etc. Only one OSL sample was collected at any location, so the OSL sample name is the same as the core name or sample location, e.g.

OSLBSGA, CORICC.

2.6 Radiocarbon Dating

Radiocarbon dating is a common method for dating events in the Quaternary and relies on the presence of organic material. Radiocarbon (14C) has a half-life of 5,568 years (Mook, 1986) and is effective for dating organic material up to ~40,000 years old.

Sediment from cores at stratigraphically significant locations, e.g. unit boundaries, were wet-sieved to 0.063 mm and rinsed in deionized (DI) water. Samples with a high clay content were treated for 1–8 hours with sodium hexametaphosphate to disperse the fine particles before wet-sieving. Samples were examined under a microscope for terrestrial plant macrofossils, with a preference for wood fragments. Organic material was photographed and cleaned with DI water. When applicable, organic material was stored

in a 4o C cooler in a plastic vial of DI water treated with one drop of 10% HCl. Organic material was dried in a muffle furnace at 50o C, weighed, and sent to either Woods Hole

Oceanographic Institute in Massachusetts or Beta Analytics, LLC in Florida for dating.

2.7 Optically Stimulated Luminescence Dating

OSL ages are useful because they date when sediment, e.g. quartz sand, was buried and does not require the presence of organic material (Murray and Olley, 2002;

Rittenour, 2008). Samples for OSL dating were collected in hand-dug pits, at outcrops, and from unopened sections of aluminum vibracore tubes. In hand-dug pits and outcrops, samples were collected at least 1 m below the ground surface. Samples from sediment collected using vibracore systems were taken from the lowermost ~30 cm of the unopened vibracore tube. Sediments just above the lowermost 30 cm of the tube were assumed to be the same as those found within the part of the tube sent for analysis. All

OSL samples were sent to Dr. Ken Lepper at North Dakota State University for analysis.

2.8 Particle Size Analysis

Particle size analysis was performed on samples from hand-augured sites in the study area. Samples were air-dried overnight in the GLASS lab. The dried samples were gently ground using a mortar and pestle. Samples were sieved for particles > 2 mm in diameter and placed in conical centrifuge tubes. Dry samples in the centrifuge tubes were homogenized by shaking the tubes, and ~2 g of dried sample was weighed out. The 2 g samples were treated with ~1 ml dose of 30% H2O2 until effervescence ceased, and sometimes left overnight, to remove organic carbon. Excess H2O2 was boiled off and the

samples were treated with sodium hexametaphosphate overnight to disperse clay-sized particles. Samples were added to the Mastersizer 2000 particle analyzer in ~800 ml of DI water using a plastic dropper or pipette; no particles over 2 mm were introduced to the

Mastersizer. Samples were drawn into the plastic dropper or pipette while using the Lab

Dancer vortex device to ensure the material was homogenous. The samples were analyzed on a program created by Dziekan (2017), in which the Mastersizer takes three separate readings from each sample and averages them. Detailed procedures for the

Mastersizer can be found in the user guide for the Mastersizer 2000. The data for each sample was exported into Microsoft Excel. Percentages based on Modified Wentworth particle sizes were calculated for sand (0.062–2.00 mm), silt (0.004–0.062 mm), and clay

(<0.0002–0.004 mm) and plotted on a ternary soil diagram from the USDA (2018).

Weight percentages for sediment > 2 mm were calculated separately.

2.9 Bathymetry Map of Clear Lake

The bathymetry map of Clear Lake was generated by collecting depth and geospatial data at 325 locations using a Garmin 64s GPS unit (x and y values) and an

Eagle FishEasy depth finder (z values). The xyz values from each location were exported as CSV files and imported in ArcMap as tables. The xyz data were converted to points and contours were drawn using ArcMap’s 3D Analyst tools. Contours were manually adjusted and superimposed over satellite imagery downloaded from the U.S. Geological

Survey (USGS) National Map Viewer Download Platform (2018).

2.10 Surficial Geology Map

A surficial geology map of the Burnside 7.5’ Quadrangle was developed using data from the U.S. Department of Agriculture (USDA) Web Soil Survey (WSS) (2016), the state of Michigan’s Wellogic water-well database (2018), geomorphology from

LiDAR-derived DEMs, and field work. Over 37,600 soil shapefiles from the WSS were uploaded to ArcMap. The shapefiles were color-coded based on the parent material, e.g. till, outwash, muck, associated with each soil-type. A feature class was created for each color-coded unit within the boundary of the Burnside Quadrangle, resulting in a preliminary surficial geology map. The feature classes were adjusted over the course of a year based on field work, interpretations of LiDAR DEMs, and boring data. The template for the final map was downloaded from the USGS’s National Map Viewer Download

Platform (2018), and included hydrologic and cultural features, as well as elevation contours.

2.11 Geologic Cross-Sections

Surface topography was extracted from the LiDAR-derived DEM in ArcMap using 3D analyst and exported to Microsoft Excel. Depths of lithologic units in water- wells from Michigan’s Wellogic database were converted to elevations in Excel and placed at appropriate distances along the profiles. When possible, water-wells that penetrated to bedrock and were within 150 m of the topographic profile were used.

Previously collected, unpublished geotechnical borings provided by Genesee County

(PSI, 2013), Michigan were used for the Clear Lake Road cross-section.

2.12 Ground Penetrating Radar

Ground penetrating radar (GPR) is useful for determining the structure and geometry of sand and gravel deposits like strandlines (Bristow and Jol, 2003). GPR transmits radio waves into the subsurface. When the transmitted energy encounters changes in electrical conductivity, e.g. due to particle size, a reflection returns to the receiver on the GPR instrument (Bristow and Jol, 2003). GPR surveys in the study area used 100 or 200 MHz Sensors and Software antennae. Step size was 0.25 m, antenna spacing was 1 m (100 MHz) to 0.75 m (200 MHz), and two-way velocity was estimated to be 0.06 nm/s. Particle size was verified at GPR sites using hand-dug or augured holes.

GPR data were processed and corrected for topography using data from the LiDAR- derived DEM of Lapeer County within the Sensors and Software computer program.

2.13 Bedrock Topography and Drift Thickness Map

A bedrock topography and drift thickness map were created in ArcMap using

Wellogic water-wells that penetrated to bedrock. The bedrock map was created by joining a table of geospatial points representing water-wells to a standalone table that included lithology and depth in ArcMap using the ‘WellID’ field. The “Extract Values to

Points” tool was used in ArcMap to add elevation data to the joined table from the DEM.

The resulting table contained lithology, depths, surface elevations of the wells, and geographic coordinates. The joined table was exported into Microsoft Excel and was manipulated so that the first bedrock surface was the only depth associated with each well. Depth to first bedrock surface was subtracted from each water-well’s surface

elevation in Excel. The Excel file was imported to ArcMap and converted to a point feature class. The “Geostatistical Wizard” was used to generate a continuous bedrock surface by interpolating the area between each water-well point with ordinary kriging, using the first bedrock elevation as the z-value. The drift thickness map was created in the same manner as the bedrock topography map, except that the depth values to bedrock were interpolated as the z-value rather than the elevation of a first bedrock surface.

2.14 Electrical Resistivity

Electrical resistivity surveys rely on the conductive properties of buried sediments to image the subsurface. Electricity travels more readily through conductive sediment, e.g. clay minerals, than through resistive sediments like sand, gravel or bedrock (Zhou et al., 2000). Resistivity surveys were set up using the Schlumberger array (Loke, 2001) with an AGI Super Sting. The Schlumberger survey design includes two current electrodes (A and B) on the outside and two potential electrodes (M and N) in the middle.

The current electrode (A-B) spacing began at 1 m for the first measurement from a central point and expanded laterally from the central point and expanded laterally in a logarithmic progression: 1.3, 1.6, 2, 2.5, 3.2, 4, 5, 6.3, 8, 10, 13, 16, 20, 25, 32, 40, 50,

63, 80 m, etc. The potential electrode (MN/2) spacing began with 0.2 m on each side about a central point, i.e. 0.4 m spacing between the M and N electrodes. The M-N spacing was moved to 1 m at AB/2 = 5 m; to 4 m at AB/2 = 20 m; and as needed beyond

40 m. More information about survey design and theory can be found in Loke (2001) and

Stierman (2011). Data were saved in the Super Sting and plotted on log-log graph paper.

Data were processed using IX1D software and compared to Wellogic water-well data for correlative stratigraphy.

2.15 Horizontal to Vertical Spectral Ratio Survey

Horizontal to vertical spectral ratio (HVSR) methods are being increasingly used to map depth to bedrock in previously glaciated areas. HVSR is a passive seismic method that relies on picking up resonant S-wave frequencies of unconsolidated sediments overlying hard bedrock (Chandler and Lively, 2016). The seismometer was pressed into bare ground in drainage ditches along roads to avoid trespassing on private land. The device was aligned to the north and ran for 16 minutes per site. After the reading the soil was checked to ensure the device was coupled well with the ground, which was determined by the impression of the rectangular base plate. Readings were collected only when wind speeds were less than 10 mph, as verified from local weather stations at the time of each reading. When possible, readings were taken close to Wellogic water-wells that penetrated to bedrock. Data were processed using Grilla software by undergraduate researcher Brian Samsen, following the methodology of Chandler and Lively (2016) and the Grilla user guide. Depth to bedrock data resulting from the HVSR survey was plotted on transects with bedrock depths calculated from water-well data.

2.16 Rebound Data

Strandlines associated with highest Lake Maumee, lowest Lake Maumee, and

Lake Whittlesey were traced in ArcGIS from an arbitrary starting point near Bryan, Ohio.

Strandlines were traced to the northeast following the methodology of Breckenridge

(2013) and Dilworth (2017). Strandlines for each lake stage were saved as separate polyline shapefiles and converted to points equally spaced at 50 m. The “Extract Values to Points” tool was used in ArcMap to extract elevations for each point from a LiDAR-

derived DEM provided by John Dilworth (personal communication, 2018). The “Snap” tool was used to align the points to a line perpendicular to isobases, i.e., the direction of maximum uplift (DMU), in the western ALE basin. Isobases were established by locating strandlines of equal elevation associated with each ALE stage on either side of the ALE basin, using maps generated by Forsyth (1959). The line representing the DMU in this study has a compass orientation of ~18o northeast, which is similar to the 23o orientation of the line representing the DMU in a study of isostatic rebound in the ALE basin

(Dilworth, 2017). The “Near” tool was used to determine the point spacing between each line after they were snapped to the line representing DMU. Datasets for the three stages of ALE were plotted against distance along the DMU. The 2nd order polynomial for each dataset was calculated in Excel. The y-intercept for each ALE dataset was selected by finding the values of each strandline, i.e. highest and lowest Maumee and Lake

Whittlesey, nearest the Ft. Wayne outlet. Also plotted were elevation data from outlets associated with Lake Maumee at Ft. Wayne, Indiana and in Lapeer County, Michigan.

Strandlines for highest Lake Maumee were adjusted for rebound by subtracting the highest and lowest values from each other, i.e. those from Ft. Wayne (y-intercept; 240 masl) from those near Imlay (260 masl). The elevation at Imlay was selected by creating an idealized line for data points associated with highest Lake Maumee, i.e. using the equation of its associated 2nd order polynomial: y=0.0002x2+0.0215x+240 with an R2 equal to 1; this was done so that it would be easier to select a data point from Imlay. The water plane for highest Maumee was set at 240 masl, with the assumption that no uplift occurred at the outlet at Ft. Wayne, and 20 m was subtracted from the sill at the Imlay outlet. Lowest Maumee and Whittlesey strandlines were flattened out as well, with

elevations relative to the outlet with closest proximity to each strandline, i.e. Ft. Wayne for lowest Maumee and the Imlay for Whittlesey.

Chapter 3

Results and Interpretations

3.1 Introduction

In an effort to gain a better understanding of the chronology and sedimentology of the study area, the main focus of which is the Imlay channel (Figure 3.1), numerous field sites were visited. Several new radiocarbon and optically stimulated luminescence

(OSL) ages came from these sites within the channel, as well as from Deanville Mountain and a channel running parallel to it. Field work, i.e. coring, hand-auguring, and outcrop observation, assisted in the generation of a digital surficial geologic map of the Burnside

7.5’ Quadrangle, as did the interpretation of water-well records and USDA Soil Survey data. Geophysics were used at several locations within the study area, including one where Leverett and Taylor (1915) identified a strandline of highest Lake Maumee.

Geographic information systems also helped to provide a better understanding of the study area’s bedrock geology and complex geomorphology.

3.2 Geomorphology

The study area is ~510 km2 and lies completely within Lapeer County, Michigan

(Figures 3.1 and 3.2). The Burnside Quadrangle (BQ) is ~141 km2 and is located in the northeastern portion of the study area. Many of the geomorphic features in the study area were identified by Leverett and Taylor (1915). However, the identification of new geomorphic features was aided by unpublished LiDAR data from Lapeer County, which

Figure 3.1. Labeled hillshade DEM of the study area in northeastern Lapeer County, Michigan.

Figure 3.2. Hillshade DEM of the study area in northeastern Lapeer County showing locations of close-up figures.

was only recently made available to the public (USGS, 2018). Geomorphic features include morainal ridges, channels and valleys, esker ridges, hummocky topography, streamlined hills, strandlines, and kettle lakes.

Located in the northwest portion of the study area (Figure 3.1), the Otisville

Moraine is an approximately 6.2 km long ridge oriented northeast-southwest. The North

Branch and Elm Creek channels are located partially within the northwestern section of the BQ. The North Branch channel is ~10 km long and runs east to west, while the ~11.5 km long Elm Creek channel is occupied by the underfit creek of the same name. About

16 km of the Deanville Moraine is located east of Deanville Mountain and is oriented north-south. West of the Deanville Moraine, a channel is oriented northwest-southeast along the eastern flank of Deanville Mountain, which rises steeply to the west (Figure

3.1), with gradients of ~5–8%, as calculated from the DEM. A previously unidentified morainal ridge sits on the northern edge of Deanville Mountain and is ~7.7 km long

(Figure 3.1). Another morainal ridge identified in this study is oriented northwest- southeast, in the broad, V-shaped valley between the Deanville Moraine and Deanville

Mountain, and into the lowlands near the Weaver Drain site (Figure 3.1). The Goodland

Moraine lies just south of the BQ and runs north-south for ~8.5 km. Two streamlined hills are cut into the Goodland Moraine (Figure 3.1); both hills taper to the west. The

Imlay Moraine is ~ 9.3 km long and strikes north-south as well (Figure 3.1). Two streamlined hills are cut into the Imlay Moraine, with the tapered end of each directed to the north.

Figure 3.3. Map (A) and oblique 3D model (B) illustrating alluvial fans in the Imlay channel. View is to the southeast on the 3D model in B. See figure 3.2 for location.

The modern surface expression of the Imlay channel is flat-bottomed and contains numerous alluvial fans, especially near its topographic divide (~250 masl), where it is flanked by highlands of the interlobate moraine to the west and Deanville

Mountain to the east (Figures 3.1 and 3.3). The Imlay channel runs for ~28.6 km within the study area and trends northwest-southeast until it reaches the topographic divide, near the intersection of Van Dyke Road (M-53) and Clear Lake Road, where it turns southward (Figures 3.1 and 3.3). Approximately 14.8 km south of the topographic divide, the channel is ~1,500 m wide and has an elevation of ~244 masl, but constricts to a width of ~900 m about 2.5 km to the north (Figure 3.1). The channel maintains a width of ~900 m for about 11 km to the north, where it is bounded by the interlobate moraine in the west and the Imlay Moraine in the east (Figure 3.1). At the southern boundary of the BQ,

~4 km south of the topographic divide, the channel is ~900 m wide and has an elevation of ~242 masl. Here, the channel is flanked to the east and west by Deanville Mountain and the interlobate moraine, respectively, which both rise to elevations >275 masl. About

2 km northwest of the topographic divide a ~3 km long valley, starting at Clear Lake in the southwest, enters the Imlay channel (Figures 3.1 and 3.3). Like the Imlay channel, several alluvial fans spread out into the valley leading to Clear Lake (Figure 3.3). The

Imlay channel maintains a width of <1,000 m for the remainder of its course through the study area. The Imlay channel intersects the Elm Creek channel ~10.3 km northwest of the topographic divide and the North Branch channel ~3 km further to the northwest.

Figure 3.4. The interlobate moraine to the west of the Imlay channel contains numerous kettle lakes, eskers, and tunnel channels. See figure 3.2 for location.

The interlobate moraine is the area of land between the Huron-Erie and Saginaw lobes of the Laurentide Ice Sheet (LIS). Like other interlobate areas (Carlson et al., 2004) the geomorphology is complex, with numerous eskers, channels and valleys, and kettle lakes (Figure 3.4). A prominent feature within the interlobate moraine is the Lum channel, which is oriented northwest to southeast and is ~16 km long (Figure 3.1). The channel’s margins are ~2 km wide, before they become indistinguishable ~2 km beyond the study area’s northwest boundary. The topographic divide of the Lum channel is ~270 masl, which is located ~3 km southwest of Clear Lake (Figure 3.1). As noted by Leverett

and Taylor (1915), considering that this elevation is ~10 m higher than the highest

Maumee strandline near the Goodland Church (Figure 3.1), it is unlikely that the Lum channel was able to accommodate water from any stage of Lake Maumee, but may be a meltwater (or ice marginal ?) channel from a previous glacial advance. The strandline at the Goodland Church is ~260 masl and sits atop a ridge composed of till. Strandlines associated with Lake Whittlesey pass to the south of the mouth of the Imlay channel, ~30 km south of the topographic divide (Figure 1.7) and appear to build off the far eastern arm of the Deanville Moraine ~35 km southeast of the study area (Figure 3.5).

Strandlines associated with Lake Whittlesey north of this area were not identified due to lack of LiDAR data.

3.3 Imlay Channel and Clear Lake Sedimentary Sequences

Four coring sites and one hand-dug pit were taken within ~4 km of the modern drainage divide of the Imlay channel (Figure 3.1). Three of the coring sites had similar sedimentary units that were classified into five sedimentary facies (Figure 3.6).

Facies Descriptions. Facies A (Figure 3.6) is mostly massive gravelly sand, but in some cases it is weakly bedded. Munsell colors are gray (10YR 6/1) and light brownish- gray (10YR 6/2) when unoxidized. When oxidized, the facies is pale yellow (2.5Y 7/3) and pale brown (10YR 7/3). Lower contacts are abrupt and sometimes erosional. Sand in facies A is fine to very coarse grained, subangular to subrounded, and is comprised of quartz grains, crystalline lithics (i.e. igneous/metamorphic), and plates of mica (i.e. muscovite and biotite). Gravel is very fine to fine pebbles, subangular, and comprised mostly of sedimentary (i.e. sandstone/shale/limestone/dolomite) and crystalline clasts.

Figure 3.5. Non-LiDAR DEM (A) showing two spits associated with Lake Whittlesey building off the eastern arm of the Goodland Moraine. A LiDAR-derived DEM (B) shows the northern-most spit identified and strandlines of lower lake levels. Both images are “flooded” (blue) to an elevation of 233 m.

Figure 3.6. Sedimentary facies and legend for cores taken within the Imlay channel: CORICA, CORICB, and CORICC. Legend is also relevant for other figures, where noted in the caption.

The depositional environment of facies A is interpreted as a fluvial environment based on facies descriptions characterized by Miall (1977) and Rust and Koster (1984).

Facies B (Figure 3.6) consists of alternating beds and laminations of silty sand and gravelly sand, the latter similar to facies A. The silty sand is light gray (10YR7/1), massive and has an abrupt flat, and sometimes erosional, lower contact with the gravelly sand. Within the facies, beds and laminations of gravelly sand are massive and have erosional lower contacts with the silty sand. The depositional environment of facies B is interpreted as being one of moderate energy, possibly near the mouth of a river or stream in a shallow lake or pond (Miall, 1977; Rust and Koster, 1984). Shallow lakes and/or ponds could form in abandoned fluvial channels (Miall, 1977; Rust and Koster, 1984;

Nanson and Croke, 1992), or temporary lakes may form due to alluvial fans intermittingly damming water drainage routes in the Imlay channel creating temporary,

shallow lakes and/or ponds as they build toward the center of a valley or channel (Figure

1.9).

Facies C (Figure 3.6) is a sandy gravel and is similar to facies A, especially the sand component, except the gravel is more abundant and coarser than that of facies A, i.e. very fine to coarse pebbles. The facies is pale yellow (2.5Y 7/8), massive, pebbles are subangular to subrounded and contain sedimentary and crystalline clasts. The facies is found in only one core and has an abrupt lower contact. A depositional environment for facies C is interpreted to be fluvial (Miall, 1977; Rust and Koster, 1984).

Facies D (Figure 3.6) is a massive gravelly and sandy mud found in one core.

Sand is the same as facies A and is supported in a matrix of silt and clay. Gravel is subrounded, very fine to coarse pebbles, is composed of sedimentary and crystalline clasts, and the lower contact is abrupt and angular. The depositional environment of facies D is interpreted as one of ponded water, possibly fill of a scour or abandoned channel due to flooding (Miall, 1977; Rust and Koster, 1984; Nanson and Croke, 1992).

Facies E (Figure 3.6) is fibrous peat with abrupt and gradational lower contacts.

In many cases the water table is ~0.30 m below the upper surface of the peat.

Topographic Divide. A vibracore, CORICC, was taken in a bog at the Imlay channel’s topographic divide (Figure 3.7). The 3.2 m long core is composed primarily of facies A and B and is capped with 0.7 m of facies E. Facies A is weakly bedded at ~2 m depth. Visible organics in the upper portion of facies A appear to increase upwards, suggesting facies E has gradational lower contact with facies A. However, the lower

Figure 3.7. Field location (A) of CORICC (B) at Imlay channel’s topographic divide. Refer to Figure 3.5 for facies description.

contact of facies E was not recovered as a result of hand auguring too far into the peat in preparation for vibracoring. Interbedded facies B within facies A suggests a brief period of deposition within a shallow lacustrine (facies B) environment, departing from a fluvial environment (facies A). Weak bedding within facies A suggests fluctuation in energy as it was being deposited, as do rip-up clasts of facies B within it. An OSL age (Table 3.1) from >3.09 m below the ground surface within facies A indicates deposition by 14.1 ±

1.4 ka.

Northwest of Topographic Divide. A 4.4 m long vibracore, CORICB (Figure 3.8), was taken in a bog within the Imlay channel ~1.5 km northwest of the topographic divide. There are two sedimentary facies present in the core, including two sequences of facies B which is interbedded with one sequence of facies A (Figure 3.8). Overlying the uppermost sequence of facies B is likely facies E. As with CORICC (Figure 3.7), facies B in CORICB is interpreted as representing a shallow lacustrine depositional environment, while facies A is interpreted as a fluvial depositional environment. An OSL age from sediment at a depth of 4.25 m yields an age of 14.9 ± 1.4 ka (Table 3.1).

Southeast of Topographic Divide. Vibracore CORIC1A (Figure 3.9) was taken in the Imlay channel ~1 km southeast of the topographic divide. The 1.3 m long core consists of four distinct, interbedded sedimentary facies. The lower contact of facies A is buried. The erosional lower contact and the larger gravel component of facies C compared to facies A suggests a sudden increase in the amount energy in the depositional environment. The abrupt lower contact of facies D suggests energy decreased suddenly.

The lower contact of facies E is irregular and gradational, with macrofossil content increasing upward. A radiocarbon age of 11,050 ± 35 14C yrs. (12.8–13.0 kcal yrs. BP)

Figure 3.8. Field location (A) of CORICB (B), ~1.5 km northwest of the Imlay channel’s topographic divide. Refer to Figure 3.5 for facies description.

Table 3.1. OSL ages. Sample ID Age (ka)* Error (ka) † Year Stratigraphy Depth Collected (facies)† † (m) Laminated sand underlying massive OSLBSGA 70.3 6.9 2017 1.86 sand & gravel on proximal end of kame deposit Sand underlying OSLBSGB 60.6 6.0 2017 sand & gravel on 10.2 distal end of kame deposit Bedded gravelly OSLCORICB 14.9 1.4 2017 sand & laminated 4.25 silty sand (Facies B) Bedded gravelly sand & laminated OSLCORICC 14.1 1.4 2017 3.09 silty sand underlying peat (Facies A) Sand underlying OSLJTEA 10.2 0.9 2017 1.05 gravelly sand in an alluvial fan deposit *Lab data (i.e. ages and chemistry) can be found in Appendix A. †Fully-propagated age uncertainty (Append. B, Aitken, 1985). ††If applicable. See Figure 3.5 for facies description.

(Table 3.2) from wood collected in the uppermost portion of facies D (Figure 3.9B and

C) is a maximum age for the clastic sediment of facies A, C, and D. A radiocarbon age of

10,250 ± 35 14C yrs. (11.8–12.1 kcal yrs. BP) from the basal peat indicates suggests that the peat was forming by this time. The two ages imply that the transition from a fluvial to a paludal environment occurred ~13 thousand years ago.

Alluvial Fan. An OSL age was collected from an alluvial fan deposit along the southwestern edge of Deanville Mountain at site JTEA, ~700 m east of CORICC at the topographic divide (Figure 3.10A). The OSL sample was collected from a unit of sand with an exposed thickness of 0.50 m. Sand was mostly very fine to medium grained, subangular to subrounded, with crystalline and quartz grains. Overlying the sand is ~0.75 m of sand and gravel, with topsoil developing within the uppermost ~0.35 m of the deposit (Figures 3.10B and C). An OSL sample from a depth of ~1.05 m yielded an age of 10.2 ± 0.9 ka (Table 3.1). This age suggests that alluvium from Deanville Mountain

Figure 3.9. Field location (A) of CORICA (B), ~1 km southeast of the topographic divide. Organic material (C) sent for radiocarbon dating. Refer to Figure 3.5 for facies description.

Table 3.2. Radiocarbon ages from the study area in Lapeer County, Michigan. 2 Sigma Lab 14C Age Year Stratigraphy Depth Sample ID Material Δ13C Calibrated ID Age Error Collected (Facies)† (m)‡ Age* Laminated marl & CORCL1B- OS- Wood/ Not 12.7 - 12.8 5.75- 10,850 55 2017 sapropel 1 137075 Plant Provided kcal BP 5.85‡ overlying sand & gravel Sand and gravel below CORCL1B- OS- Wood/ Not 14.1 - 14.9 6.61 - 12,400 65 2017 laminated 2 137076 Plant Provided kcal BP 7.32‡ marl & sapropel Sand and gravel below CORCL1B- Beta- Wood/ 13.8 - 14.1 6.61 - 12,090 40 -26.2 2017 laminated 3 491846 Plant kcal BP 7.33‡ marl & sapropel Basal laminated CORCL1B- Beta- Wood/ 14.0 - 14.3 marl & 6.49 - 12,250 40 -25.8 2017 4 491847 Plant kcal BP sapropel 6.61‡ overlying sand & gravel Basal peat just CORICA1- OS- Wood/ 11.8 - 12.1 overlying above 10,250 35 -24.8 2017 1 135589 Plant kcal BP gravelly mud 0.83 (Facies E) m Uppermost just pebbly mud CORICA1- OS- 12.8 - 13.0 below Wood/Plant 11,050 35 -25.63 2017 underlying 2 135590 kcal BP 0.83 peat m (Facies D) Diamicton (debris CORSL2A- OS- Not 14.0 - 14.6 flow?) Wood/Plant 12,300 55 2018 2.52 1 139410 Provided kcal BP underlying mud & organics Diamicton Beta- CORSL2A- 16.6 – 17.0 underlying 492602 Wood/Plant 13,870 40 -25.1 2018 2.52 3 kcal BP mud &

organics

Diamicton CORSL2A- Beta- 13.1 – 13.4 Wood/Plant 11,420 40 -24.3 2018 underlying 2.41 4 496778 kcal BP marl

Silt 2.48 CORSL2A- Beta- 14.2 – 14.8 underlying Wood/Plant 12,420 40 -24.3 2018 5 496779 kcal BP and overlying

diamicton *Radiocarbon ages were converted to calibrated ages using Calib V7.10 using the IntCal 13 calibration curve. †Where applicable. See Figure 3.5 for facies descriptions. ‡Depths from below the lake bed in Clear Lake with a water depth of ~8 m.

Figure 3.10. Field location (A) of OSLJTEA (B and C), ~700 m east of the topographic divide. Photo facing north. Colored lines distinguish unit boundaries. continued to be deposited into the Holocene, ~10–11 thousand years ago, and implies that alluvial fans occupying the uppermost elevations of the channel stabilized ~1–3 thousand years after peat (Figure 3.8) began to accumulate in the valley occupied by the Imlay channel.

Clear Lake. Clear Lake (Figures 3.1 and 3.11) is privately owned and was without a bathymetric map prior to this study. Clear Lake appears in plat books for Lapeer

County as early as 1874 and is likely one of the only natural lakes within the boundaries of the BQ. The lake appears to be primarily spring-fed, which is supported by the observation of artesian wells near the lake. One creek does enter the lake; however, aerial imagery suggests it originates from a spring ~1 km to the southeast of its mouth. Clear

Lake has one outlet, Cedar Creek, which flows northeast toward the Imlay channel. The underfit creek runs within the Imlay channel until ~14 km to the northwest, where it widens and becomes known as the North Branch of the Flint River. The lake lies in a depression at the head of a valley ~3 km southwest of the Imlay channel, which connects to it (Figures 3.1 and 3.12A). Clear Lake has an elevation of ~252 masl and is elongated in an east-west direction (Figures 3.11 and 3.12A). The western half of the lake is deeper than its eastern side with a maximum depth of 45 feet (~13.7 m).

Core CORCL1B (Figure 3.12) was taken with a Livingstone corer with hydraulic assist in the eastern section of Clear Lake. Coring was halted upon refusal at a depth of

~8 m below the lake bed, which resulted in several bent thrust rods. The core consists of two recovered units. Unit a is gray (1GLEY 6/N), normally graded (i.e. fining upwards) gravelly sand. Pebble sized gravel is observed most abundantly in the lower portion of unit a, is very fine to very coarse, and is of sedimentary and crystalline lithology. Sand within unit a is fine to very coarse grained, subangular to rounded, and consists of quartz and crystalline grains. Unit b is a shelly marl, i.e. lime mud with subordinate clastic material (Cohen, 2003). Unit b is interbedded and laminated with sapropel, and is

Figure 3.11. Bathymetry map of Clear Lake in Lapeer County, Michigan.

estimated to be ~6.62 m thick, although the majority of it was not collected. The marl is light grey (1GLEY 7/N) and the sapropel is greenish black (1GLEY 2.5/10GY).

Radiocarbon ages (Figure 3.12B, C; Table 3.2) from terrestrial plant matter in unit a are 12,400 ± 65 14C yrs. (14.1–14.9 calibrated ka BP) and 12,090 ± 40 14C yrs. (13.8–

14.1 calibrated ka BP). Woody material near the base of unit b was dated at 12,250 ± 40

14C yrs. (14.0–14.3 calibrated ka. BP). Pine (spruce?) needle fragments in unit b at a depth of ~5.8 m yielded an age of 10,850 14C ± 55 yrs. (12.7–12.8 calibrated ka BP).

Stratigraphy and chronology of CORCL1B (Figure 3.12) indicates that sand and gravel of unit a was deposited sometime after ~14 ka, likely in a fluvial environment as indicated by the stratigraphy of alluvial fans which grade into the Clear Lake valley (Figure 3.3). A radiocarbon age from organics at the base of unit b is about the same age as organics from unit a and suggests water rose in the Clear Lake basin shortly after deposition of unit a, i.e. ~14 ka, becoming a lacustrine depositional environment. Sometime before

~12.8 ka Clear Lake was a biologically productive lake.

3.4 Deanville Mountain, Weaver Drain, and Soil Texture

Deanville Mountain (Figures 3.1 and 3.13) was identified by Leverett and Taylor in 1915, who interpreted the features in the area as kame deposits. At least four gravel pits currently operate in the area, and are located ~1.6 km southwest of the Weaver Drain site (Figure 3.1). Access to most walls within the Bechtel pit (BSG) was not provided because they were too steep, therefore observations at BSG rely primarily on photographs. Sedimentary beds within the pit are either clast-supported sand and gravel or mostly sand (Figure 3.14C). Faults (Figure 3.14A, B, C) and deformation features, i.e.

Figure 3.12. Field location (A) of CORCL1B (B), ~3.5 km southwest of the topographic divide. Organic material (C) sent for radiocarbon dating. See Figure 3.6 for legend of core log (B). Water depth is 8.5 m.

load structures (Anketell et al., 1970) (Figure 3.13D), suggest that the deposit was in an ice-marginal environment (McDonald and Shilts, 1973). Faults strike approximately east- west (Figure 3.14A and B) and northwest-southeast (Figure 3.14 C), roughly perpendicular to the dip of the beds. Much of the pit is capped with a ~2 m thick matrix- supported gravel, i.e. diamicton (Figures 3.14A, B and 3.15).

Bechtal Sand and Gravel. An OSL sample (OSLBSGA; Figure 3.16) was taken at the top, proximal end of the kame on an upper ledge where a few meters of sediment have been removed (Figure 3.14A). The exposure (Figure 3.16) contains a deposit of clean, laminated sand underlying a bed of sand and gravel where soil development occurs. The sample was taken from the laminated sand with an exposed thickness of

~0.40 m. The sand was very fine to very coarse grained, subangular to rounded, and contains quartz and crystalline grains. The laminations are horizontal and some thin bedding, i.e. >10 mm, is observed at the outcrop. The overlying sand and gravel is ~1.29 m thick, has an erosional lower contact, and is comprised of clast supported gravel and sand. Sand is the same as that found within the underlying laminated sand. Very fine pebbles to small cobbles are subangular, and comprised of crystalline and sedimentary clasts. Bedding is horizontal and appears to be normally graded. Topsoil development is

~0.30 m thick and occurs in the uppermost sand and gravel. The OSL sample resulted in an age of 70.3 ± 6.9 ka. (Figure 3.16; Table 3.1). Diamicton is observed both overlying and underlying the sand and gravel (Figure 3.15) on the ledge at a site ~88 m north of where the OSL age was collected (Figure 3.14). This exposure shows the till extending

~0.70 m below the topsoil, which appears to be developed within clast-supported sand and gravel.

Figure 3.13. A DEM illustrates the Weaver Drain, alluvial fans filling a channel adjacent to Deanville Mountain, the Bechtal Sand and Gravel pit, and a previously unidentified morainal ridge between the channel and the Deanville Moraine to the east. See figure 3.2 for location.

arrow

D gure facing is west. to southwest

facesand southeast, the Fi

OSLBSGA was collected, and the blue arrow (A) approximates site with good exposure and approximates OSLBSGA(A) blue the arrow was collected,

Figure 3.14. Bechtal gravel pit facing east (A). The red arrow points to diamicton (A and (A B), yellow points arrow Bechtal 3.14. diamicton The red to pit gravel facing (A). Figure east site the where to (A) faults and (C) micro was Faultsan (B), where collected. OSLBSGA unit of north ofm ~88 diamicton C (D). load i.e. structures Figure deformation, tall. in and bar and m ~1 is D red (A B) arrow diamicton points Thered to

Figure 3.15. Location within Bechtal Sand and Gravel, ~88 m north of OSLBSGA (Figure 3.16) pit showing diamicton interpreted as till overlying sand and gravel deposits. Photo is facing east and red arrows point to larger clasts. Colored lines distinguish unit boundaries. The shovel on the right is ~ 1 m tall.

Figure 3.16. Field location (A) of OSLBSGA (B and C), ~1.6 km southwest the Weaver Drain site and ~500 m east of BSGB. Photo (C) is facing approximately to the south and shovel is ~1 m tall. Colored lines distinguish unit boundaries.

Diamicton (Figures 3.14 and 3.15) is interpreted as till due its lateral continuity throughout the pit at high elevations along the property line where the ground has not been excavated (Figure 3.14A). USDA Soil Survey data also indicates that the soil overlying the BSG pit is from a loamy till, which will be discussed later in this chapter.

The erosional lower contact of the diamicton, as well as the overlying sand and gravel

(glaciofluvial), also supports its interpretation as a till. The old age of OSLBSGA (Table

3.1) could indicate that the deposit at BSG predates late Wisconsinan glaciation.

However, the age may represent a sample that was insufficiently bleached from lack of exposure to sunlight due to a number of factors, including: rapid transport and deposition of the sediment at an ice margin or insufficient amount of light penetrating the water column, possibly due to the turbidity of the water near the glacier’s margin (Rittenour,

2008).

A second OSL sample (OSLBSGB; Figure 3.17) was taken at BSG, but at the distal end of the kame ~500 m west of the site where OSLBSGA was collected (Figure

3.16). About 1 m of cross-bedded sand and gravel is exposed along a lower face of the pit. Above the cross-bedded sand and gravel is ~0.90 m of clean, horizontally bedded and laminated sand. Sand is very fine to very coarse grained, subangular to subrounded, and is comprised of quartz and crystalline grains. About 9.7 m of sand and gravel, estimated from photos and the DEM, overlies the clean sand. The OSL sample was collected from the bedded and laminated sand unit (Figure 3.17) and resulted in an age of 60.6 m ± 6.0 ka.

Figure 3.17. Field location (A) of OSLBSGB (B and C), ~1.6 km southwest the Weaver Drain site and ~500 m west of BSGA. Photo (C) is facing approximately to the southeast and shovel is ~1 m tall. Colored lines distinguish unit boundaries.

Evidence that the outcrop at the Bechtal pit is a kame deposit comes from observations of faults and deformation (load) structures (Figure 3.14). A diamicton interpreted as till caps most of the pit at high elevations (Figures 3.14 and 3.15), except where OSLBSGA was collected (Figure 3.16). Since the till appears to laterally continuous south of where the OSLBSGA was collected (Figure 3.14) and at least ~88 m to the north (Figures 3.14 and 3.15), one possibility is that the till was eroded away.

Indeed, several gullies on the steep eastern slope lead to a channel flanking Deanville

Mountain in the east (Figure 3.13). Another possibility for the absence of till at the site where OSLBSGA was collected is that the ice margin never reached this far. The OSL ages from the samples collected at OSLBSGA and OSLBSGB suggests that the kame was deposited before late Wisconsinan glaciation occurred or, possibly, that sediment has been insufficiently bleached because it was deposited under the ice, in turbid waters, or due to rapid deposition and burial at the ice margin (Rittenour, 2008).

Channel. A channel runs along the eastern flank of Deanville Mountain, ~2 km south of the Weaver Drain site (Burgis, 1970) (Figure 3.13). In plan-view, the sigmoidal shaped feature is ~4 km long and strikes approximately northwest-southeast. East of the channel is a morainal ridge, which is ~2.5 km long and has a maximum elevation of about 260 masl (Figure 3.13). Topographic profiles from the DEM indicate the channel is flat along its longitudinal axis in the north for the first 460 m. The channel rises to a maximum elevation of ~270 masl about 2 km from its northern end and begins to slope to the southeast with a grade of ~0.4%. However, the DEM shows several alluvial fans entering the channel from the highlands that make up Deanville Mountain (Figures 3.1 and 3.13), making the interpretation of paleowater flow direction difficult.

A vibracore, CORSL2A (Figure 3.18), was taken in a bog ~1,400 m southeast of the channel’s northernmost end, where the feature was easily accessible from the road. The purpose of coring this site was to collect material for radiocarbon dating. The vibracore contains three units. Unit a has a recovered thickness of ~0.20 m (Figure 3.18B), is a fossiliferous diamicton, and is classified as a loamy sand on a USDA soil texture diagram

(Figure 3.19, sample 5). Unit a contains very fine to coarse pebbles and is bluish gray

(2GLEY 5/5B). Unit b is 0.06 m thick fossiliferous silt and plots as a silt loam on the

USDA soil texture diagram (Figure 3.19, sample 12) and is dark bluish gray (2GLEY

4/5B). However, the Munsell color for unit b was not checked until after several months of storage in a 4 oC cooler, and may have changed since CORSL2A was first opened. An approximately 0.25 m thick sequence of unit a overlies unit b. Unit c is 0.39 m thick, is comprised of laminated marl and sapropel, and has a gradational lower contact. Light laminations in Unit c are white (2.5YR 8/1) and dark lamina are reddish gray (2.5YR

6/1). Unit d is ~1.81 m thick peat interbedded with unit c in its lowermost section.

Organics collected from unit a at a depth of ~2.52 m yielded ages of 12,300 ± 55 14C yrs.

BP (14.0–14.6 calibrated ka BP) and 13,900 ± 40 yrs. 14C yrs. BP (16.7–17.0 calibrated ka BP) (Figure 3.18B and C; Table 3.2). A twig from unit b at a depth of ~2.47 m yielded a radiocarbon age of 12,400 ± 40 14C yrs. BP (14.1–14.8 calibrated ka BP) and a pine cone from the uppermost sequence of unit a at a depth of ~2.40 m, resulted in an age of

11,400 ± 40 14C yrs. BP (13.1–13.3 calibrated ka BP). Diamicton from unit a is interpreted as colluvium. Evidence that the material in unit a is colluvium from a debris flow rather than a glacial deposit (i.e. till) is based on the two radiocarbon ages in the lowest section of unit a, both at a depth of ~2.52 m but with an age difference of ~2,500

Figure 3.18. Field location (A) of CORSL2A (B), ~2 km south of Burgis’ (1970)

Weaver Drain site. Organic material (C) that was radiocarbon dated. See Figure 3.6 for legend of core log (B).

rich rich

es collected from the Burnside Quadrangle. USDA soil texture diagram diagram Burnside the (B)from soil es texture Quadrangle. collected USDA

plotted and average textures classifications of Burgis’ (1970) sandstone plotted(1970) classificationsBurgis’ textures of and average

richtills.

. Location (A) of sediment sampl Location . of sediment (A)

(USDA, 2018) with 12 sediment samples samples with 2018) sediment 12 (USDA, andlimestone 3.1 Figure

calibrated years. Also, the diamicton from unit a plots as loamy sand on the USDA soil chart (sample 5; Figure 3.19), while most of the till in the area plotted as silt or clay loam.

The steep eastern slope of Deanville Mountain adjacent to the channel provides further evidence that till capping the area may have slumped into the channel as a debris flow

(Figure 3.13). The debris flow which deposited the colluvium occurred at least twice before ~13 ka, i.e. the age from the uppermost part of unit a (Figure 3.18). The oldest age in the core (16.6–17.0 calibrated ka BP) is similar to the age Burgis (1970) (16.1–17.3 calibrated ka BP) from material collected in the Imlay channel. In this study, the oldest age from CORSL2A is interpreted to represent a minimum age for deglaciation in the area, since it likely came from till deposited on the western edge of Deanville Mountain prior to the debris flow that deposited the diamicton in the channel.

Weaver Drain. The Weaver Drain is an engineered ditch that begins its course in the highlands between Deanville Mountain in the west and the Deanville Moraine in the east (Figure 3.13). The site where Burgis (1970) collected organics for radiocarbon dating is described as being located ~1 km northwest of BSG on a hummocky plain just east of a morainal ridge that was identified on the DEM (Figure 3.13). Two field sites within ~500 m of the location identified by Burgis were visited in this study and neither displayed the stratigraphy discussed by Burgis (1970). Hand auger samples were collected at two other sites near the Weaver Drain for particle size analysis, plotting as silt loam (sample 6) and sandy loam (sample 11) (Figure 3.19). A shovel was used to clear debris from the edge of the Weaver Drain at the first site, WDB (Figure 3.20). The lowermost sand and gravel were encountered at a depth of ~1.40 m. Some silt and sand were recovered below 1.40 m as well, but the presence of the water table at the top of the uppermost sand and gravel

layer made the recovery and depth determinations of these finer sediments difficult.

Overlying the sand and gravel is ~0.70 m of mud which plotted as a silt loam on the

USDA particle size chart (Sample 3; Figure 3.17). About 0.70 m of shelly marl overlies the mud, with the uppermost ~0.51 m of the marl is affected by soil development.

The second site, WDE (Figure 3.21), was encountered serendipitously when roadside ditches in the area were being cleared with backhoes. The lowermost sediment exposed

~0.20 m is diamicton (Figure 3.21B and C) and plotted as a silt loam on the USDA soil texture diagram (Sample 8; Figure 3.21). The diamicton was laterally continuous for at least several hundred meters west of WDE along the road. Overlying the diamicton is

~0.70 m of mud, the uppermost ~0.50 m of which was affected by soil development. The diamicton is interpreted as a till due to its lateral continuity, the hummocky topography at the site, and its proximity to the morainal ridge (Figures 3.11 and 3.21A). The V-shaped valley now occupied by the Weaver Drain would have likely been an area of ponded water if an ice dam blocked water drainage to the north and northwest. Soil survey data indicates that sediment within the valley is composed of glacial, glaciofluvial, glaciolacustrine and undifferentiated glaciofluvial and glaciolacustrine deposits. It is possible that the stratigraphy described by Burgis (1970), i.e. “beds of unoxidized clay, silt, sand, and plant macrofossils overlie sandstone till…,” is representative of glaciofluvial and/or glaciolacustrine deposits overlying till in an ice-dammed depression in the valley now occupied by the Weaver Drain rather than an outlet channel for the middle stage of Lake Maumee (Maumee III).

Figure 3.20. Field location (A) of WDB (B and C), ~500 m southeast the Weaver Drain site. Photo (C) is facing east and blue arrow points to the water table. Colored lines distinguish unit boundaries.

Figure 3.21. Field location (A) of WDE (B and C), ~500 m southwest of the Weaver Drain site. Photo (C) is facing south. Colored lines distinguish unit boundaries.

Soil texture. An attempt was made to compare soil texture from 12 sites within the

Burnside Quadrangle to Burgis’ average soil texture for the limestone-rich till, plotted as loam (sand: 45%, silt: 34%, clay: 21%) on a USDA soil texture diagram (Figure 3.19).

Also plotted on the USDA diagram is the average soil texture for sandstone-rich till

(Burgis, 1970), plotted as clay loam (sand: 29%; silt: 40%; clay: 31%) (Figure 3.19).

Eight samples from this study, i.e. samples 1, 2, 3, 4, 6, 7, 8, 12, plotted as silt loam. The remaining samples plotted as follows: silty clay loam (sample 9), clay loam (sample 10), sandy loam (sample 11), and loamy sand (sample 5). Sample 10 plotted with the average texture discussed by Burgis for limestone-rich till, and none plotted with the texture for sandstone-rich till. Furthermore, only one of the samples in this study correlate to the

USDA soil textures associated with each sampling location, i.e. clay loam or loam. All data from the particle size analysis can be found in the table in Appendix C.

The particle size analysis in this study was useful in demonstrating how the diamicton from the channel east of Deanville Mountain in the core CORSL2A (Figure

3.18) differed from other diamicton interpreted as till by the USDA Soil Survey at other sites. However, a direct comparison between samples collected in this study and those described by soil survey were not possible. One reason for this may be that the USDA uses the hydrometer method (HM) or pipette method for particle size analysis on grains

<2 mm (Gee and Or, 2000; Burt, 2011), while this study uses the laser diffraction method

(LDM). A study on particle size analysis (Stefano and Mirabile, 2010) directly compared the results of 228 soil samples using the HM and the LDM. Stefano and Mirabile (2010) found that while sand percentages were similar using either the HM or the LDM, clay distribution was overestimated for the HM, by >20% in some cases, relative to the LDM.

Burgis (1970) also used the hydrometer method to determine particle size for the limestone-rich (i.e. west of Imlay channel) and sandstone-rich (i.e. east of Imlay channel) tills, which may also explain why the soil texture results from this study differed from

her’s. Of course, only one sample of till was collected east of the Imlay channel (sample

10; Figure 3.19) for soil texture analysis in this study, which may be another reason why there appears to be a lack of correlation between soil texture results between the two studies.

3.5 Bedrock

Bedrock within the study area includes three formations which are all

Mississippian in age: the Coldwater Shale (5.7% of area), the Marshall Sandstone

Formation (83.7%), and the Michigan Formation (10.6%) (Figure 3.22B; Table 3.3)

(Westjohn and Weaver, 1996; Michigan DEQ, 1999). The oldest bedrock unit, the

Coldwater Shale, is up to ~390 m thick in eastern Michigan, and is comprised of shale, sandstone, siltstone, and carbonate units. Locally, the ~61 m thick Marshall Sandstone

Formation is comprised of sandstone, shale, siltstone, and carbonate units and overlies the Coldwater Shale (Monnet, 1948; Westjohn and Weaver, 1996). The hydraulically connected sandstone units within the Marshall Sandstone Formation form a major regional aquifer, i.e. the Marshall Aquifer, with an areal extent of ~52,000 km2. The upper confining unit for the Marshall Aquifer is the Michigan Formation, which is comprised of shale, limestone, dolomite, gypsum, siltstone, and sandstone; units and can be as thick as ~122 m (Westjohn and Weaver, 1996).

Maps. Bedrock topography (Figure 3.22B) and drift thickness (Figure 3.22C) maps were generated using ArcGIS. The long axis of a semicircular bedrock ridge trends approximately N-S through the center of the BQ (Figure 3.22B). The ridge underlies

Deanville Mountain in the center of the quadrangle. Another bedrock high is located in

the southwestern corner of the study area. Low bedrock elevations in the southwest corner of the BQ appear to underlie the area of the Lum channel (Figure 3.1). Another bedrock low occurs in the northeastern portion of the study area and runs south along the eastern side of the BQ. Unconsolidated sediment (i.e. drift) is generally thickest in areas of lower bedrock topography and thinnest in areas of high bedrock topography (Figure

3.22C). Bedrock surfaces for the 947 wells in the study area (Table 3.3) were shale

(41%), sandstone (39%), sandstone and shale undifferentiated (14%) (Table 3.3), and 6% of the wells indicated a first bedrock surface of either limestone or undifferentiated limestone and/or sandstone and/or shale. Minimum and maximum thicknesses of unconsolidated sediments within the study area were 8.8 m and 83.5 m, respectively, with an average thickness of 40 m (Figure 3.22C; Table 3.3). Lithology of first bedrock surfaces from 212 wells within the Burnside Quadrangle (Table 3.3) followed a similar trend: shale (47%), sandstone (33%), sandstone and shale undifferentiated (16%), and undifferentiated limestone (4%) (Table 3.3). Minimum and maximum thicknesses of unconsolidated sediment within the quadrangle were 12.2 m (40 ft.) and 73.2 m (240 ft.), respectively, with an average thickness of 33 m (111 ft.) (Table

3.3).

The bedrock topography does not suggest that the process that cut the Imlay channel affected the underlying basement material to any great degree (B-B’; Figure

2.22B). Of course, another reason why the bedrock appears unaffected below the Imlay channel in the interpolated bedrock surface may be that few water-wells penetrate to bedrock within the area occupied by the channel (Figure 3.22A). In contrast, bedrock below the Lum channel appears to mirror its surficial expression. Drift (Figure 3.22C)

in the study area and in study the area

sections

(B), and drift thickness (C) (C) and (B), drift thickness

bedrock used to create the bedrock and drift thickness maps. Also shown are shown Also maps. bedrock the are and bedrockthickness drift used create to

wells that penetrate to to wellspenetrate that

Figure 3.22. 3.22. topography of Maps bedrock topography Figure surficial (A), water of locations 3.28, 3.27, cross and Lake (Figures (3.30) 3.29) of locations longitudinal the Clear the

Table 3.3. Bedrock composition and drift thickness. Burnside Study Area Quadrangle First Bedrock First Bedrock Lithology Lithology (Wellogic Number of (Wellogic Number of Records) Wells Percent Records) Wells Percent Shale 392 41 Shale 99 47 Sandstone 370 39 Sandstone 69 33 Sandstone & Sandstone & Shale 133 14 Shale 35 16 Limestone Limestone Undifferentiate Undifferentiated 52 6 d 9 4 Total Wells 947 Total Wells 212 Marshall Coldwater Michigan Marshall Coldwater Michigan (%)* (%)* (%)* (%)* (%)* (%)* 83.7 5.7 10.6 94.3 5.7 0 Unconsolidated Unconsolidated Sediments Sediments Maximum Average Minimum Maximum Average Minimum (Feet) (Feet) (Feet) (Feet) (Feet) (Feet) 29 274 118 40 240 111 *Area of formation in study area. overlying Deanville Mountain appears to be relatively thin, possibly hung up on the bedrock high just below that area.

3.6 Geophysics

As noted earlier, few water-wells penetrate to bedrock within the boundaries of the Imlay channel (Figure 3.23A). In order to determine bedrock depth and thicknesses of unconsolidated sediment within the Imlay channel, geophysical data were collected along

Clear Lake Road (CLR), ~600 m south of the Imlay channel’s topographic divide

(Figures 3.1 and 3.23A). The two geophysical methods used were electrical resistivity, i.e. Schlumberger vertical electrical soundings (VES), and passive seismic (Tromino) surveys. CLR runs perpendicular to the Imlay channel and has abundant water-well and geotechnical data, in contrast to many of the other roads that run perpendicular to the channel. Once geophysical data were collected along CLR and correlated to water-well

and geotechnical data, it could be compared to geophysical data collected along other roads running across the Imlay channel, resulting in bedrock depth estimates of areas of the channel with sparse water-well records.

Resistivity. VES data (Figure 3.23B) were plotted along the A-A’ transect (Figure

3.23A). Depths associated with the VES signals were plotted relative to stratigraphy from

Michigan Wellogic water-well records (Figure 3.23A) and two geotechnical borings. All

Wellogic wells were within 150 m of the profile and penetrate to bedrock. Resistivity values for the VES stations were between less than 10 ohm-m to >1,000 ohm-m (Figure

3.23B). The exception is VESIC8C (Figure 3.21B), where resistivity decreases from ~80 to 60 ohm-m, i.e. the sediment becomes more conductive, at the same elevation (~235 masl) as the laminated silt and clay in the geotechnical borings. Also, two of the three

VES logs taken within the Imlay channel, i.e. VESIC7C and VESIC9C, had resistivity values >1000 ohm-m, while one, VESIC6C taken outside the Imlay channel, had resistivity values <200 ohm-m.

None of the data from the VES survey appears to correlate with logged stratigraphy along CLR, with the exception of VESIC8C (Figure 3.23B). One issue with the VES data is that, due to overlapping resistivity values, it is nearly impossible to correlate the VES signal to a specific geologic material (Schrott and Sass, 2008), e.g. unconsolidated sand and sandstone overlap across three orders of magnitude (10 – 103)

(Sheriff, 1989). Another reason for the lack of correlation between water-well data may be due to a lack of a sharp resistivity contrast between bedrock and the overlying unconsolidated sediment (Schrott and Sass, 2008). Gravelly sand interpreted as post-

channeldivide.

, two passive seismic (Tromino) (Tromino) seismic two , passive

(B)

the topographic the

m or velocity in B and C may be reversed for economy of space. be for in economy B mreversed and velocity or may C

Ohm

well logs across the Imlay channel, ~600 m south of channel, m ~600 Imlay the well logs across

ations (A) of 4 Schlumberger vertical electrical sounding vertical stations ationsof Schlumberger 4 (A)

, and , water seven

. Loc .

(C)

stations on values that scales indicate that Note 3. Figure

glacial fill was collected from two vibracores (Figures 3.7 and 3.8) and a hand-dug pit

(Figure 3.10) within the Imlay channel. The sand and gravel contains fragments of mica

(biotite and muscovite), which is common within sandstone units of the Marshall

Formation (Monnet, 1948; Westjohn and Weaver, 1996). The Marshall Formation is the most abundant formation within the study area (Figure 3.22B; Table 3.3). Due to a lack of correlation between water-well stratigraphy, bedrock elevations established with VES data in this study are rejected.

Passive Seismic. In another effort to estimate bedrock depth following the VES survey, Tromino passive seismic surveys were also used in an attempt to estimate depth to bedrock and thicknesses of unconsolidated sediments along Clear Lake Road. Depth values associated with the Tromino data (Figure 3.23C) are plotted on A-A’ (Figure

3.23A) along with Wellogic stratigraphy and two geotechnical boings. Velocities associated with TRMIC11A are: 290 ms-1 (first layer) and 715 ms-1 (second layer); and for TRMIC12A: 330 ms-1 (first layer) and 725ms-1 (second layer). Depths to bedrock for the Tromino data were 23.5 m below the surface for TRMIC11A and 28.5 m for

TRMIC12A.

Like those of the VES surveys, the Tromino data were mostly inconclusive, with data from 33 of the 35 survey points failing parameters established by SESAME (2004).

Study areas where data from passive seismic surveys did correlate well with bedrock depths had hard, flat bedrock surfaces (Chandler and Lively, 2016). The bedrock surface in Lapeer County appears to be irregular where water-well records are available (Figure

3.23), especially near the Imlay channel’s topographic divide. Likewise, a low contrast in velocities associated with soft bedrock and the overlying unconsolidated sediments may

have been responsible for the poor results for much of the passive seismic data, although descriptions of bedrock hardness are lacking in the water-well data. Due to high failure of

Tromino data relevant to SESAME standards and lack of any nearby water-wells to correlate the two data points (Figure 3.23C), bedrock surfaces established from these data are rejected in this study.

3.7 Lake Maumee Water Levels and the Imlay Channel

Lake Maumee is described in the literature as having three or four levels and is reported to have drained from two outlets: Ft. Wayne outlet (FWO) in Indiana, and the

Imlay channel on the thumb of Michigan (Figures 1.3 and 1.5). A study by Bleuer and

Moore (1971) reported that the FWO was buried by ~3 m of sediment, which implies that all levels of Lake Maumee could have drained there. In order to understand which stages of Lake Maumee might have also drained through the Imlay channel, a cross-section of the channel was taken along Clear Lake Road (CLR), near the channel’s topographic divide, as well as another cross-section running longitudinally within the channel.

Ground penetrating radar (GPR) was used on a strandline associated with the highest stage of Lake Maumee and is used in this study to estimate water levels of Lake Maumee within the cross-sections. Finally, rebound curves for highest and lowest stages of Lake

Maumee, as well as Lake Whittlesey, were calculated in the western ALE basin.

Ground Penetrating Radar. The Goodland Church, located ~6 km south of south of the Imlay channel’s topographic divide (Figure 3.1), was first described by Leverett and Taylor (1915) who reported it as the location of a strandline associated with highest

Lake Maumee (Figures 3.1, 3.24, 3.25, 3.26A). The strandline is ~1.3 km long, is

truncated by a small valley, and found at a single location within the study area (Figures

2.24 and 3.25). In order to determine whether or not the feature is a strandline, GPR was used at the Goodland Cemetery across the road from the Goodland Church. The strandline is oriented north-south, and sits atop a wider ridge (Figure 3.24 and 3.25) composed of loamy till, according to the USDA Soil Survey (2016). The strandline is

~20–40 m wide and has an elevation of ~260 masl at the Goodland Cemetery. A hand- dug pit revealed about 0.80 m of sand and gravel near the top of the ridge, with cobbles as large as 10 cm. A 100 m long GPR transect, GPRMS1C (Figure 3.26B), was run south to north and nearly parallel to the long axis of the ridge. A 40 m transect, GPRMS2C

(Figure 3.26B), intersects GPRMS1C at ~66 m and runs west to east, i.e. toward the center of the Imlay channel. Three types of radar reflections are observed in GPRMS2C: those that dip westward or are horizontal, those that dip eastward (toward the channel) or are horizontal, and those that are very faint due to attenuation of the electromagnetic waves. Eastward dipping reflections are interpreted as shoreline progradation toward the channel, while reflections dipping westward are interpreted as overwash (Taylor and

Stone, 1996; Buynevitch, 2000; Tamura, 2012). Faint reflectors occur where the signal begins to attenuate, i.e. at a depth of ~1–1.5 m. Here attenuation is interpreted as the interface of the sand and gravel of the strandline and the loamy till of the ridge it sits on

(Figures 3.24 and 3.25). The loamy till would likely have a higher concentration of fine sediments, i.e. silt and clay, which would attenuate the signal, as opposed to reflecting it back to the GPR receiver (Daniels, 2000; Bristow and Jol, 2003). Geomorphic data shows the strandline is built upon a previously deposited ridge that is up to 15 times wider than the strandline. The strandline is most likely a spit, as opposed to a wave cut terrace, since

Figure 3.24. Hillshade DEM overlain with elevation data for a strandline associated with the highest Lake Maumee stage of ALE.

Figure 3.25. Oblique 3D view of the strandline at the Goodland Cemetery in Lapeer County. Michigan.

respectively. features, sh

in shallow a of observed as and primarily sand comprised is gravel,

) ) ationand overwa

strandline(B

Strending -

TheN

3.26. A strandline associated with highest Lake Maumee (A and B) was first identified by Leverett and (A 3.26. Leverett and firstwith by identified B) Maumee was Lake associated strandline A highest

Figure Figure Taylor(1915). westward(C) contains eastward and transect stratigraphy GPRMS2C of Radar m transect. 40 the GPR the pitnear prograd as (E) dippinginterpreted reflections

its elevation is lower in the north, possibly as a result of an interaction with the valley there (Figures 3.25 and 3.26A).

Geologic Cross-Sections. Two cross-sections for the area were created by interpreting water-well records from Michigan’s Wellogic database, and one used data from previously collected geotechnical borings (Figure 3.27) (PSI, 2013). Bedrock profiles from the interpolated map (Figure 3.22B) were also plotted in the CLR cross- section and the longitudinal cross-section (LCS) of the Imlay channel.

Clear Lake Road. The ~10.5 km (6.5 mile) CLR cross-section is based on the interpretation of 50 previously collected geotechnical borings along the profile and 18

Wellogic water-wells that penetrate to bedrock and were located within 150 m of the profile (Figures 3.23 and 3.27). The surface topography of the cross-section (Figure 3.28) is plotted using colors and symbols (e.g. Qdlt, Qdct, etc.) that correlate to geologic units on the digital surficial geology map. Areas of high surface topography are predominantly loamy diamicton (Qdlt), except in the far eastern portion of the cross-section where clayey diamicton (Qdct) is present. Many valleys contain post-glacial alluvium and peat, with basal layers of peat as old as 11.8–12.1 calibrated ka BP (Figure 3.9; Table 3.2).

Undifferentiated glaciolacustrine and glaciofluvial sediment (Qglu) fills several valleys east of the Imlay channel, based on USDA soil survey data and on the geotechnical borings, which show beds of sand and gravel or silty sand at depth (Figure 3.28). Bedrock along the CLR cross-section is generally shale or sandstone and is interpreted to represent the Mississippian Marshall Formation (Westjohn and Weaver, 1996) based on data from the bedrock map (Figure 3.23B). Bedrock elevations from water-well logs vary from

~207 masl (679 ft. asl) near the valley leading to Clear Lake, to ~244 masl just west of

Figure 3.27. Location of water-wells and geotechnical borings for the Clear Lake Road cross-section (Figures 3.28 and 2.29) and the longitudinal cross-section of the Imlay channel (Figure 3.30).

section along Clear

g interpretations of interpretations g

wellborings, and geotechnical

p),water planes Lake for Erie, of four stages ancestral

surface. Figure 3.28. Surficial 3.28. cross geologic Figure 3.3) showin Lake(Figure Road water collected previously geologic topography surficial surface and (digital geology ma bedrock topography, bedrock and an interpolated

the Imlay channel’s eastern boundary (Figures 3.28 and 3.29). Also plotted is the bedrock surface from the interpolated bedrock map (Figure 3.22B). The interpolated bedrock profile has a low elevation of ~216 masl (~710 ft. asl) near the valley leading to Clear

Lake and a high elevation of ~238 masl (780 ft. asl) just east of the Imlay channel’s eastern boundary (Figures 3.28 and 3.29). Water-well records indicate the lowest bedrock elevation within the Imlay channel, i.e. ~235 masl (770 ft. asl), is located just west of a bedrock high in the center of the channel with an elevation of ~238 masl (780 ft. asl). The westernmost valley in the CLR cross-section connects with Clear Lake in the south

(Figures 3.1, 3.3, 3.12, 3.27), contains peat, muck, and post-glacial alluvium, and likely overlies diamicton (till) or bedrock. A core taken at Clear Lake (Figure 3.12) extends to an elevation of ~237 masl (777 ft. asl) where it reaches sand and gravel. The stratigraphy from Clear Lake, as well as chronology from the core and fill within the channel, suggests that the Clear Lake valley began to fill with sediments before ~14.9 ka (Tables

3.1 and 3.2). Stratigraphy from the core taken at Clear Lake was also used to estimate the depth to the top of fill in the deepest part of the lake, i.e. ~13.7 m (Figure 3.12). This estimate suggests the top of post-glacial fill in the Clear Lake valley may be ~4–5 m deeper than the highest bedrock elevation within the valley once occupied by the Imlay channel. The lowest bedrock elevation from water-well data along the CLR cross-section within the Imlay channel, i.e. 235 masl (Figures 3.28 and 3.29), is interpreted as the controlling elevation of water draining through the Imlay channel. While the interpolated bedrock profile indicates the controlling elevation maybe ~6 m lower than the one determined from water-well data, the interpolated bedrock profile along CLR differs from that of drilled wells by more than 3 m in some cases, even in areas where bedrock data

well data, and an -

units, water planes from planes from water units,

sectionLake along Clear -

surface. polatedbedrock

four stages of ancestral Lake Erie, bedrock bedrock Erie, Lake of ancestral fourstages geologic cross geologic 3.3) showing generalized (Figure Road surficialgeologic water topographyfrom inter Interpretation 3.29. of Figure surficial the

from wells is available and was used in the interpolation (Figures 3.28 and 3.29). Thus, the bedrock profile taken directly from water-well logs is given more weight than those interpolated using data entered into ArcGIS for the CLR cross-section (Figures 3.28 and

3.29).

Longitudinal Cross-Section of the Imlay Channel. Sediment thicknesses and bedrock elevations for the ~22 km (14 mile) longitudinal cross-section of the Imlay channel (LCS) (Figures 3.27 and 3.30) are based on 14 Wellogic water-well logs up to

~700 m from the transect (Figure 3.27). Sediment interpreted as diamicton has a peak elevation of 238 masl (780 ft. asl) at a distance of ~6,096 m (~20,000 ft.) along the profile. Sediment interpreted as post-glacial alluvium, i.e. sand, gravel, silt, and clay, was thickest (~39 m) in the bedrock depression at ~5,593 m (~61,000 ft.) (Figure 3.26).

Bedrock was shale or sandstone and is interpreted to represent the Mississippian Marshall

Formation (Westjohn and Weaver, 1996) (Figure 3.22B). Bedrock elevations from water- well data ranged from ~197 masl (~646 ft. asl) at Well 6 to ~235 masl (770 ft. asl), where the LCS intersects the CLR cross-section (Figures 3.22 and 3.27). Interpolated bedrock elevations range from 220 masl (720 ft. asl) in the northern part of the LCS, to ~235 masl

(770 ft. asl), just south of Well 9 (Figure 3.30). The longitudinal profile crosses a bedrock high on the interpolated bedrock map, which seems to correlate closely with the bedrock sill (Figure 2.22B). Since water-well data for the bedrock profile within the LCS is sparse, the interpolated bedrock profile is given more weight than the water-well bedrock profile, except for data from Wells 8 and 9 (Figures 3.27 and 3.30).

wellborings,

bedrock bedrock

sectionof channel Imlay the

iously collected water iouslycollected

topography, and an interpolated interpolated an and topography, Figure 3.30. Longitudinal 3.30. geologic Figure cross showingof 3.1) (Figure interpretations prev planes for topography), surface water stages, fourALE lake surface. bedrock

Water Planes. A strandline ~6 km south of the CLR cross-section (Figures 3.1 and 3.24) has an elevation of ~260 masl (850 ft. asl) and is associated with highest Lake

Maumee (Leverett and Taylor, 1915). Leverett and Taylor (1915) established lake planes for the three stages of Lake Maumee as follows: highest Maumee (Maumee I) at 244 masl (800 ft. asl); middle Maumee (Maumee III) at 238 masl (780 ft. asl), lowest

Maumee (Maumee II) at 232 masl (761 ft. asl), or ~ 6 m between each stage. The elevation of these lake stages are based on those of unwarped strandlines near the zero- isobase for Lake Maumee, located near the city of Birmingham, Michigan (Leverett and

Taylor, 1915). In order to determine what stages of Lake Maumee were able to drain through the Imlay channel, 6 m was subtracted from the strandline associated with highest Maumee (260 masl/850 ft. asl) to get the elevation of middle Maumee (254 masl/830 ft. asl), and 12 m from the highest Maumee strandline to get an elevation for lowest Maumee (248 masl/810 ft. asl). Similarly, a strandline with an elevation of ~230 masl (755 ft. asl) is located ~40 km southeast of the topographic divide (Figure 1.7) and is associated with Lake Whittlesey. The elevations of these lake planes are plotted on the

CLR (Figures 3.28 and 3.29) and longitudinal (Figure 3.30) cross-sections.

The Imlay channel is filled with clay, silt, sand, and gravel, including laminated silt and sand at a depth of ~12 m below the surface. A minimum age for the fill is ~14.9 ka from an OSL age taken within the channel (Figure 3.7; Table 3.1). The top of fill within the valley leading to Clear Lake is about as deep as the lowest bedrock elevation from the Imlay, i.e. 235 masl, but may be ~7 m deeper (Figures 3.28 and 3.29) when estimating the upper surface of fill at the deepest point in Clear Lake (Figure 3.11). This suggests that the valley leading to Clear Lake may have been created by a separate

process than the one that cut the Imlay channel. The bedrock profile from water-well records is favored along the CLR cross-section, where data is abundant and in close proximity to the transect (Figures 3.22, 3.27, 3.28, 3.29). In contrast, the interpolated bedrock surface is favored on the LCS (Figure 3.30), where water-wells can be located up to 700 m from the transect (Figure 3.27), except at the intersection of the LCS and CLR cross-section where two elevations come from water-wells (Wells 8 and 9; Figures 3.27 and 3.30) in close proximity to the LCS. One of these water-wells (Well 9) suggests the controlling sill of the Imlay channel is bedrock with an elevation of 235 masl. Water planes for three stages of Lake Maumee were calculated based on the elevation of a strandline associated with highest Lake Maumee (Figure 3.24) (Leverett and Taylor,

1915) near the study area, while the elevation for Lake Whittlesey comes from a strandline ~20 km to the south of the study area. Plotting the water planes with the controlling bedrock elevation for the Imlay channel suggests that water from three stages of Lake Maumee, uncorrected for isostatic rebound, could have drained through the

Imlay channel, while water from Lake Whittlesey could not (Figures 3.28, 3.29, 3.30).

Glacial Isostatic Adjustment. When continental ice sheets blanketed North

America, the weight of the ice depressed the underlying mantle. As the ice melted, the mantle adjusted to the unloading, a process known as glacial isostatic adjustment (GIA) which continues to affect the Great Lakes region today (Mainville and Craymer, 2005).

While strandlines for a single ALE stage may be at different elevations today, depending on their locations within the ALE basin, prior to the effects of GIA these strandlines were at the same elevation when that particular lake stage existed (Breckenridge, 2013).

Highest (244 masl), middle (238 masl), and lowest (232 masl) stages of Lake Maumee

were reported (Leverett and Taylor, 1915) to be unaffected by isostatic adjustment in the western ALE basin until just south of Birmingham, Michigan, a suburb of Detroit. These elevations are used in the literature (Leverett and Taylor, 1915; Eschman and Karrow,

1985) to describe the water planes for each stage of ALE. However, there has never been a comprehensive study of strandline elevations in the western ALE basin, which has been made easier in recent years with the advent of geographic information systems and

LiDAR-derived DEMs.

Rebound Curves. Rebound curves were created by tracing strandlines on LiDAR- derived DEMs in the western basin of ALE for highest and lowest Maumee, as well as

Lake Whittlesey. Highest Lake Maumee strandlines are represented by 499 elevation points (Figure 3.31). The lowest elevation plotted is 241 masl and the highest is 260 masl.

Strandline elevations for highest Maumee generally increase to the northeast in the direction of maximum uplift (DMU). Strandlines corresponding to lowest Lake Maumee are represented by 556 elevation points (Figure 3.31). The lowest elevation plotted is 230 masl and the highest is 243 masl. Strandline elevation increases to the northeast in the direction of maximum uplift (DMU).

Strandlines associated with Lake Whittlesey are represented by 787 elevation points. The lowest elevation plotted is 221 masl and the highest is 232 masl (Figure 3.31).

Strandlines for highest Maumee are difficult to trace just north of the Michigan-Ohio border for ~40 km until just south of Tecumseh, Michigan. DEMs appear to show a large valley running parallel to the margins of ALE, which terminates as a delta within the

ALE basin just south of the western arm of the Defiance Moraine. It is possible that sediments in this valley post-date highest Maumee and may have obscured the highest

o plotted is the elevation the is plotted o of till the

ee, and Lake Whittlesey along the direction of maximum uplift. and along Als direction the Lake of Whittlesey ee, maximum

sill at Ft. Wayne, Indiana, the Imlay channel’s bedrock sill and the topographic divide, and the locations of locations the and divide, topographic sill and the channel’s bedrock Imlay Indiana, the Wayne, Ft. sillat Maum Michigan for reference. Michigan and Detroit, Toledo,Ohio, Plymouth, lowest Lake in Strandline 3.31. elevations western the Maumee, Lake ALE for basin highest Figure

Maumee strandlines there. Highest and lowest Maumee strandlines are also difficult to trace for ~45 km beginning near the city of Plymouth, Michigan (Figure 3.31), and is likely due to urban development in the Detroit metropolitan area. Strandlines for Lake

Whittlesey were traced to the northern border of Sanilac County, above which LiDAR data is conveniently unavailable. Whittlesey strandlines have been noted for their prominence as compared to other strandlines in the area, e.g. lakes Maumee and Arkona

(Forsyth, 1958), which may account for why they were traceable through the Detroit metropolitan area while those associated with the Lake Maumee stages of ALE were not.

The sill at the Ft. Wayne outlet is composed of till (Bluer and Moore, 1972) and plots at an elevation of 226 masl (Figures 3.31 and 3.32). The bedrock sill at the Imlay channel has an elevation of ~235 masl based off the lowest bedrock surface indicated from water-wells in a cross-section along Clear Lake Road (Figures 3.27, 3.28, 3.29).

Also plotted is the topographic divide at an elevation of 250 masl, which is located almost directly above the bedrock sill (Figures 3.28, 3.29, 3.30).

Adjusted Water Planes. The sill of the Imlay outlet and the highest Lake Maumee stage were adjusted for ~20 m of total rebound (Figure 3.32), i.e. the difference in elevation between strandlines of highest Lake Maumee at Ft. Wayne, Indiana and the

Imlay channel in Lapeer County, Michigan (Figure 3.31). Elevations for lowest Maumee were adjusted ~ 4m relative to their current position above the Ft. Wayne Outlet. The reason for this is because strandlines for lowest Maumee grade to the Ft. Wayne outlet, but are difficult to trace beyond ~30 km south of the Imlay outlet. Whittlesey strandlines become untraceable ~50 km northeast of Ft. Wayne, but are traceable within ~30 km

Figure 3.32. Water planes of highest and lowest Lake Maumee, Lake Whittlesey, and the outlets at Imlay and Ft. Wayne adjusted for post-glacial isostatic rebound. Profiles of the north half of the Ft. Wayne and the east half of the Imlay are also shown.

of the bedrock sill of the Imlay channel. Thus, the Whittlesey strandlines are adjusted relative to their position with the sill at the Imlay outlet.

This model (Figure 3.32) assumes no isostatic adjustment occurred at the outlet in

Ft. Wayne, Indiana. The bedrock sill at the Imlay outlet sits ~11 m lower than the sill of the Ft. Wayne outlet, i.e. at an elevation of 215 m. The model also suggests that neither the Imlay outlet, nor the one in Ft. Wayne, Indiana, would be able to accommodate water from Lake Whittlesey. This agrees with morphological evidence (Figure 1.7), which shows continuous strandlines from Whittlesey (Figure 1.7) passing south of the mouth of the Imlay channel ~30 km south of the sill near CLR and ~50 km northeast of the Ft.

Wayne outlet.

3.8 Surficial Geologic Map

Accompanying the thesis on a CD-ROM is a full-size, digital map of the Burnside

7.5’ Quadrangle and unit descriptions. Geologic units were interpreted and mapped using data from the USDA Web Soil Survey (2016), water-well records from the Michigan

Wellogic database (2018), unpublished geotechnical borings from the Genesee County

Drain Commission (PSI, 2018), unpublished LiDAR from FEMA (now available from the USGS), and field work.

Map Units

Peat and muck (Qp) spans the latest Pleistocene and Holocene in age, as determined by radiocarbon (Figure 3.4; Table 3.2) and OSL dating (Table 3.1). Qp is found in areas of low topography, including valleys, and may overlie Holocene or

Pleistocene alluvium as well as Pleistocene diamicton (Qdlt, Qdct), glaciolacustrine (Ql), glaciofluvial (Qsg2), and undifferentiated glaciofluvial-glaciolacustrine (Qglu) deposits.

Alluvium (Qal) is stratified sand, silt, clay, mud, and/or gravel. Radiocarbon

(Table 3.2) and OSL dates (Table 3.1) indicate alluvium can be Holocene or late

Pleistocene in age. Nearly all of the alluvium was described as glaciofluvial or glaciolacustrine or undifferentiated glaciofluvial-glaciolacustrine deposits in the USDA soil survey. However, chronology (Tables 3.1 and 3.2) in the Imlay channel suggests this sediment was deposited by streams after deglaciation.

Loamy diamicton (Qdlt) was identified by the USDA Soil Survey and accounts for a majority of the surficial units within the quadrangle. In the field, the diamicton is observed to have a substantial silt and sand component with lesser portions of clay and gravel (e.g. pebbles, cobbles and boulders). Qdlt is found on morainal ridges and

hummocky surfaces and may have overridden older glaciofluvial deposits. Qdlt is associated with the Huron-Erie and Saginaw lobes of the LIS.

Clayey diamicton (Qdct) was identified with the USDA Soil Survey data. In the field, Qdct is predominantly clay-rich and silt-rich with lessor portions of sand and gravel. Qdct is found primarily on morainal ridges in the eastern portion of the quadrangle, but also occurs on hummocky surfaces. Qdct is associated with the Huron-

Erie and Saginaw lobes of the LIS.

Wave-worked diamicton (Qdwt) is found overlying clay diamicton (Qdct) in small areas on the east-central portion of the quadrangle. Qdwt is matrix (i.e. silt and sand) supported to clast (i.e. gravel) supported and overlies dense clay-rich diamicton

(Qdct). It is found at elevations of ~252 masl (830 ft. asl) that correlate with the middle

Maumee stage of ancestral Lake Erie.

Glaciolacustrine deposits (Ql) are comprised of sand, silt and clay, according to

USDA Soil Survey data. A digital elevation model of Lapeer County that was generated using unpublished LiDAR data shows several depressions on till plains and morainal ridges that are interpreted as Ql in this quadrangle.

Glacial outwash (Qsg2) is stratified sand and gravel. In low areas Qgs2 may be overlying or underlying diamicton (Qdlt, Qdct, Qdwt) or found proximal to glaciolacustrine (Ql), undifferentiated glaciofluvial-glaciolacustrine (Qglu), or alluvial

(Qal) deposits. On slopes, Qsg2 may be found as kames with evidence of ice-collapse features, e.g. faults and deformation features exposed in gravel pit walls (Figure 312).

Undifferentiated glaciofluvial-glaciolacustrine sediment (Qglu) is comprised of laminated silt and clay and bedded sand and gravel. Silt and clay may show signs of

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deformation (Anketell et al., 1970), e.g. sand pillows in underlying mud or disturbed laminations in slit and clay, and may overlie or underlie bedded sand and gravel.

3.9 Summary

This study in northeastern Lapeer County produced 5 OSL and 10 radiocarbon ages, including minimum ages for fill within the Imlay channel, as well as a minimum age for local deglaciation that duplicates that of a previous study (Burgis, 1970), although the stratigraphy near the Weaver Drain site where organics were collected by Burgis could not be reproduced.

Geomorphic data was greatly aided in this study due to a LiDAR survey of Lapeer

County and resulted in the identification of landforms, e.g. morainal ridges, channels, that may not have been apparent to earlier workers who had to rely on topographic maps for elevation data. A new map (digital attachment) of the surficial geology of the Burnside

Quadrangle helps to illustrate the complex suite of glacial and post-glacial deposits within the study area.

Numerous water-well records and geotechnical borings helped to reveal the stratigraphy of the channel and that a bedrock sill at 235 masl likely controlled the drainage of water from ALE through the outlet. Water-well records were also instrumental in developing drift thickness and topographic maps of the study area.

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Chapter 4

Discussion

4.1 Introduction

The results obtained in this study have implications beyond the bounds of the

Imlay channel and Lapeer County, Michigan. Abundant subsurface data from geotechnical borings and water-well records, and was unavailable to previous workers, helped establish that the sill of the Imlay channel is buried by many meters of post- glacially deposited debris. Likewise, new OSL and AMS radiocarbon data from the study area have helped to clarify when local deglaciation occurred and have implications for regional deglaciation as well.

4.2 Deglaciation

Deglaciation for the thumb of Michigan and limiting chronology for deglaciation of the Saginaw Lobe since the 1970’s has depended upon a single radiocarbon age from plant matter collected at the Weaver Drain site (Burgis, 1970). A radiocarbon age from wood within colluvium within a channel (Figures 3.13 and 3.18) ~1 km northeast of the

Weaver Drain site in this study is 16.6–17.0 calibrated ka BP (Figure 3.18; Table 3.2).

This value is consistent with the radiocarbon sample collected by Burgis (1970) of 16.1–

17.3 calibrated ka BP. The colluvium is probably, in part, reworked glacial till, along with sand and gravel, that was derived from the eastern slope of Deanville Mountain and deposited in the channel. Gullies (Figure 3.13) suggest that Deanville Mountain has undergone erosion since deglaciation of the area. Based on the proximity of the Weaver

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Drain site (Burgis, 1970) and the newly discovered morainal ridge, it is likely that the material recovered by Burgis for radiocarbon dating is associated with that moraine located ~1 km to the west. Direction of water flow through the channel at the base of

Deanville Mountain, which was most likely a conduit for glacial meltwater as the ice margin was retreating to the east, is difficult to interpret because the bottom of the channel appears to be covered by sediments that likely post-date deglaciation, including alluvial fans, which enter the channel from gullies on Deanville Mountain (Figure 3.13).

The radiocarbon age from material within the channel from this study (16.6–17.0 calibrated ka BP) and from the Weaver Drain (16.1–17.3 calibrated ka BP) (Burgis,

1970) both represent minimum deglacial ages for northeastern Lapeer County.

The minimum ages for deglaciation from this study and Burgis’ (1970) are younger than expected when compared to ages associated with deglaciation of the

Saginaw Lobe (Figure 4.1). In a recent study, Horton (2015) dated the Sturgis Moraine,

~230 km southwest of the Weaver Drain, resulting in a minimum-limiting age of 16.3–

16.9 calibrated ka BP. The radiocarbon age from the Sturgis Moraine and Weaver Drain suggest that ice from the Saginaw Lobe retreated at an average rate of ~480 m/year. A more recent investigation (Dziekan, 2017) of the Sturgis and Shipshewana moraines, also associated with the Saginaw Lobe, used OSL to obtain minimum ages for the moraines, resulting in ages of 19.0–21.8 ka and 22.5–24.7 ka, respectively (Figure 4.1). The new

OSL age from the Sturgis Moraine suggests the average rate of ice retreat was ~89 m/year, and within or below the range of values reported for retreat of the LIS in other studies, e.g.: 155 m/year New England (Ridge, 1997); 105–270 m/year, Des Moines Lobe

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(Lepper et al., 2007); 23–300 m/year, Ft. McMurray, Alberta (Fisher et al., 2009); 150–

250 m/year, Thunder Bay, Ontario (Lowell et al., 2009).

Leverett and Taylor (1915) suggested that oscillatory ice-margins in northeastern

Lapeer County resulted in cyclic water-level fluctuations in the ALE basin. The geomorphology of the study area suggests that one readvance of ice may have occurred.

Evidence of the readvance comes from an absence of strandlines associated with Lake

Maumee to the east of the interlobate moraine, except for the single strandline near the

Goodland Cemetery. Of course, there is a possibility that no readvance ice margin ever occurred and that the Imlay, Goodland, and Deanville moraines represent a recessional event where the ice margin simply receded from the area (Figure 1.8D).

Two possibilities exist for the timing of kame deposits on Deanville Mountain

(Figure 3.13). The first is that the kames were deposited during the Early Wisconsinan glacial period at ~70 ka (Elyes and Westgate, 1987; Mickelson and Colgan, 2003; Dyke,

2004), which corresponds to Marine Isotope Stage 4. The two OSL ages from the kame at the BSG pit fit within this early Wisconsinan time frame: 70.3 ± 6.9 ka collected from sand at the top, proximal position of the deposit (Figure 3.16); and 60.6 ± 6.0 ka (Figure

3.17). Another possibility is that the kame deposits are Late Wisconsinan in age, i.e. less than ~30 ka (Elyes and Westgate, 1987), but sediment within the deposits were insufficiently bleached prior to deposition, possibly due to the turbidity of proglacial meltwater in close proximity to the ice margin (Rittenour, 2008). This second possibility also makes sense when considering the preservation of the feature. Presumably, if the kame at BSG was deposited 60,000–70,000 years ago it would have been overrun by an unknown number of major advances of the LIS. The fact that the loosely consolidated

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Figure 4.1. Plot showing minimum ages of deglaciation for the Saginaw Lobe (Sturgis and Shipshewana moraines), Huron Lobe (Port Huron Moraine), the Weaver Drain, and Lake Maumee. Also illustrated are average ice-margin retreat rates between several locations and GRIP data from Rasmussen et al. (2006). Figure was modified from Fisher et al. (2015).

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sediment of this kame is still preserved at the BSG pit suggests it was most likely deposited sometime during the Late Wisconsinan.

4.3 Channel Fill and Chronology

The floor of the Imlay channel is filled with sediment (Figures 3.27, 3.28, 3.29;

Tables 3.2 and 3.2) that postdates accepted ages of deglaciation in the region (Figure 4.1)

(Burgis, 1970; Horton, 2015; Dziekan, 2017). The lithology of sediment within the Imlay channel is similar to that of infilled spillway valleys in other studies, i.e. clay, silt, sand, and gravel. Valleys associated with other spillways that have similar stratigraphy and geomorphology as the Imlay channel, with flat bottoms and/or alluvial fans, include the

Wabash Valley in Indiana (Bleuer and Moore, 1971) and the Souris Spillway in North

Dakota (Kehew and Boettger, 1986).

The sediment filling the channel is likely reworked glacial and glaciofluvial sediment eroded from the highlands that constitute Deanville Mountain and the interlobate area (Figure 3.1). The loamy till identified by the USDA Soil Survey is a good candidate for the source of this channel-fill material and is abundant in both these areas

(Digital Map). Geomorphic evidence that fill within the channel comes from the surrounding highlands is illustrated in the DEMs, which show numerous gullies and fans leading to and building into the channel (Figure 3.1 and 3.3). Studies in paraglacial environments in which the landforms are adjusting to post-glacial conditions of deposition and erosion (Ballantyne, 2001), suggest that alluvial fans in alpine valleys continued to be deposited and did not fully stabilize for thousands of years following local deglaciation (Lian and Hickin, 1996). The timing of sedimentation within the Imlay

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channel may have also been prolonged due to the effects of permafrost on the landscape.

Estimates of ~4,000 years have been made for the thawing of late Pleistocene permafrost following the deglaciation of southern Wisconsin after the last glacial maximum (Clayton et al, 2001). Indeed, the uppermost sediment from alluvial fans (Figures 3.3, 3.9, 4.1) on the topographic floor of the Imlay channel (Figure 3.10) appears to be ~1,800 years younger than the uppermost channel fill (Figure 3.9), and ~4,700 years younger than sand and silt recovered about 5 m lower within the channel (Figure 3.4). Ballantyne (2001) noted that initial sedimentation within paraglacial environments is very rapid at the onset and slows toward the end as vegetation becomes established. The youngest alluvial fans, dated at 10.2 ± 0.9 (Figure 3.10; Table 3.1), fully stabilized in the Holocene and may be indicative of this differential sedimentation.

The timing of sediment fill within the channel is interesting as well. The most limiting minimum OSL age of ~14.9 ka from the channel suggests that valley fill began sometime prior to the Bolling-Allerod interstadial. Ten OSL and radiocarbon ages from the study area fall within or just outside of the Bolling-Allerod (Figure 4.1). The cluster of ages near the Bolling-Allerod interstadial may indicate that the landscape began to stabilize around this time, with vegetation taking hold during the warmer interstadial.

In total, an ~6,000-year time lag in landscape equilibrium from rapid post-glacial sedimentation rates in the Imlay channel valley following deglaciation to the final stabilization of the landscape in the Holocene. Similar lag periods following deglaciation appear to have occurred in southern Michigan and northern Indiana, in the area of land once occupied by the Saginaw Lobe. As discussed earlier, Dziekan (2017) found that the minimum age of retreat for the Saginaw Lobe at the Sturgis Moraine in southern

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Michigan, based on OSL ages, was ~6,000 years older than a prior study on the moraine

(Horton, 2015) that used radiocarbon dating exclusively.

Clear Lake. The depression now occupied by Clear Lake (Figures 3.1 and 3.11) likely acted as a backwater to the Imlay channel when it was active. Sand and gravel deposits from a Livingstone core taken in the lake are similar (Figure 3.3.12) to sand and gravel from Facies A (Figure 3.6) taken within the channel. A radiocarbon age from woody material collected in the basal sand and gravel of Clear Lake suggests that it was deposited after 14.1–14.9 calibrated ka BP (Figure 3.12; Table 3.2). This radiocarbon age agrees with an OSL age from gravelly sand (Facies B; Figure 3.6) collected within the

Imlay channel of 14.9 ± 0.6 ka (Figure 3.8; Table 3.1). Clear Lake likely formed when an alluvial fan (or fans) (Figure 3.3) blocked the drainage route of streams flowing out of the depression before ~14.9 ka. Water probably came from streams flowing into the Clear

Lake depression, as well as groundwater from a sand and gravel unit about 220 – 223 masl (Figure 3.29), which is probably the source of groundwater entering the lake today.

The valley now occupied by Clear Lake (Figure 3.1) appears to have been created sometime before the Imlay channel was cut. A core taken in the shallow, eastern basin of

Clear Lake (Figures 3.11 and 3.12) suggests that the valley leading to Clear Lake is located at an elevation of ~231 masl, or ~4 m lower than the bedrock sill of the outlet to the east (Figures 3.28 and 3.29). It is not clear why the valley leading to Clear Lake did not fill with sediments in a manner similar to the Imlay channel, but one possibility is that buried ice in Clear Lake valley from a previous ice advance melted out ~14.9 ka, deepening the lake at that time. The interpolated bedrock map (Figure 3.22B) suggests that Clear Lake’s valley is not incised into bedrock. However, the degree to which

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bedrock may be scoured within the Clear Lake valley is unknown due to sparse water- well data in this area (Figure 3.22).

4.4 Water Drainage Through the Imlay Channel

The geologic cross-sections along Clear Lake Road (Figures 3.28 and 3.29) and the longitudinal profile of the Imlay channel (Figure 3.30) shed light onto which levels of

ALE might have drained through the channel. A bedrock sill within the Imlay channel near the topographic divide (Figure 3.1) is the most likely candidate for the controlling elevation of 235 masl for water from Lake Maumee. Evidence of a bedrock sill for the

Imlay channel comes from a water-well in the center of the channel that extends to bedrock (Figures 3.28, 3.29, 3.30). The topographic profile of the interpolated bedrock surface in the Clear Lake Road cross-section (Figures 3.24 and 3.25) indicates the bedrock just below the silt and clay laminations may be as low as 228 masl, meaning the

Imlay channel could accommodate waters from the Whittlesey stage of ALE. Of course, the interpolated bedrock profiles on both the Clear Lake (Figures 3.24 and 3.25) and longitudinal (Figure 3.30) cross-sections of the Imlay channel appear to differ by up to ~6 m compared with actual bedrock depths from water-well data. Furthermore, it is not possible that Lake Whittlesey drained through the Imlay channel because continuous strandlines for the lake can be traced eastward below the southernmost boundary of the

Imlay channel (Figure 1.7).

The isolation of the highest Maumee strandline near the Goodland Cemetery (260 masl) (Figures 3.24, 3.25, 3.26), which was used to calculate the water levels of the three levels of Lake Maumee, is somewhat problematic. Ice cover on lakes has been suggested

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as a being a mechanism for the suppression of beach ridge development in other proglacial lakes (McMillan and Teller, 2011). Another possible reason for the isolation of the strandline is that most it may have been destroyed by a readvancing ice margin.

LiDAR DEMs from the Fort Wayne outlet in Indiana show ~4 km long, bow-shaped strandlines that appear to grade toward that outlet. Similarly, ~5 km long strandlines associated with Lake Whittlesey (Figure 3.5) bow northward toward the Ubly outlet, which is reported to have drained that stage of ALE (Eschman and Karrow, 1985). In contrast, ~500 m north of the Goodland Cemetery, the strandline appears to grade downslope ~200 m eastward, into the valley that truncates the 1.3 km strandline (Figures

3.24 and 3.25). Because of this change in elevation, the strandline is most likely a spit, as opposed to a wave-cut terrace, and appears to be interacting with valley when it was deposited. While there is no evidence to corroborate that the lake that constructed the strandline at the Goodland Cemetery was highest Lake Maumee, or that whatever ALE lake stage that constructed it drained through the Imlay channel, it is the only strandline near the channel and will be used to reconstruct paleolake levels of ALE in this study.

4.5 Rebound Curves

Rebound curves for highest and lowest levels of Lake Maumee have similar slopes, suggesting these lake stages formed under similar rebound conditions at roughly the same time. In contrast, the shallower slope of the Whittlesey strandline indicates that it was not affected by glacial isostatic adjustment over as long a period of time as those of the Maumees. Rebound velocities of ~12.5 cm/century are estimated with the model in this study (Figure 3.31). However, the ICE-3 model used by Mainville and Craymer

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(2010) to estimate isostatic rebound in the Great Lakes Region suggested velocities of

~18 cm/century or ~28 m of total rebound between Ft. Wayne, Indiana and Imlay City.

Even though rebound velocities from this study are derived from elevations from LiDAR data, difficulty tracing the highest and lowest Maumee strandlines through the Detroit

Metropolitan area, and the highest Maumee strandline in general, may mean that the estimates of rebound rates from this study are inaccurate.

The effects of isostatic rebound were removed between the Ft. Wayne and Imlay outlets using the elevations of the sills in each location (Figure 3.32). This model suggests the bedrock sill at Imlay was ~11 m lower than the sill composed of till at the Ft.

Wayne outlet (Bleuer and Moore, 1972). Like the geologic cross-section at Clear Lake

Road (Figures 3.27, 3.28, 3.29), the lake level model adjusted for isostatic rebound suggests that all stages of Lake Maumee would have been able to drain from the Imlay channel once the area was free of ice, coeval with the outlet in Indiana. Little research is available that discusses two outlets draining a proglacial lake in unison. However, it is likely that the Imlay channel was the controlling outlet, once uncovered by the ice, since it would likely be rebounding at a faster rate than the outlet at Ft. Wayne to the south.

The Imlay and Ft. Wayne outlets appeared to be active coeval with one another. Other large lakes with two active outlets draining water at the same time include Lake Nipissing

(Thompson et al., 2011) and Lake Agassiz (Fisher et al., 2011; Lepper et al., 2013).

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Chapter 5

Summary, Conclusions, and Future Work

5.1 Summary

The Imlay channel has been studied sporadically in the last 100 years, which contributed to the ambiguity surrounding its role in the drainage of ALE. This study has greatly improved the understanding of the Imlay channel and its role in the drainage of

ALE due to recent infrastructure development in the study area, e.g. water-wells and geotechnical borings. These data suggest the sill of the Imlay channel is made of bedrock and is located just below the topographic divide at an elevation of ~235 masl. Cores taken within the Imlay channel for this study validate the stratigraphy of the previously collected geotechnical borings and water-well records, and chronology from those cores suggest that the valley once occupied by the Imlay channel is now filled with a sedimentary package that post-dates accepted ages of deglaciation in the area.

A better understanding of the surficial and bedrock geology was aided by the use of water-well records, soil survey data, and publicly available LiDAR-derived DEMs.

Identifying several geomorphic features, e.g. morainal ridges, valleys, and alluvial deposits would have been impossible without these data. The DEM also helped to trace strandlines of several stages of ALE that were used to construct a model of regional isostatic rebound, as well as another model that was used to remove such effects.

Based on the evidence from the analysis of rebound curves and adjusted rebound models, geologic cross-sections, geomorphology, bedrock and surficial geologic mapping, and chronologic data collected and analyzed in this study, the original

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hypothesis that “…water from any stage of Lake Maumee in the ALE basin was drained by the Imlay channel.” is supported.

5.2 Conclusions

All levels of the Maumee stage of ALE, i.e. highest, middle, and lowest, could have drained through the Imlay channel once it was free of ice, while there is no evidence that any other stage of ALE could have done the same. A model that adjusted sills of the

Imlay channel and Ft. Wayne outlet for the effects of glacial isostatic rebound suggests that it is possible that both outlets drained Lake Maumee.

Chronologic and stratigraphic data suggests that the valley once occupied by the

Imlay channel was infilled with sediment prior to 14.9 ka ± 0.5 ka and that the bedrock sill of the channel has a modern elevation of 235 masl, almost directly below the topographic divide identified by Burgis (1970).

Geomorphic evidence suggests that a local readvance of ice in the area was possible, but is not conclusive, as the preserved moraines may have been deposited in a simple recessional event as the ice left the area. The age of a radiocarbon sample collected by Burgis (1970) of 16.1–17.3 calibrated ka BP, is interpreted to represent minimum age for deglaciation in this study area and was replicated in this study with an age (16.3–16.9 calibrated ka BP) taken from organic material in reworked till deposited as colluvium in an outlet channel near the Weaver Drain.

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5.3 Future Work

The Pleistocene geology of the Imlay channel, and the broader area surrounding it, would be greatly improved through surficial geology mapping of adjacent quadrangles, as well as additional chronologic controls for moraines in the area. The nearest mapped

USGS 7.5’ quadrangle is located on the opposite end of Lapeer County and no other quadrangles are mapped with surficial geology north, south, or east of the Burnside. Few kettle and scour lakes exist on Deanville, Imlay, and Goodland moraines east of the Imlay channel, limiting the number of sites for coring and, thus, for chronologic data. One exception is Petz Lake, located ~6.5 km northwest of Imlay city, which appears to be a kettle on the Goodland Moraine. It is not recommended that geophysical methods like

VES or passive seismic be used in the area due to their inability to differentiate between bedrock and the overlying unconsolidated sediments. Similarly, GPR is not recommended for any future work in the study area due to the high silt content of the soils, which means the signal attenuates at shallow depths of ~1–1.5 m. Any future classifications of soil texture in the area are recommended to be performed using the hydrometer or pipette method for fines and sieving for coarser material so that results can be directly compared to data from the USDA, which uses such methods for textural classifications of soil.

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Appendix A

OSL Lab Results

Table A.1. Original Lab results for OSL data.

1 2  4 6 Sample ID N M/m t Dc Equivalent Dose Rate Age Age Dose5 (Gy/ka) (ka) Uncert.7 (Gy) (ka)

UT1703JL 84/96 1.01 0.347 1.2% 76.139 ± 1.084 ± 70.3 6.9 2.879 0.098 UT1704JL 92/95 1.04 0.258 3.3% 16.141 ± 1.579 ± 10.2 0.9 0.443 0.138 UT1705JL 94/96 1.00 0.362 3.4% 15.507 ± 1.103 ± 14.0 1.4 0.580 0.099 UT1706JL 93/96 1.00 0.380 0.7% 16.276 ± 1.095 ± 14.9 1.4 0.642 0.094 UT1707JL 91/94 1.04 0.443 2.3% 56.104 ± 0.926 ± 60.6 6 2.607 .081

[1] No. of aliquots used for OSL De calculation / no. of aliquots from which OSL data was collected (filtering criteria given in Lepper et al., 2003). [2] Mean/median ratio: a measure of dose distribution symmetry/asymmetry (see supplement to Lepper et al., 2007). [3] Total dose distribution data dispersion (Std. dev./Mean). [4] Dose recovery fidelity (refer to "check dose" in Lepper et al., 2000 and supplement to Lepper et al., 2007). [5] Equivalent doses are based on the mean and std. err. of the OSL

De distribution. [6] Dose rates calculated following the methods described in Aitken (1985; 1998) and Prescott and Hutton (1988; 1994). [7] Fully-propagated age uncertainty (Append. B, Aitken, 1985).

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Table A.2. Elemental Analysis for OSL data.

Field Lab depth (m) H2O (%) K Rb Th U Sample Sample concentrati concentrati concentrati concentrati ID ID on (ppm) on (ppm) on (ppm) on (ppm)

BSG A1 UT1703 1.86 10 ± 3 7866.43 ± 24.71 ± 2.26 ± 0.55 ± JL 799.88 3.38 0.21 0.05

JTEA UT1704 1.05 10 ± 3 12390.25 45.40 ± 2.44 ± 0.93 ± JL ± 1178.76 4.62 0.23 0.08

ICC 1 UT1705 2.95 30 ± 3 11081.91 33.86 ± 1.70 ± 0.74 ± JL ± 1074.76 2.58 0.152 0.05

ICB 1B UT1706 4.00 30 ± 3 11052.17 34.04 ± 2.48 ± 0.56 ± JL ± 1002.79 3.96 0.23 0.06

BSG B1 UT1707 10.20 12 ± 3 7713 ± 27.39 ± 1.53 ± 0.63 ± JL 752 3.88 0.14 0.05

INAA Irradiations were performed at The Ohio State University Nuclear Reactor Lab. INAA data reduction was carried out by Scientific Consulting Services, Dublin, OH.

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Appendix B

Particle Size Analysis

Table B.1. Results of Mastersizer Particle Size analysis (Modified Wentworth Scale). Sample number correlates to those in Figure 3.19.

Sample Site Depth %Sand %Silt %Clay USDA East/ %> 2 # (m) (0.062- (0.004- (<0.0002- Soil West mm 2 mm) 0.062 0.004 Texture Channe mm) mm) l 1 HASMU 3 37 53 10 clay loam E 27 XBRWN CTY1 2 HASVE 2.2 35 55 10 loam E 19 LTMAR SH1 3 HASFO 3.4 19 71 10 loam E 0 STBRW NCTY1 4 CORSL 2.58 75 21 4 muck E 5 2A (Diamict on) 5 HASWD 3.9 27 60 12 sandy E 15 C(3) loam 6 HASWD 0.7 56 36 8 Sand E 8 B(2) 7 HASWD 1.95 20 67 13 clay loam E 17 A(1) 8 HASWD 1.6 39 51 10 sand/ E 3 E gravel 9 HASMU 3.4 15 70 15 loam E 1 XBROO K1 10 HASAN 4.4 16 55 29 loam E 4 KLINCk 1 11 HASAN 2.6 30 41 28 loam W 37 KCLRL K1 12 CORSL 2.48 29 55 16 Muck E 1 2A (Silt)

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Appendix C

Ground Penetrating Radar

Results of ground penetrating radar survey at Goodland Cemetery not shown in Figure 3.26. GPRMS1C (A) runs parallel to the strandline and GPRMS3C (B) runs perpendicular.

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Appendix D

Vertical Electrical Soundings

IX1D results for VESIC6C (A), VESIC7C (B), VESIC8C(C), VESIC9C (D) in Figure 3.23B.

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Appendix E

Tromino Stations

Scans of Tromino data processed using Grilla software for TRMIC11A (A) and TRMIC12A (B) in Figure 3.23C.

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Appendix E

Field Locations

Table E.1. Field locations for all sites in this study.

SiteID Feature or Purpose Lat. Long. North. East. Year Ele. Ele. Lithology (North) (West) (UTM) (UTM) (m) (ft.) Z: 17N Z: 17N COR Clear Lake Sediment 43.14129 83.11706 4778679 327835 2017 252 828 CL1B core COR Imlay Sediment 43.15226 83.06769 4779797 331880 2017 248 814 ICA Channel core COR Imlay Sediment 43.16192 83.09557 4780926 329640 2017 249 816 ICB Channel core COR Imlay Sediment 43.15622 83.07865 4780259 331000 2017 250 819 ICC Channel core COR Scoured Sediment 43.18181 83.04388 4783031 333897 2017 263 862 SL2A Feature core HAS Loamy Till Hand 43.18947 83.08931 4783973 330225 2018 266 874 LINCKA Auger/ Shovel HAS Alluvium Hand 43.14361 83.09307 4778888 329792 2018 278 911 STINEA Auger/ Shovel HAS Clay Till Hand 43.20164 83.0175 4785181 336093 2018 257 845 BCTYA Auger/ Shovel HAS Loamy Till Hand 43.18063 83.02399 4782861 335510 2018 260 853 BCTY1B Auger/ Shovel HAS Clay Till Hand 43.18202 83.02476 4783017 335451 2018 257 844 BCTY2B Auger/ Shovel HAS Loamy Till Hand 43.2306 83.11128 4788586 328555 2018 252 828 MARSHA Auger/ Shovel HAS Loamy Till Hand 43.15254 83.09185 4779878 329916 2018 274 900 CLRLK Auger/ RDA Shovel HAS Imlay Hand 43.12415 83.05778 4776656 332609 2017 244 800 ICA Channel Auger/ Shovel HAS Maumee Hand 43.09324 83.07125 4773249 331429 2017 255 836 MS1B Strandline Auger/ (Armstrong Shovel Rd.)

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HAS Maumee Hand 43.09307 83.07201 4773232 331367 2017 257 845 MS2B Strandline Auger/ (Armstrong Shovel Rd.) HAS Loamy Till Hand 43.23009 83.0032 4788313 337331 2018 248 813 BROOKA Auger/ Shovel HAS Alluvium Hand 43.22435 82.99916 4787668 337644 2017 245 802 CADEA Auger/ Shovel HAS Weaver Hand 43.20162 83.03964 4785223 334295 2016 251 823 WDA Drain Auger/ Shovel HAS Weaver Hand 43.19666 83.0375 4784668 334455 2016 251 824 WDB Drain Auger/ Shovel HAS Weaver Hand 43.20037 83.03944 4785084 334308 2016 252 827 WDC Drain Auger/ Shovel HAS Road Cut Out Crop 43.19678 83.04271 4784691 334032 2017 254 832 WDE OSL Kame OSL 43.18575 83.04975 4783481 333430 2017 304 998 BSGA OSL Kame OSL 43.18654 83.05584 4783581 332937 2017 282 925 BSGB OSL Alluvium OSL 43.1564 83.07046 4780262 331666 2017 251 824 JTEA VES Imlay VES 43.15416 83.05165 4779977 333190 2017 269 881 IC6C Channel VES Imlay VES 43.15185 83.0743 4779766 331342 2017 250 821 IC7C Channel VES Imlay VES 43.1518 83.07066 4779753 331638 2017 247 810 IC8C Channel VES Imlay VES 43.15404 83.06878 4779998 331797 2017 251 823 IC9C Channel TRM Imlay Tromino 43.15286 83.0709 4779872 331621 2017 247 812 IC11A Channel TRM Imlay Tromino 43.15292 83.06231 4779872 331471 2017 248 813 IC12A Channel GPR Strandline GPR 43.09324 83.07125 4773251 331429 2017 258 845 MSC

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