Upper Jurassic-Lower Cretaceous Palynostratigraphy of the Aklavik Range, Northwest Territories, Arctic Canada

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2019-04-26 Upper Jurassic-Lower Cretaceous palynostratigraphy of the Aklavik Range, Northwest Territories, Arctic Canada

Nguyen, Anne Van

Nguyen, A. V. (2019). Upper Jurassic-Lower Cretaceous palynostratigraphy of the Aklavik Range, Northwest Territories, Arctic Canada (Unpublished master's thesis). University of Calgary, Calgary, AB. http://hdl.handle.net/1880/110242 master thesis

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UNIVERSITY OF CALGARY

Upper Jurassic-Lower Cretaceous palynostratigraphy of the Aklavik Range, Northwest

Territories, Arctic Canada

Anne Van Nguyen

A THESIS

SUBMITTED TO THE FACULTY OF GRADUATE STUDIES

IN PARTIAL FULFILMENT OF THE REQUIREMENTS FOR THE

DEGREE OF MASTER OF SCIENCE

GRADUATE PROGRAM IN GEOLOGY AND GEOPHYSICS

CALGARY, ALBERTA

APRIL, 2019

ABSTRACT

Quantitative palynostratigraphy of Lower Cretaceous rocks from the Aklavik Range, Northwest

Territories, is used to provide insight into paleoenvironmental conditions in the Boreal Realm during the Jurassic-Cretaceous transition. Paleoenvironmental reconstruction of this time interval is based on palynoassemblages preserved in the Husky Formation (upper Tithonian – lower

Berriasian). Relative abundance of ecologically important spore and pollen taxa such as

Cupressaceae-Taxaceae and Classopollis classoides pollen reveal increasing humid conditions with a seasonally arid phase during the early Berriasian. Dinoflagellate cyst assemblages preserved in the purported Martin Creek Formation (lower Berriasian) include potential biostratigraphically significant species Oligosphaeridium cf. tenuiprocessum. A revised late Albian age is proposed for the succession, which may be stratigraphically attributed to the Arctic Red Formation. High relative abundance and diversity of fern spores indicate that high moisture conditions prevailed.

Trends in the relative abundance of gymnosperm pollen suggest a transition from cool to warm temperatures during this time period.

ACKNOWLEDGMENTS

I would first and foremost like to extend my appreciation to my research supervisor, Dr. Jennifer

Galloway, who has been a constant source of support and encouragement. Thank you for being such an exemplary role model, for encouraging me to step out of my comfort zone and become a better scientist. Without you, I would have never been able to experience so many incredible things during my graduate studies. I would also like to thank my co-supervisor, Dr. Benoit Beauchamp, who I hold in great esteem, for his support and guidance, fun anecdotes and overall positive attitude. I am also grateful to the many people at the Geological Survey of Canada, particularly

Dr. Thomas Hadlari for helping me collect samples on the steeper slopes I was wary about climbing; Dr. Larry Lane for being so patient in explaining the geology and logistics of the study area; Dr. Terry Poulton for his advice on the biostratigraphy, and who was so open and willing to help with this thesis; Dr. Rob Fensome for teaching me about dinoflagellate cysts and patiently helping with species identification when I struggled with them; Dr. Manuel Bringué who helped with one of my figures; and Leanne Tingley who processed my samples despite the long period of construction at the GSC. I would also like to thank Drs. Alex Dutchak and Charles Henderson for agreeing to be a part of my review committee and for providing useful feedback. Thank you as well to my colleagues from the Arctic Geoscience Laboratory at the University for making the windowless lab room seem less dreary, and to all the professors and fellow graduate students with whom I worked as a teaching assistant, who showed me how fun teaching could be.

iii

DEDICATION

For

My beloved family

and

Connie, Emmée, and Eli – proof that mind melds can work long-distance

TABLE OF CONTENTS

ABSTRACT ...... ii ACKNOWLEDGMENTS ...... iii DEDICATION ...... iv LIST OF TABLES ...... viii LIST OF FIGURES ...... ix LIST OF PLATES ...... x CHAPTER 1: INTRODUCTION ...... 1 History of the Jurassic and Cretaceous Periods ...... 1 Defining the Jurassic-Cretaceous boundary ...... 2 Nomenclature and Placement of the J/K boundary ...... 8 The Boreal Realm of Canada ...... 9 Thesis Structure and Role of Student ...... 18 References ...... 21 CHAPTER 2: PALYNOSTRATIGRAPHY OF THE JURASSIC-CRETACEOUS TRANSITION IN THE NORTHERN RICHARDSON MOUNTAINS OF THE NORTHWEST TERRITORIES, ARCTIC CANADA ...... 29 ABSTRACT ...... 29 INTRODUCTION ...... 30 GEOLOGIC SETTING AND STUDY AREA ...... 34 Beaufort-Mackenzie Region ...... 37 Husky Formation ...... 42 Thermal Maturity ...... 46 Ammonite and Buchia Biostratigraphy in Arctic Canada ...... 47 Previous Palynological Studies on the Jurassic-Cretaceous Boundary in Arctic Canada ..... 50 METHODS ...... 55 Palynology ...... 55 Multivariate Statistical Techniques ...... 58 RESULTS ...... 60 Spores and Pollen ...... 60 Dinoflagellate Cysts ...... 61 Multivariate Statistical Analyses: Q- and R-mode Cluster Analysis, CONISS ...... 67 Detrended Correspondence Analysis (DCA) ...... 73

DISCUSSION ...... 76 Biostratigraphy ...... 76 Dinoflagellate Cysts ...... 76 Spores and Pollen ...... 81 Paleoecology ...... 85 Bryophytes and Lycophytes ...... 85 Ferns ...... 86 Conifers ...... 87 Cycads, Ginkgos, and Gnetophytes ...... 90 Paleoenvironment ...... 91 Q-mode Cluster Analysis and Detrended Correspondence Analysis (DCA) ...... 91 R-mode Cluster Analysis ...... 93 Stratigraphically Constrained Cluster Analysis (CONISS) ...... 95 CONCLUSION ...... 100 REFERENCES ...... 101 CHAPTER 3: LOWER CRETACEOUS PALYNOSTRATIGRAPHY OF THE AKLAVIK RANGE, NORTHWEST TERRITORIES, ARCTIC CANADA ...... 119 ABSTRACT ...... 119 INTRODUCTION ...... 120 GEOLOGIC SETTING AND STUDY AREA ...... 123 Husky Formation ...... 126 Martin Creek Formation ...... 131 McGuire Formation ...... 132 Kamik Formation ...... 133 Mount Goodenough Formation ...... 134 Rat River Formation ...... 135 Rapid Creek Formation and Albian Flysch ...... 136 Boundary Creek Formation ...... 136 METHODS ...... 137 Palynological Samples ...... 137 Multivariate Statistical Techniques ...... 140 RESULTS ...... 143 Dinoflagellate Cysts ...... 143

Spores and Pollen ...... 148 Cluster Analyses: Q-, R-mode and CONISS ...... 152 Detrended Correspondence Analysis...... 158 DISCUSSION ...... 158 Biostratigraphy ...... 158 Dinocyst Zonations ...... 158 Age Estimation ...... 165 Paleoecology ...... 174 Bryophytes, Lycophytes and Monilophytes ...... 174 Cycads and Ginkgos ...... 176 Conifers ...... 176 Gnetophytes...... 178 Paleoenvironment ...... 179 Q-mode Cluster Analysis ...... 179 Detrended Correspondence and R-mode Cluster Analyses ...... 180 Stratigraphically Constrained Cluster Analysis (CONISS) ...... 186 Angiosperms ...... 188 CONCLUSION ...... 190 REFERENCES ...... 191 CHAPTER 4: SUMMARY AND FUTURE WORK ...... 214 SUMMARY ...... 214 FUTURE WORK ...... 217 REFERENCES ...... 219 PLATES ...... 224 SPORES AND POLLEN ...... 225 DINOFLAGELLATE CYSTS ...... 243 UNKNOWN PALYNOMORPHS ...... 261 APPENDICES ...... 263 Appendix 1: Count data from the Husky Formation at the Martin Creek section ...... 264 Appendix 2: Count data from the Treeless Creek section ...... 265 Appendix 3: RStudio scripts used for quantitative statistical analyses ...... 267 Appendix 4: DCA site ordination of Treeless Creek section ...... 268

vii

LIST OF TABLES

Chapter 2 Table 1: Biological nomenclature and taxonomic authority of identified spores and pollen grains in the Husky Formation ...... 62 Table 2: Biological nomenclature and taxonomic authority of identified dinoflagellate cysts in the Husky Formation ...... 65

Chapter 3 Table 1: Biological nomenclature and taxonomic authority of identified dinoflagellate cysts in the Treeless Creek section...... 146 Table 2: Biological nomenclature and taxonomic authority of identified spores and pollen grains from the Treeless Creek section...... 149 Table 3: Paleontological reports and age estimations based on fossil assemblages collected from various localities in the Treeless Creek area ...... 168

viii

LIST OF FIGURES

Chapter 1 Figure 1: Current markers for the J/K boundary in the Tethyan Realm and various areas in the Boreal Realm ...... 6 Figure 2: Polar map of the distribution of Boreal and Tethyan seas in Canada and estimated limits of the Boreal Realm worldwide ...... 10 Figure 3: Simple cladogram of plant evolution ...... 14

Chapter 2 Figure 1: Plate reconstruction at 145 Ma, and estimated southern limits of the Boreal Realm during the J/K transition ...... 35 Figure 2: Maps of the Beaufort-Mackenzie region and Martin Creek study area ...... 38 Figure 3: Modified lithostratigraphy of the northeastern Richardson Mountains ...... 40 Figure 4: Photograph of the study area, Martin Creek ...... 43 Figure 5: Uppermost Jurassic - Lower Cretaceous biostratigraphic correlation chart ...... 48 Figure 6: Lithostratigraphic column of the composite section measured at Martin Creek...... 56 Figure 7: Q-mode cluster analysis ...... 68 Figure 8: CONISS combined with R-mode cluster analysis ...... 70 Figure 9: DCA plot of Martin Creek section ...... 74 Figure 10: First and last occurrences of identified dinocyst species ...... 77

Chapter 3 Figure 1: Maps of the Beaufort-Mackenzie region and Treeless Creek study area ...... 124 Figure 2: Extensive outcrops of the Husky Formation at Treeless Creek ...... 127 Figure 3: Lower Cretaceous lithostratigraphy of north and northwestern Canada ...... 129 Figure 4: Lithostratigraphic column and composite image of Treeless Creek section ...... 138 Figure 5: Marine-to-terrestrial ratio of the Treeless Creek section and relative abundance of dinocysts identified to the species level ...... 144 Figure 6: Q-mode cluster analysis ...... 153 Figure 7: CONISS combined with R-mode cluster analysis ...... 155 Figure 8: DCA plot of Treeless Creek section and coastal wetland profile ...... 159 Figure 9: First and last occurrences of identified dinocyst species ...... 163 Figure 10: Possible revision of the geology in the Treeless Creek area ...... 172

LIST OF PLATES

Spores and Pollen Plate 1: Divisions Marchantiophyta, Bryophyta ...... 225 Plate 2: Division Tracheophyta, Class Lycopodiopsida, Order Lycopodiales ...... 227 Plate 3: Division Tracheophyta, Class Lycopodiopsida, Orders Selaginellales, incertae sedis ...... 229 Plate 4: Division Tracheophyta, Class Polypodiopsida, Order Cyatheales ...... 231 Plate 5: Division Tracheophyta, Class Polypodiopsida, Orders Gleicheniales, Osmundales . 233 Plate 6: Division Tracheophyta, Class Polypodiopsida, Orders Polypodiales, Schizaeales, incertae sedis ...... 235 Plate 7: Division Tracheophyta, Class incertae sedis ...... 237 Plate 8: Division Tracheophyta, Classes Cycadopsida, Ginkgoopsida, Gnetopsida ...... 239 Plate 9: Division Tracheophyta, Class Pinopsida ...... 241

Dinoflagellate Cysts Plate 10: Order Gonyaulacales, Family Ceratiaceae ...... 243 Plate 11: Order Gonyaulacales, Family Gonyaulacaceae, Subfamilies Cribroperidinioideae, Gonyaulacoideae ...... 245 Plate 12: Order Gonyaulacales, Family Gonyaulacaceae, Subfamily Gonyaulacoideae ...... 247 Plate 13: Order Gonyaulacales, Family Gonyaulacaceae, Subfamily Leptodinioideae ...... 249 Plate 14: Order Gonyaulacales, Family Gonyaulacaceae, Subfamily Leptodinioideae ...... 251 Plate 15: Order Gonyaulacales, Family Gonyaulacaceae, Subfamily Leptodinioideae ...... 253 Plate 16: Order Gonyaulacales, Family Gonyaulacaceae, Subfamily Leptodinioideae ...... 255 Plate 17: Order Gonyaulacales, Family Pareodiniaceae ...... 257 Plate 18: Order Gonyaulacales, Family incertae sedis...... 259

Unknown Palynomorphs Plate 19: Unknown palynomorphs ...... 261

CHAPTER 1: INTRODUCTION

History of the Jurassic and Cretaceous Periods

The Cretaceous Period was first named by J.M.J. d’Omalius d’Halloy in 1822 as “terrains crétacés” in reference to the extensive Upper Cretaceous chalk (creta, in Latin) deposits in NW

Europe (Cope, 2008; Gradstein et al., 2012). The concept of the Jurassic Period originated from the Jura Mountains in the Western Alps: Alexander von Humboldt first used the term “Jura

Kalkstein” in 1799 to describe carbonate shelf deposits in the Jura Mountains, although he mistakenly thought they were older in age (Gradstein et al., 2012). Alexandre Brongniart subsequently used the term “terrains jurassiques” in 1829 when correlating the “Jura Kalkstein” facies to the Lower Oolite Series in Britain (Cope, 2008; Gradstein et al., 2012). Definitions of the

J/K boundary were not considered at the time of conception of the two Mesozoic systems. In the

1840’s and 1850’s, Alcide d’Orbigny proposed the division of the Jurassic and Cretaceous into stages using ammonite and other fossil assemblages in France, with the Jurassic having 10 stages and the Cretaceous having five (Cope, 2008; Gradstein et al., 2012). While many of these stages are still used, and in some cases formalized, today, definitions have changed from their inception

(Gradstein et al., 2012).

The chronostratigraphic framework established by d’Orbigny (1840, 1842, 1847, 1849-

1850, 1849-1851, 1852) established many of the stages of the Jurassic and Cretaceous, which were subsequently modified by Alfred Oppel (1856-1858), who subdivided the Jurassic into biostratigraphic zones using ammonite species as zonal indices, successfully correlating Jurassic strata in northwestern Europe and formalizing the eleven Jurassic stages seen today (Cope, 2008;

Gradstein et al., 2012). The Cretaceous was originally divided into five stages, with the Neocomian as the basal stage. This was further subdivided into the Berriasian, Valanginian and Hauterivian

stages (Barbier and Thieuloy, 1965; Gradstein et al., 2012). Zonal division using ammonite species of the Cretaceous followed after the success of Oppel’s biostratigraphic zones in the Jurassic

System (Cope, 2008).

Defining the Jurassic-Cretaceous boundary

A chronostratigraphic boundary is defined using its base, which requires finding a standard to which all other sections can be compared and correlated (Cope, 2008). These standards are called Global Stratotype Sections and Points (GSSPs) and are ratified by the International Union of Geological Science (IUGS). Consequently, definition of the J/K boundary is dependent on defining the base of the Cretaceous and therefore the base of the Berriasian Stage. However,

GSSPs have not been established for all the Cretaceous stages: according to the International

Subcommission on Cretaceous Stratigraphy of the IUGS, only five out of the twelve stages have ratified GSSPs. Establishing GSSPs for the stages of the Lower Cretaceous is particularly problematic; only the Albian Stage has a recently ratified GSSP out of the six Lower Cretaceous subdivisions (Kennedy et al., 2017). The Berriasian Stage was originally defined based on a limestone succession near Berrias in southeast France, but the lower boundary of this stratotype lacked a major faunal turnover (Gradstein et al., 2012). To circumvent this problem, the base of the ammonite subzone Berriasella jacobi Mazenot, 1939 was decided as the standard for defining the Berriasian Stage in 1975 (Gradstein et al., 2012). However, a recent study suggests that most specimens identified as B. jacobi in past literature may be misidentified, making its use as a biostratigraphic marker untenable (Frau et al., 2016). Since 2016, the primary marker for the J/K boundary has been shifted to the base of the calpionellid Calpionella alpina Lorenz, 1902 Subzone, as discussed in further detail below.

Defining the J/K boundary is highly problematic due to a variety of factors, including distinct biogeographical provincialism, an absence of a major faunal turnover at the boundary, the lack of reference sections caused by the “Purbeckian regression”, and a general lack of consensus on the chronostratigraphic framework for the transition (Remane, 1991). Biogeographical provincialism refers to the isolation of geographical regions, causing the proliferation of endemic faunas and floras. Endemism was pervasive during the Late Jurassic and Early Cretaceous due, in part, to the rearrangement of continents during the Mesozoic associated with the break-up of

Pangea and the “Purbeckian regression” that further isolated regions (Remane, 1991). Other causes of provincialism may include temperature, geographic barriers, salinity, and environmental instability (Hallam, 1975). This provincialism resulted in the division of the world into broad geographic realms generally referred to as the Austral (southern), Tethyan (equatorial), and Boreal

(northern) realms, wherein each realm had distinct groups of ammonites and a lack of shared species. Most studies researching a proper definition of the J/K boundary are focused on the

Tethyan Realm as J/K strata are well represented in the area and are therefore logistically easier to study. The Tethyan Realm’s larger geographic area contributes to a greater general diversity as well as a greater ammonite diversity (Hallam, 1975). In comparison, the Boreal paleobiogeographic realm is limited to the northern part of the Northern Hemisphere and is primarily defined by the occurrence of much fewer ammonite families and subfamilies (Hallam,

1975). Biostratigraphic correlation is further complicated by the complex nature of each realm and the endemic faunas of separate regions. The Boreal Realm, for instance, could have at least six different ammonite or bivalve zonations: the Arctic, Western Canada, East Greenland, Russia,

Siberia, and the North Atlantic (Wimbledon et al., 2011).

Definition of the J/K boundary has therefore shifted focus from endemic ammonites to other potential biostratigraphic markers. In 2011, the Berriasian Working Group (WG) of the

International Subcommission on Cretaceous Stratigraphy (ISCS) proposed these potential primary markers for the J/K boundary in the Tethyan Realm: 1) the base of the Calpionella Lorenz, 1902

Zone – alpina Subzone; 2) the spike in abundance of Calpionella alpina; 3) the first appearance datum (FAD) of the nannoconid species, Nannoconus steinmannii minor Deres and Achéritéguy,

1980 and Nannoconus kamptneri minor Bralower et al., 1989; and 4) the base of the magnetozone

M18r (Wimbledon et al., 2011). Potential markers for the J/K boundary are primarily biostratigraphic and/or magnetostratigraphic as the J/K transition lacks significant chemostratigraphic events and isotope excursions (Price et al., 2016; Wimbledon, 2017).

Calpionellids have been increasingly considered as the most useful biostratigraphic marker in defining the J/K boundary. However, calpionellid zones are not correlative with the boundaries of any ammonite zones. Despite this, only calpionellids are known to exhibit a widespread and consistent turnover from Crassicollaria and large Calpionella to smaller, round Calpionella alpina, which is preserved in the rock record (Wimbledon et al., 2013; Schnabl et al., 2015;

Wimbledon, 2017). This turnover occurs within the magnetozone M19n.2n, which has recently been considered in place of magnetozone M18r. While M18r is easily identifiable, it does not correlate with any biostratigraphic markers (Wimbledon et al., 2013; Schnabl et al., 2015). As of

2016, the base of the Alpina Subzone has been considered as the primary marker for the J/K boundary, although recent studies (e.g., Pruner et al., 2010; Wimbledon et al., 2013) have shown that the Jacobi Subzone occurs within the Crassicollaria intermedia Durand Delga, 1957 Subzone rather than within the Calpionella alpina Subzone (Frau et al., 2016; Schneider et al., 2018). The

FADs of the calcareous nannofossils Nannoconus steinmannii minor and N. kamptneri minor were

thought to occur directly below and above the Alpina Subzone boundary, respectively, but recent research suggests some modification to their age ranges. While in some places (e.g., Theodosia,

Puerto Escano, Kurovice and Rio Argos) the FAD of N. steinmannii minor is known to occur within magnetozone M19n.2n, the age range for N. steinmannii minor appears mainly limited to magnetozones M19 to M17r (Wimbledon, 2017). Likewise with Nannoconus kamptneri minor: its

FAD has been recorded in mid M19n.2n (e.g., Theodosia), but its age range is primarily restricted between M19n and M17r (Wimbledon, 2017). The FAD of Hexalithus strictus (H. geometricus)

Bergen, 1994 has been considered as a marker instead due to its restricted stratigraphic distribution, occurring solely within magnetozone M19n.2n and encompassing the base of the

Alpina Subzone (Wimbledon, 2017).

Despite these advances in research, all the aforementioned markers for the Tethyan J/K boundary are not preserved in Boreal strata (Harding et al., 2011). Currently, defining the J/K boundary in the Boreal Realm depends mainly on ammonites, bivalves and belemnites. In terms of ammonite zonation, the basal Cretaceous is encompassed by the Subcraspedites preplicomphalus Swinnerton, 1935 Zone in the North Sea; in the Russian Platform, the J/K boundary lies near the base of the Craspedites nodiger Eichwald, 1868 Zone; in Northern Siberia, the boundary lies near the base of, Craspedites taimyrensis Shul’gina, 1962 Zone (Rogov and

Zakharov, 2009; Fig. 1). Due to the endemism of ammonite species, and their paucity in some regions, bivalves are commonly used where ammonites are scarce, as is the case in northern

Canada. The J/K boundary in northern Canada is primarily defined by the base of the Buchia okensis Pavlow, 1907 Zone, which also occurs throughout the Boreal Realm such as in Northern

Siberia, eastern Greenland, northeastern Russia, etc. (Zakharov, 1987; Urman et al., 2014).

Figure 1: Current markers for the J/K boundary in the Tethyan Realm and various areas in the

Boreal Realm (modified from Davies, 1983 and Wimbledon, 2017).

In general, biostratigraphic correlation of the J/K boundary in a global context, required for ratification of the GSSP, depends first on defining the boundary in the Tethyan Realm, and subsequently requires matching berriasellid ammonites, calpionellids and radiolarians from the

Tethyan Realm with craspeditid ammonites, buchiid bivalves, belemnites, and possibly other fossil groups from the Boreal Realm (Hedinger, 1979; Wimbledon et al., 2011; Georgescu and Braun,

2013; Vishnevskaya, 2017). Currently, the definition of the boundary between the Jurassic and

Cretaceous systems remains unresolved worldwide, with the base Cretaceous System being the only Phanerozoic system boundary that remains undefined in the Geologic Time Scale.

Nomenclature and Placement of the J/K boundary

The difficulties in defining the J/K stages coupled with the marked provincialism of ammonite assemblages led to much debate among the scientific community wherein different stage and substage nomenclature were used depending on the region. Prior to the acceptance of the

Tithonian Stage as the uppermost Jurassic and the Berriasian Stage as the basal Cretaceous, three different sets of stages were used to refer to Upper Jurassic – Lower Cretaceous strata: the

Portlandian and Purbeckian stages were used in northwestern Europe (d’Orbigny, 1842-1851); the

Volgian and Ryazanian stages for Russia, Poland, and parts of the Arctic (Nikitin, 1881;

Bogoslovsky, 1897); and Tithonian and Berriasian stages for the area of the Tethys Ocean (Cope,

2008; Gradstein et al., 2012). The Ryazanian and Volgian stages were often used in literature pertaining to the Boreal Realm, and still are to some degree. These stages are approximately equivalent to the middle – late Berriasian and early Tithonian – early Berriasian, respectively

(Harding et al., 2011). Although these terms are useful for studies within the Boreal Realm, they also present a problem in terms of correlation as the Volgian/Ryazanian boundary lies above the

J/K boundary in the Tethyan Realm (Cope, 2008).

There still exists some controversy in terms of defining the stages of the Jurassic and

Cretaceous. Some researchers argue for the J/K boundary to be placed at the Berriasian-

Valanginian boundary as originally proposed by d’Orbigny and Oppel in the 19th century (Granier,

2019); others argue for the reintegration of the Volgian Stage in place of the Tithonian, effectively placing it wholly within the Jurassic Period in the General Stratigraphic Scale of Russia

(Vishnevskaya, 2017).

The Boreal Realm of Canada

The precise geographic limits of the Boreal Realm at a given time cannot be determined as the boundary is gradational and changed throughout time. In the Callovian and early Oxfordian,

Boreal ammonites extended into southern Europe, but subsequently returned to northern Europe by the late Oxfordian (Hallam, 1975). Unlike these migratory ammonites, Bathonian and upper

Bajocian ammonites were mostly restricted to the Arctic (Hallam, 1975). In general, the southern boundary of the Boreal Realm lies through southern Europe, the North Pacific Region, and between Japan and eastern Siberia (Hallam, 1975; Fig. 2). Biota with Boreal affinities in the

Canadian Arctic are part of the North American Province of the Boreal Realm, otherwise known as the Chukchi-Canadian biotic province (Jeletzky, 1973; Saks, 1975). During the latest Jurassic to Valanginian, the Canadian portion of this province was restricted to the Canadian Arctic

Archipelago and parts of Canadian Arctic mainland, mainly the Beaufort-Mackenzie area of

Yukon and the Northwest Territories (Jeletzky, 1973). Simultaneously, the Tethyan Realm (North

Pacific Province of the Tethyan Realm) consisted of the Western Cordillera of Canada and the

Peace River region (Jeletzky, 1973). These two realms were eventually connected by the Western

Interior Seaway in the Late Cretaceous and contributed to the faunal intermingling of previously endemic invertebrates (Jeletzky 1973, 1984).

Figure 2: Polar map of the distribution of Boreal and Tethyan seas in Canada and estimated limits of the Boreal Realm worldwide (modified from Jeletzky, 1971, 1973; Saks, 1975; Zakharov,

2012). Red square denotes the region of interest and, within it, Jeletzky’s (1971, 1973) estimated extent of Boreal seas in the Late Jurassic-Early Cretaceous based on the distribution of ammonite and Buchia species. Maximum extent of Berriasian Boreal seas may have included all purple areas, which is consistent with Zakharov’s (2012) interpretation. Base map created from ODSN website

(http://www.odsn.de).

Palynology and Plant Evolution

Palynology is the study of palynomorphs such as fossil spores, pollen grains and dinoflagellate cysts (i.e., dinocysts). Palynomorphs are ubiquitous in Mesozoic strata, their fossils occurring in both Boreal and Tethyan realms in high abundances. In addition, they have a high preservation potential that allows them to be easily collected for statistical analyses. Pollen analysis is dependent on the assumption of uniform dispersal of a high abundance of palynomorphs over a large area (Fægri and Iverson, 1989). Assuming palynomorphs are evenly distributed, they are therefore all comparable and the quantity of individual species can be expressed as a percentage of a total sum (Fægri and Iverson, 1989). The absence of such microfossils may therefore indicate that the parent plant did not occur locally (Fægri and Iverson, 1989). Despite complex taphonomic processes, the relative abundance of palynomorphs preserved in sediment has been shown to accurately reflect the distribution of onshore vegetation (e.g., Muller, 1959; Heusser and Balsam,

1977; Mudie, 1982; Heusser, 1983; Hooghiemstra et al., 1986; Mudie and McCarthy, 1994; Sun et al., 1999; van der Kars, 2001; van der Kaars and De Deckker, 2003; Hooghiemstra et al., 2006;

Montade et al., 2011; Luo et al., 2014; Zhao et al., 2016). The distinct morphology of palynomorphs allows for the identification of different species and can subsequently be used to provide information on Mesozoic depositional environments, paleoclimates, and past terrestrial and marine ecosystems in which they occurred. These inferences are based on the ecological tolerances of the parent plants from which they were derived. Palynomorphs are not as affected by the endemism pervasive in ammonite occurrences during the Mesozoic due to the warm equable climate characteristic of the Cretaceous Period (Hallam, 1985; Fægri and Iverson, 1989; Hay,

2008). Palynomorphs therefore have the potential to be used for biostratigraphic correlation between the Boreal and Tethyan realms to define the J/K boundary.

Reconstructing paleoenvironmental and paleoclimatic conditions using palynomorphs is dependent on a solid understanding of the parent plants and protists from which they were derived

(Fig. 3). The evolution of these previously existing parent plants or protists likely began with cyanobacteria in the Precambrian, blue-green photosynthetic prokaryotes that were the constituents of the microbial mats that gave rise to stromatolites. Prokaryotes are the most ancient organisms on Earth, existing as simple cells that lack a nuclear envelope and complex chromosomes. The Great Oxygenation Event led to the evolution of the first eukaryotic cells 2.1 billion years ago: eukaryotes had cells with nuclear envelopes, complex chromosomes and membrane-bound organelles (Raven et al., 2013). Eukaryotes include algae, a generalized term that can encompass euglenoids, dinoflagellates, heterokonts, red algae, and green algae.

Dinoflagellates are unicellular eukaryotes with two flagella that are used for motion. One of the flagella encircles the circumference of the dinoflagellate, having a ribbon-like appearance, and the other lies perpendicular to it, allowing the dinoflagellate to spin as it moves in the water column (Evitt, 1985; Raven et al., 2013). Extant dinoflagellates are mostly marine and haploid, their reproductive strategy being predominantly asexual. Reproduction is by longitudinal cell division wherein each daughter cell inherits one flagella and part of the theca that forms the cell wall (Raven et al., 2013). A few modern species are haplontic, meaning their life cycle is predominantly haploid with brief instances of diploidy, wherein the dinoflagellate undergoes zygotic meiosis to create zygotes called resting cysts, a dormant stage in their lifecycle induced when environmental conditions are unfavourable (Raven et al., 2013). When conditions improve, the dinoflagellate exits through a hole in the theca, called the archeopyle, with the removed plate referred to as the operculum, and the cyst is abandoned. These cysts are readily preserved due to the resilient nature of dinosporin that make up the cyst walls. Traditionally, the cyst wall was

Figure 3: Simple cladogram of plant evolution highlighting common Phyla seen in this thesis (after

Raven et al., 2013).

considered as having been composed of sporopollenin, but research shows that the organic compound of the resting cyst wall is distinct from sporopollenin and unique to dinoflagellates

(Kokinos et al., 1998). The fossil record is therefore heavily biased towards those dinoflagellate species capable of creating resting cysts, which makes their evolutionary history difficult to recreate. Dinocysts are most common in Upper Triassic rocks and exist to this day to a lesser extent, although dinoflagellates likely existed prior to the appearance of land plants. The identification of dinocysts rely heavily on the placement of the archeopyle-operculum and the number and arrangement of thecal plates (Traverse, 2007), which becomes difficult to determine in older rocks due to the limitations in preservation and a limited number of modern analogues to which species can be compared.

Land plants, or embryophytes, likely evolved from a single ancestor of charophycean green algae in the Late Ordovician (Traverse, 2007). The first assemblage of embryophytes predominantly consisted of bryophytes, fern allies, and ferns. These all produce spores as a part of their reproductive strategy. Spores are typically haploid and unicellular, and can grow into intermediate, multicellular haploid plants in the life cycle of the organism called gametophytes.

Gametophytes have reproductive parts that produce sperm and eggs (gametes) that combine to form a diploid zygote that develops into an adult plant capable of producing spores (sporophyte).

This type of life cycle is referred to as “alternation of generations.” The gametophyte generation is dominant in bryophytes, whereas the sporophyte generation is dominant in fern and fern allies.

Embryophytes later included gymnosperms and angiosperms, the seed plants, which generated pollen grains that contain a reduced male gametophyte that can produce sperm to fertilize an egg in the female structure (e.g., pistil in flowering plants). Unlike spores, the pollen grain’s gametophyte does not grow into a multicellular intermediate body. The then fertilized egg

produces an embryo within a seed that then grows into an adult plant. Embryophytes are differentiated from their algal ancestors by having several characteristics, including the development of male and female gametangia, multicellular embryos, multicellular diploid sporophytes, multicellular sporangia, and the production of sporopollenin, a highly-resistent polymer that constitutes the outer wall (exine layer) of spores and pollen grains (Raven et al.,

2013). Sporopollenin evolved as a response to a threat in dessication, and the exine layer is often sculptured in taxonomically unique ways, permitting differentiation between species.

Bryophytes include all non-vascular plants such as liverworts, mosses, and hornworts. It is important to note that bryophytes do not exclusively comprise those taxa belonging to the phylum

Bryophyta. Only mosses belong to the phylum Bryophyta. Liverworts belong to the phylum

Marchantiophyta, while hornworts belong to Anthocerotophyta. Extant bryophytes lack a true xylem and phloem, which are respectively used for transportation of water and food (Raven et al,

2013). Due to the lack of a true vascular system, bryophytes are generally smaller in size and grow close to the ground. Their reproductive strategy involving production and dispersal of spores is therefore mainly dependent on water transportation.

Tracheophytes are differentiated from bryophytes by the presence of a true vascular system consisting of a xylem and phloem. The presence of a true xylem and phloem allows these plants to grow larger in size than bryophytes, transport fluids and sugars, and develop leaves.

Tracheophytes include lycophytes (fern allies), monilophytes (ferns), gymnosperms (seed plants), and angiosperms (flowering plants). Lycophytes and monilophytes depend on water dispersal for reproductive purposes and can be differentiated by a variety of factors, including the evolution of megaphylls. Lycophytes have microphylls, leaves that only have one leaf vein and are therefore

typically narrow in size. Monilophytes have megaphylls, which have multiple veins that allow them to grow broader leaf shapes.

Gymnosperms are the first group of extant seed plants. Unlike the more advanced angiosperms, gymnosperms develop a “naked” seed, meaning they lack a protective surrounding structure, the fruit. Gymnosperms are all heterosporous: they develop megaspores and microspores that give rise to female and male gametophytes, respectively (Raven et al., 2013). The male gametophyte, the pollen grain, is released when matured for reproduction and then aerially transported to the female gametophyte, which is retained in the sporophyte (i.e., cone) for increased nutrition and protection (Raven et al., 2013). This process is termed pollination, leading to the fertilization of the ovule and the subsequent formation of the seed (Raven et al., 2013). Pollen grains first appeared in the Late Devonian and are defined in palynological terms as the gametophyte that develops within the microspore wall of seed plants and the microspore wall that is ultimately preserved as a microfossil (Traverse, 2007).

Angiosperms are the most reproductively advanced plants to have evolved. They are thought to have originated in the Jurassic or Early Cretaceous, with a closely related predecessor in the Middle Triassic (Hochuli and Feist-Burkhardt, 2013). Unlike gymnosperms, angiosperms have protected seeds that are encased in fruit. Angiosperms now dominate the modern floral assemblage due to their highly successful reproductive strategies involving various forms of pollination, including wind, insect, and animal pollination.

Thesis Structure and Role of Student

This thesis is divided into four chapters: an introductory chapter (Chapter 1), two independent papers addressing the research problem (Chapters 2 and 3), and a concluding chapter

(Chapter 4).

Both papers (Chapters 2 and 3) are co-authored by Jennifer Galloway, Larry Lane, Thomas

Hadlari, Terry Poulton, Rob Fensome, and Benoit Beauchamp. The study areas of both papers are located in the Mackenzie Delta area in the northern Richardson Mountains, Northwest Territories of Canada. The first paper (Chapter 2) is based on the “Martin Creek” section where the J/K boundary is purportedly preserved in the Canadian Arctic. This paper presents a detailed palynological study using multivariate statistical analyses to determine if a palynological signature could be used to distinguish the J/K transition. The second paper (Chapter 3) is focused on Lower

Cretaceous palynology in the “Treeless Creek” section using similar methods of pollen analysis and uses dinoflagellate cysts for age determination of the formation from which samples were collected.

My role as the student was to travel to the study areas to collect mudstone samples from the formations that purportedly preserved the J/K boundary. I assisted with sample processing for palynological analyses at the Geological Survey of Canada in Calgary. Using the resulting microscope slides, I compiled quantitative data by identifying spores, pollen grains, and dinocysts to the genus or species level. I then used this data to calculate relative abundance, and conducted all multivariate statistical analyses through use of previously existing computer programs.

Compilation, synthesis, and interpretations of the data were completed by me in consultation with my supervisor, Dr. Jennifer Galloway.

Galloway provided advice, reviews and edits, and scientific support relating to palynology, and is the official co-supervisor of this project and student. Lane assisted in selection of the study site for detailed stratigraphic research and provided insight into the structural geology and tectonics of the area. Fensome provided oversight in the identification of dinocysts found in the collected samples. Hadlari provided lithostratigraphic interpretation and assisted in sample collection from

the Martin Creek section. Poulton identified the macrofossils found at the sites and assisted in formulation of the macrofossil-based biostratigraphy. Beauchamp is the official co-supervisor with

Galloway of this project and student, assisted with project design and provided research direction.

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404 (Part B), pp. 44-56.

CHAPTER 2: PALYNOSTRATIGRAPHY OF THE JURASSIC-CRETACEOUS

TRANSITION IN THE NORTHERN RICHARDSON MOUNTAINS OF THE

NORTHWEST TERRITORIES, ARCTIC CANADA

ABSTRACT

Definition and recognition of the Jurassic-Cretaceous transition is unresolved due to various problems in terms of correlation that make the boundary difficult to define. Correlation of the boundary between the Tethyan and Boreal realms is problematic due to provincialism of age diagnostic taxa. Quantitative palynostratigraphy of the Husky Formation in the Aklavik Range,

Arctic Canada, thought to preserve the Jurassic-Cretaceous transition, is herein studied to provide insight into paleoenvironmental conditions in the Boreal Realm during this time. Miospores such as Concavissimisporites and dinoflagellate cysts such as Cribroperidinium cf. jubaris are considered for their potential importance in biostratigraphic correlations. Hierarchical cluster analysis is used to delineate stratigraphically unique terrestrial palynoassemblages whose compositions reflect moisture and temperature as interpreted using detrended correspondence analysis. Stratigraphically constrained cluster analysis reveals four groups containing palynomorphs with relatively distinct humidity and temperature preferences within upper

Tithonian to lower Berriasian strata of the Husky Formation. Two major palynological events are interpreted for the Canadian Arctic region: 1) an increase in humidity expressed by predominance of Cupressaceae-Taxaceae pollen, which is thereafter succeeded by 2) an increase in aridity as represented by increased abundances of Classopollis pollen. Paleoenvironmental reconstruction of the upper Tithonian – lower Berriasian Husky Formation is interpreted to have been characterized by predominantly humid conditions with a phase of a seasonal aridity.

INTRODUCTION

Detailed chronostratigraphy of Jurassic-Cretaceous strata in the Richardson Mountains,

Northwest Territories, is limited by the unknown position of the Jurassic-Cretaceous (J/K) boundary in Canada and globally. The J/K boundary remains the last Phanerozoic systems boundary to be defined by a Global Boundary Stratotype Section and Point (GSSP) in the Geologic

Time Scale. Its definition remains problematic due to the development of pronounced biogeographical provincialism of stratigraphically important fossils between the northern circumpolar Boreal Realm and the more southern Tethyan Realm following the breakup of the

Supercontinent Pangea during the Jurassic (Remane, 1991). Early Cretaceous climate change was closely related to the breakup of Pangea and predominantly characterized by greenhouse conditions and a humid climate, caused by increased moisture transfer from the oceans towards the continents (Főllmi, 2012). Climate change across the J/K boundary in the Boreal Realm is not widely understood, with much of the research on paleoclimate and boundary definition during this interval focused on areas from the Tethyan Realm due to its larger geographic area and greater accessibility (Wimbledon et al., 2011; Főllmi, 2012).

The J/K boundary in the Tethyan Realm is traditionally defined by the base of the ammonite Subzone Berriasella jacobi Mazenot, 1939 or the Pseudosubplanites grandis Mazenot,

1939 zone (Zakharov et al., 1996; Michalík and Reháková, 2011; Wimbledon et al., 2011; Schnabl et al., 2015). However, recent research suggests that most specimens labelled as B. jacobi in past literature may have been misidentified, making its status as an index fossil questionable for defining the base of the Berriasian Stage as its age range is poorly understood (Frau et al., 2016;

Schneider et al., 2018a). The Jacobi Subzone therefore cannot be used as a GSSP marker without more analysis (Wimbledon, 2017). In 2011, the Berriasian Working Group of the International

Subcommission on Cretaceous Stratigraphy proposed these primary markers for the J/K boundary in the Tethyan Realm: 1) the base of the Calpionella Lorenz, 1902 Zone – alpina Lorenz, 1902

Subzone; 2) the sudden bloom of Calpionella alpina; 3) the first appearance datums (FAD) of

Nannoconus steinmannii minor Deres and Achéritéguy, 1980 and N. kamptneri minor Bralower,

1989; and 4) the base of magnetozone M18r (Wimbledon et al., 2011). While correlations using ammonite zonation can be constrained by magnetostratigraphy where successions are complete and unaffected by diagenesis, the base of M18r does not correlate with any known biostratigraphic markers (Schnabl et al., 2015 and references therein). Magnetozone M19n.2n has instead been considered as a viable marker for the J/K boundary (Wimbledon et al., 2013; Schnabl et al., 2015).

In calpionellid research, the J/K boundary is drawn between the Crassicollaria Remane,

1962 and Calpionella zones (C. alpina Subzone), an increasingly recognized alternative to ammonite zonation in recent years due to the presence of a widespread and consistent turnover from Crassicollaria and large Calpionella to the small, globular Calpionella alpina that occurs within magnetozone M19n.2n (Michalík and Reháková, 2011; Wimbledon et al., 2011; Schnabl et al., 2015 and references therein; Wimbledon, 2017). The base of the Alpina Subzone is currently considered as the primary marker for the J/K boundary by the International Union of Geological

Science (IUGS) Berriasian Working Group. However, the boundaries of calpionellid zones are not well understood (Remane, 1991). The explosive bloom of Calpionella alpina has been suggested as an alternative indicator for the J/K boundary in the Tethyan Realm as it is both distinctive and widespread within the Tethyan Realm (Michalík and Reháková, 2011 and references therein).

Other markers include microfossils such as calcareous nannoplankton, which rapidly diversified during the J/K transition. Nannoconus steinmannii minor and N. kamptneri minor both have a FAD in M19n.2n (Wimbledon, 2017). The FAD of Hexalithus strictus Bergen, 1994 has been

considered as an alternate marker due its comparatively smaller stratigraphic distribution restricted to magnetozone M19n.2n (Wimbledon, 2017).

Despite these advances in research, one of the biggest challenges remains correlation of the J/K boundary between the Tethyan and Boreal realms. While correlation of Upper Jurassic and

Lower Cretaceous units have been extensively and successfully documented between northwest

Europe and the Russian Platform, correlation of the uppermost Jurassic to lowermost Cretaceous successions proves problematic due to the isolation of the Russian Platform during this time interval (Harding et al., 2011). Through the use of magnetostratigraphy, the base of the Berriasian in the Tethyan Realm correlates broadly with the Craspedites nodiger Eichwald, 1868 Zone

(Volgian) or with the Craspedites taimyrensis Shul’gina, 1962 Zone (upper Volgian) in the

Nordvik Peninsula of the Russian Platform (Houša et al., 2007; Harding et al., 2011; Bragin et al.,

2013). Biostratigraphic correlation using ammonite zonation is limited due to the endemism of ammonite species during the Mesozoic, and biostratigraphic correlation between the Tethyan and

Boreal realms remains difficult on a regional scale (Frebold, 1961). Many of the promising markers in the Tethyan Realm are absent or rare in the Russian Platform, an important area for

Boreal and Tethyan correlations (Harding et al., 2011). This provincialism likely resulted from the breakup of Pangea that allowed little to no faunal intermingling due to geographical barriers through intermittent marine connections and generally isolated seas (Remane, 1991). Due to all these complications, there exists no ammonite species or any alternative biomarkers that can singularly define the J/K boundary globally (Schnabl et al., 2015). In Northern Siberia, lower

Berriasian strata are defined by the Chetaites sibiricus Shul’gina, 1962 and Hectoroceras kochi

Spath, 1947 zones (Jeletzky, 1973), with the zonal boundary between Chetaites sibiricus and C. chetae Shul’gina, 1962 being representative of the Volgian/Ryazanian boundary, which was found

to correspond with the magnetozone M17r (Zakharov and Rogov, 2008; Bragin et al., 2013). In

Canada, definition of the Berriasian Stage is dependent on the Buchia okensis Pavlow, 1907 and

Craspedites (Subcraspedites) aff. suprasubditus Bogoslovsky, 1972 zones (Jeletzky, 1973).

Palynology has been used to supplement biostratigraphic correlation of the J/K boundary in the Boreal Realm and is particularly useful for linking non-marine facies where marine fossils are otherwise absent (Norris, 1970; Hunt, 2004). In comparison to ammonites, palynomorphs are highly resistant microfossils (e.g., pollen grains, spores, dinocysts, etc.) that are exceptionally well- preserved and abundant in strata that lack other fossils. Palynomorphs can be used to reconstruct the paleoenvironment where they were deposited based on the presumed affinity to parent plants and their inferred ecological preferences (Fægri and Iversen, 1989). Dinocyst species, in particular, can be used for more precise age determination in strata of Late Triassic age or younger because of their rapid evolution and wide geographic distribution (Evitt, 1985). For example, while the

Jacobi Subzone of the Tethyan Realm was previously correlated with the base of the

Subcraspedites primitivus Swinnerton, 1935 Zone of the Boreal Realm, the Jacobi Subzone is now considered to correspond to part of the Subcraspedites preplicomphalus Swinnerton, 1935 Zone based on dinocyst and miospore biostratigraphy from Dorset and the Spilsby Province (Hunt,

2004). Likewise, the first occurrence datum of the palynomorphs Warrenia californica Monteil,

1992, Dichadogonyaulax bensonii Monteil, 1992 and Apiculatisporis verbitskayae Dorhofer, 1976 have been found in both the Subcraspedites preplicomphalus Zone of the North Sea Basin as well as in the Purbeck sequence of the Hard Cockle Beds at Durlston Bay, both of which occur in the middle of magnetozone M19n (Hunt, 2004; Wimbledon et al., 2011). The first occurrences of the less widespread dinocysts Dapsilidinium warrenii (Habib, 1976) Lentin and Williams, 1981 and

Cantulodinium arthuriae van Helden, 1986 also appear at the top of M19n in Durlston

(Wimbledon et al., 2011). The calcareous dinoflagellate cyst species, Stomiosphaerina proxima

Řehánek, 1987, has also been considered as a potential secondary marker for the J/K boundary as its FAD is concurrent with the FAD of Calpionella alpina in Siberia (Reháková, 2000;

Vishnevskaya, 2017). Above the J/K boundary, the terrestrial palynomorphs Matonisporites elegans Hunt, 1985 and Aequitriradites spinulosus Cookson and Dettman, 1961 were found to have a first occurrence datum at magnetozone M17r in the Durlston Cherty Freshwater beds and in the Subcraspedites lamplughi Spath, 1924 Zone in the North Sea Basin (Abbink et al., 2001b;

Wimbledon et al., 2011).

In this study, quantitative palynology is used to supplement the paucity of age diagnostic macrofossils in the Richardson Mountains and to improve regional correlation in the Canadian

Arctic and elsewhere. Multivariate statistical analyses are used to determine if the study area preserves the local J/K boundary as defined by the base of Buchia okensis Zone, and to evaluate paleoenvironmental conditions that may have been prevalent during the earliest Berriasian.

Palynological investigations undertaken in the Canadian Arctic will be compared along with other

Boreal areas such as northwestern Europe and parts of Russia and Siberia. In doing so, a palynological signature can be identified to help correlate the J/K boundary across larger regions, define the J/K boundary in the Boreal Realm, and reconstruct the paleoclimate of the late Tithonian to earliest Berriasian in the Canadian Arctic.

GEOLOGIC SETTING AND STUDY AREA

In Canada, Upper Jurassic to Lower Cretaceous strata with faunas of Boreal affinities can be found in the Sverdrup Basin, Porcupine Plateau-Richardson Mountain Trough, and parts of the

Arctic Coastal Plain (Jeletzky, 1973; Dixon, 1992; Fig. 1). During the Jurassic-Early Cretaceous, the Boreal Realm in Canada was restricted to the Canadian Arctic Archipelago and the Beaufort-

Figure 1: Northern circumpolar maps showing a) plate reconstruction at 145 Ma, and b) estimated southern limits of the Boreal Realm during the J/K transition (modified from Jeletzky, 1971, 1973;

Saks, 1975; Zakharov, 2012). Red square denotes the region of interest and, within it, the estimated local extent of Boreal seas in the Late Jurassic-Early Cretaceous based on the distribution of ammonite and Buchia species (Jeletzky, 1971, 1973). Maximum extent of Berriasian Boreal seas may have included all purple areas, which is consistent with Zakharov’s (2012) interpretation.

Plate reconstruction and polar base maps created from ODSN website (http://www.odsn.de).

Mackenzie region of Yukon and the Northwest Territories (Jeletzky, 1973; Fig. 1). The Western

Cordillera of Canada and the Peace River region were constituents of the Tethyan Realm of Canada

(Jeletzky, 1973). The study area is located in the northern Richardson Mountains, west of the

Mackenzie Delta in the Northwest Territories of Canada (68°09’36”N, 135°40’12”W; Figs. 1, 2), where Lower Cretaceous rocks are widely exposed. Changes in tectonic regimes during the

Cretaceous are recorded through shifts in provenance, from an east-southeastern to south- southwestern provenance during the Aptian/Albian (Dixon, 1992). A regional unconformity separates the Lower from Upper Cretaceous rocks (Dixon, 1992).

Beaufort-Mackenzie Region

During the J/K transition, the North American craton was bordered by mostly marginal seas and marine deposition was primarily restricted to the northwestern shelf where the Brooks-

Mackenzie Basin was located (Balkwill et al., 1983; Fig. 2). The Brooks-Mackenzie Basin was bordered to the east and southeast by the North American craton, to the north by the Canada Basin, and to the south and southwest by the Western Cordillera (Fensome, 1987). The basin contains a succession of Mississippian to Hauterivian rocks, with the Mesozoic fill characterized by a predominance of clastic rocks derived from upland and cratonic sources (Balkwill et al., 1983;

Fig. 3). Strata of Berriasian to Hauterivian age were mainly deposited in shallow marine or deltaic settings (Balkwill et al., 1983); the Beaufort-Mackenzie region is therefore one of the few localities in North America that contain a continuous succession of marine rocks across the approximate J/K boundary (Fensome, 1983). From the Valanginian to Hauterivian, a major regressive episode caused progradation of sandy deltas northwestward away from basin margins (Balkwill et al.,

1983). A major sub-Hauterivian unconformity formed throughout the Mackenzie Delta area, marking the onset of rifting of the Canada Basin as continental breakup and subsequent sea floor

Figure 2: a) General map of the Beaufort-Mackenzie region (i.e., Brooks-Mackenzie Basin) in the northwesternmost area of the Northwest Territories, with area of interest boxed in red (modified from McNeil et al., 2001; Fensome, 1987; Dixon and Dietrich, 1990). b) geologic map of the

Martin Creek study area with outcrop location (68°09’36”N, 135°40’12”W) represented by the red dot (modified from Lane, 2005).

Figure 3: Modified lithostratigraphy of the northeastern Richardson Mountains (modified from

Dixon, 2004). The rock unit of interest is the mudstone-dominated Husky Formation, separated into four members. A Jurassic-Cretaceous boundary is thought to be preserved at, or near, the contact between the arenaceous and red-weathering members.

41 spreading progressed (Dixon, 1982). This unconformity also occurs as the break-up unconformity in the Sverdrup Basin (Galloway et al., 2013; Hadlari et al., 2016). In some areas of the Richardson

Mountains, normal faulting occurred simultaneously with deltaic deposition, the largest fault block of which is the Cache Creek Uplift that occurred in the mid-Hauterivian (Balkwill et al., 1983;

Fig. 2). This period of rifting resulted in the formation of the Arctic continental margin in post-

Albian time, which was followed by regional folding and faulting during the Late Cretaceous and

Paleogene (Lane, 2002). Compressional tectonic activity began to encroach on northern Yukon during the late Aptian to Albian, allowing for the Canada Basin margins to drift (Dixon, 1992).

The main phase of drifting occurred during the Late Cretaceous (Cenomonian to late

Maastrichtian) along with compression and some strike-slip motion that extended until the late

Miocene (Dixon, 1992).

Husky Formation

The Husky Formation is a mudstone-dominated rock unit of Late Jurassic to Early

Cretaceous age deposited along the southern margin of the Brooks-Mackenzie Basin. This succession contains five major transgressive-regressive (T-R) cycles (Poulton et al., 1992;

Georgescu and Braun, 2014; Figs. 3, 4). The Husky Formation is interpreted to have been deposited in a generally calm water, marine setting based on the prevalence of mudstone and marine fossils (Dixon, 1982). Extensive bioturbation has been recorded in the Husky Formation, representing well-oxygenated bottom conditions (Dixon, 1982). In the Mackenzie Delta-

Tuktoyaktuk Peninsula area, the formation has a thickness of up to 640 metres in outcrops, and

750 metres in the subsurface (Dixon, 2004). The Husky Formation overlies the sandstone- dominant Aklavik Formation of the Jurassic Bug Creek Group and underlies the sandstone-

Figure 4: Photograph of the study area, Martin Creek (68°09’36”N, 135°40’12”W). The red dotted line represents a fault. Yellow lines show two measured sections from which samples were collected. 1 - Lower member of the Husky Formation; 2 - Arenaceous member of the Husky

Formation; 3 - Red-weathering member of the Husky Formation; 4 - Upper member of the Husky

Formation; 5 - Martin Creek Formation.

dominant Martin Creek Formation (Figs. 3, 4). The base of the Husky Formation is diachronous, ranging from late Callovian to early Kimmeridgian in age (Jeletzky, 1967; Pocock, 1976). The

Husky Formation can be differentiated from the Bug Creek Group by the presence of Buchia

Rouillier, 1845 fauna. The formation was previously known as the “lower shale-siltstone division”

(Jeletzky, 1967). In outcrop, the unit is now typically divided into four informal members: in descending order they are the “upper”, “arenaceous”, “red-weathering” and “lower” members. In the Aklavik Range, the Husky Formation is predominantly composed of dark-coloured mudstone and siltstone, with higher sand content in the “lower” and “arenaceous” members. The “lower member” is the thickest of the four divisions, ranging from 227 to 245 metres thick in the Aklavik

Range (Hedinger, 1993). The contact between the underlying Aklavik Formation is sharp and likely disconformable, representing a period of rapid marine transgression (Hedinger, 1993). The

“arenaceous member” is composed of coarser material in comparison to the adjacent members

(Hedinger, 1993). The “arenaceous” and “lower” members represent transgressive beds (Dixon,

1982), and are considered Oxfordian to Tithonian in age based on dinocyst and bivalve assemblages (Brideaux et al., 1975, 1976). A major transgression can be observed in the uppermost portion of the “arenaceous member”, which may represent the J/K boundary (Dixon, 1992). The

J/K boundary is thought to exist at, or near, the abrupt contact between the “red-weathering” and

“arenaceous” members based on the presence of the bivalve Buchia okensis, which is diagnostic of the base of the Berriasian Stage in the area (Jeletzky, 1958; Brideaux, 1976). The “red- weathering member” is composed of recessive weathering shale with abundant layers of ironstone concretions while the “upper member” has a characteristic coarsening upward cycle (Hedinger,

1993). The “red-weathering” and “upper” members are primarily composed of mudstone deposited in an offshore setting that represent the regressive phase of a T-R cycle in the Berriasian, and are

part of a large coarsening upward cycle that goes into the overlying Martin Creek Formation

(Dixon, 1982). The “upper member” is therefore gradationally overlain by the nearshore to shoreline sandstones of the Martin Creek Formation, representing a facies boundary (Dixon, 1982;

Dixon, 2004). The “upper member” coincides with the upper Berriasian Buchia n. sp. aff. volgensis

Jeletzky non Lahusen, 1888 Zone (Jeletzky, 1958, 1960; Brideaux, 1976).

Thermal Maturity

The Husky Formation is a source rock for the Beaufort-Mackenzie Basin (Tang and

Lerche, 1991). The organic matter in the formation is dominated by Type III (gas prone) kerogen

(Dixon et al., 1985; Link et al., 1989; Tang and Lerche, 1991). The petroleum source potential in the upper and lower 100 metres of the formation has been referred to as “fair” by Link et al. (1989).

The lowermost beds of the Husky Formation have a total organic carbon (TOC) value of 4.5% that decreases upward (Link et al., 1989). The middle portion of the formation has a moderate TOC of

0.9% but a “poor” petroleum source potential (Link et al., 1989). The middle to upper part of the formation has recorded an increase in TOC values to 2.7% (Link et al., 1989). TOC values then decrease towards the overyling Martin Creek Formation (Link et al., 1989). The formation is therefore considered marginally mature, with a vitrinite reflectance value of 0.53% (Link et al.,

1989; Link and Bustin, 1989). Geochemical data suggest that the Husky Formation was mature at the end of the Early Cretaceous and during the Paleogene (Langhus, 1980). Peak hydrocarbon generation likely occurred during the Late Eocene to Oligocene (Tang and Lerche, 1991), wherein the Husky Formation produced significant liquid hydrocarbons with an approximate vitrinite reflectance value of 0.70% (Langhus, 1980).

Ammonite and Buchia Biostratigraphy in Arctic Canada

Macrofossils from the Canadian part of the North American Boreal Province are characterized by a relatively low diversity and are often rare or absent in a studied area (Jeletzky,

1973). In areas where ammonites are sparse, bivalves such as Buchia are commonly used for biostratigraphic correlation in the Canadian Arctic (Fig. 5) because of their abundance, facies tolerance and diversity. Jeletzky (1958, 1973, 1984) attempted to refine the J/K boundary in the

Canadian Arctic by creating a zonal scheme using Buchia, and the Husky Formation is dated primarily using this bivalve genus due to the absence age-diagnostic ammonites. The upper

Tithonian succession in the Canadian Arctic areas is defined, in part, by the Arenoturrispirillina waltoni Chamney, 1971 Assemblage Zone, which is characterized by the occurrence of

Craspedites (Taimyroceras?) canadensis Jeletzky, 1966 and Buchia terebratuloides Lahusen,

1888 sensu lato zones (Chamney, 1973). The Berriasian is defined by the occurrence of Buchia okensis and B. sp. nov. aff volgensis zones, which are represented by the Lituotuba gallupi

Chamney, 1971 assemblage and the Gaudryina milleri Tappan, 1955 assemblage Zone (Chamney,

1973; Brideaux, 1976). The base of the Buchia okensis Zone is the preferred marker for the J/K boundary in many parts of the Boreal Realm, including Arctic Canada, as it closely corresponds to the base of the Hectoroceras kochi d'Orbigny, 1849 ammonite Zone (Zakharov, 1987). The base of the H. kochi Zone is located near the boundary between the Volgian and Boreal Berriasian stages (Zakharov, 1987).

Field notes written by GSC geologist E.W. Mountjoy in 1962 record a measured section

(107D2) from the Jurassic Bug Creek Group up to the Lower Cretaceous Martin Creek Formation in the Aklavik Range. These field notes are stored at the Geological Survey of Canada, Calgary

Office. A preliminary fossil report by Jeletzky for this section can be found in Report no. KM-3-

Figure 5: Uppermost Jurassic - Lower Cretaceous biostratigraphic correlation chart comparing standard ammonite zonation schemes with pelecypod assemblages and dinocyst asssemblages in

Arctic Canada.

1963 under the Field no. 151-154MJ. The recorded presence of Buchia fossils in the Husky

Formation was used to distinguish those beds from the Bug Creek Group where Buchia fossils do not occur. In the Husky Formation, the “lower member” contains Buchia piochii Gabb, 1864 sensu lato and Buchia mosquensis von Buch, 1844 sensu lato and its allies, which are early Portlandian-

Kimmeridgian in age. The “arenaceous member” contains Buchia fischeriana d’Orbigny, 1845 sensu lato. The “red-weathering member” corresponds to the Buchia okensis Zone in that it contains Tollia (Subcraspedites) aff. suprasubditus Bogoslovsky, 1902 and T. (S.) aff. hoeli

Frebold, 1930. The “upper member” corresponds to the Buchia cf. uncitoides Pavlow, 1907 sensu lato and Buchia volgensis Lahusen, 1888 zones. The “red-weathering” and “upper” members were dated by Jeletzky (1967) as early Berriasian based on the known ranges of Buchia okensis, B. uncitoides and B. volgensis zones, although defining the basal Cretaceous is dependent specifically on the appearance of abundant and typical representatives of Buchia okensis sensu lato (Dixon

1982; Zakharov, 1987).

Previous Palynological Studies on the Jurassic-Cretaceous Boundary in Arctic

Canada

Published palynological information from the Canadian Arctic remains sparse as few

Arctic localities are sampled in comparison to European localities. The following are relevant works on J/K boundary strata in the Canadian Arctic: In northern Canada, Jeletzky (1958, 1960) reported sections that yielded strata that include the J/K boundary in a conformable sequence

(Pocock, 1967). The palynological assemblages in these sections are unfortunately in nearshore facies and contain relatively few microplankton (Pocock, 1976). Species in the Kimmeridgian assemblage in these sections include the dinocysts Endoscrinium luridum (Deflandre, 1939)

Gocht, 1970, Gonyaulacysta jurassica (Deflandre, 1929) Norris and Sarjeant, 1965, and

Pareodinia ceratophora. The Volgian assemblage contain fewer marine palynomorphs but

Gonyaulacysta jurassica and Pareodinia ceratophora remain common (Pocock, 1967). In comparison to microplankton, terrestrial palynomorphs are abundant, notably those pollen and spores from conifers and ferns such as Protoconiferus Bolchovitina spp., Rogalskaisporites Danze-

Corsin and Laveine, 1963 spp., Classopollis classoides, Applanopsis Döring, 1961 spp., and

Cerebropollenites mesozoicus (Pocock, 1967). Pocock (1967) notes that the Ryazanian assemblage is similar to the Volgian assemblage with the exception of the appearance of two

Cretaceous taxa: Cicatricosisporites dorogensis Potonié and Gelletich, 1961 and Arcellites Ellis and Tschudy, 1964 spp. Pocock (1967) therefore suggests that the first appearance of these palynomorphs could be used as a possible indicator for the basal Cretaceous. The signature produced by the Valanginian palynological assemblage of Jeletzky’s (1958, 1960) sections resembles the palynological signature of the J/K boundary in many parts of Europe. Some notable palynological trends include: a decrease in the abundance of Classopollis classoides pollen and

Schizaceae spores; the disappearance of Platysaccus Biard, 1963 and Rogalskaisporites spores; the first appearances of the spores Appendicisporites Weyland and Greifeld, 1953, Pilosisporites verus Delcourt and Sprumont, 1955, P. trichopapillosus, Januasporites Pocock, 1962 and dinocysts, Gardodinium eisenackii Alberti, 1961 and Cyclonephelium compactum Deflandre and

Cookson, 1955 at the base of the Valanginian (Pocock 1967).

Brideaux and Fisher (1976) were the first to publish information on Late Jurassic dinocyst assemblages in Arctic Canada. They described Oxfordian – Berriasian dinocyst assemblages in the

Beaufort-Mackenzie region and in the Canadian Arctic Archipelago, interpreting the

Paragonyaulacysta? borealis assemblage as a possible palynological marker for the J/K boundary in the area. Other dinocyst species included in this assemblage are: Paragonyaulacysta capillosa

(Brideaux and Fisher, 1976) Stover and Evitt, 1978, Epiplosphaera saturnalis, and Tetrachacysta spinosigibberosa (Brideaux and Fisher, 1976) Backhouse, 1988, which are also known from Upper

Jurassic to Lower Cretaceous successions in Arctic Canada (Brideaux and Fisher, 1976).

Brideaux (1976) extrapolated further on the Brideaux and Fisher (1976) publication and described Berriasian dinocysts collected from the “red-weathering” and “upper” members of the

Husky Formation, and the basal beds of the overlying Martin Creek formation in the Martin Creek area. Tetrachacysta spinosigibberosa appeared in the basal beds of the “red-weathering member”.

The confirmed age range of T. spinosigibberosa is therefore late Oxfordian to early Berriasian.

Occurrences of T. spinosigibberosa in younger strata were interpreted as having been derived from re-working from older rocks (Brideaux, 1976). Seven other taxa are known to commonly occur in other basal Cretaceous strata in the Canadian Arctic: Schizosporites reticulatus Cookson and

Dettmann, 1959 spores, and the dinocysts Pareodinia ceratophora, Sirmiodinium grossii,

Tubotuberella rhombiformis, Caligodinium aceras (Manum and Cookson, 1964) Lentin and

Williams, 1973, Circulodinium distinctum (Deflandre and Cookson, 1955) Jansonius, 1986 and

Scriniodinium campanula Gocht, 1959. Brideaux (1976) notes that the dinocyst assemblage from

Arctic Canada is distinct from other areas, but likely has Boreal affinities and does not share any affinities with Tethyan assemblages.

Pocock (1976) presented a preliminary Upper Jurassic – Lower Cretaceous dinocyst zonation scheme in the Canadian Arctic and compared these zonal indices with key Canadian ammonite and pelecypod faunas and those from the European zonations (Fig. 5).

Brideaux (1977) published a paper on the taxonomy of Early Cretaceous dinocysts and acritarchs from two subsurface sections in the Richardson Mountains in the northern Canadian mainland, but did not correlate these to any zonation schemes.

Davies (1983) studied Toarcian to Valanginian dinocysts in the Sverdrup Basin and identified seventeen Oppel-zones based on these palynomorphs (Fig. 5). These dinocyst assemblages exhibit relatively low diversity during the Tithonian to Berriasian, with most taxa having long ranges. Oppel-zones that range from the Tithonian to Berriasian include Oppel-Zone

J: Epiplosphaera saturnalis – Escharisphaeridia pocockii (Sarjeant, 1968) Erkmen and Sarjeant,

1980 (middle Kimmeridgian – upper Tithonian); Oppel-Zone K: Cribroperidinium jubaris (lower

– upper Tithonian); Oppel-Zone L: Prolixosphaeridiopsis spissum (McIntyre and Brideaux, 1980)

Hogg and Bailey, 1997 – Paragonyaulacysta capillosa (uppermost Tithonian – lower Berriasian);

Oppel-Zone M: Trichodinium erinaceoides – Tetrachacysta spinosigibberosa (uppermost

Tithonian – lower Berriasian); Oppel-Zone N : Cyclonephelium cuculliformis Davies, 1983 –

Paragonyaulacysta? borealis (Berriasian).

Bujak and Scott (1984) Research Ltd. provided an atlas of palynology and micropaleontology of the Cretaceous to Pleistocene strata based on both published and unpublished data from offshore and onshore wells, and surface sections exposed in the Beaufort-

Mackenzie area (Fig 5). They outlined age-diagnostic dinocysts found in the area and proposed a corresponding palynological zonation. According to the palynological zonation scheme proposed by Bujak and Scott (1984), the Berriasian Stage is entirely defined by the Tetrachacysta spinosigibberosa Zone based on published data from Brideaux, 1976, Brideaux and Fisher, 1976, and McIntyre and Brideaux, 1980. Notable taxa also present in this Zone include:

Paragonyaulacysta capillosa, which has its last occurrence in this Zone; Circulodinium distinctum, Scriniodinium campanula, Oligosphaeridium vasiformum (Neale and Sarjeant, 1962)

Davey and Williams, 1966, and Sentusidinium separatum (McIntyre and Brideaux, 1980) Lentin and Williams, 1981 all have a first occurrence in this Zone.

Fensome (1983) completed a PhD dissertation on the J/K boundary in Arctic Canada that included a comprehensive taxonomic review of miospores of strata preserved in the Aklavik

Range. This work was further expanded upon in a paper on the prominence of schizaealean spores that occur for the first time in J/K boundary strata in the Aklavik Range (i.e., Husky, Martin Creek,

McGuire, Fault Creek, Lower Canyon formations), and included a detailed taxonomic review of those species (Fensome, 1987). Biostratigraphically important spore taxa for defining the J/K boundary included Cicatricosisporites Potonié and Gelletich, 1933, Concavissimisporites

Delcourt and Sprumont, 1955, and Pilosisporites Delcourt and Sprumont, 1955. The proposed zonation scheme based on these palynomorphs are: the Concavissimisporites montuosus Döring,

1965 Zone (upper Oxfordian – Kimmeridgian); Cicatricosisporites abacus Burger, 1966 Zone

(Kimmeridgian – lowermost Ryazanian); Cicatricosisporites purbeckensis Norris, 1969 Zone

(lower Ryazanian to Hauterivian); Pilosisporites delicatulus Norris, 1969 Subzone (upper

Ryazanian – lower Valanginian).

Galloway et al. (2013) published a study on the Middle Jurassic – Lower Cretaceous palynostratigraphy of the Sverdrup Basin, and interpreted paleoenvironment and paleoclimate using multivariate statistical analyses. Aalenian to Albian palynoassemblages were analyzed.

Aalenian to Bathonian palynoassemblages were interpreted to have been derived from a coniferous forest with an understory of diverse fern, fern allies, and mosses, which suggests a humid and warm climate at the time. The Callovian to late Valanginian/early Hauterivian were thought to have been seasonally warmer and drier based partly on the relatively high abundance of

Classopollis pollen from Cheirolepidiaceae parent plants. Late Valanginian to Albian strata were characterized by an increase in Cupressaceae-Taxaceae pollen, which is indicative of cooling

54 conditions and an increase in moisture. These cooling conditions were inferred to have been a part of a general cooling trend experienced in the Canadian Arctic during the Early Cretaceous.

METHODS

Palynology

Forty-seven mudstone samples were collected for this project across the interpreted horizon of the J/K boundary in the Husky Formation preserved in the Aklavik Range of the Richardson

Mountains, Northwest Territories (Fig. 6). The section at Martin Creek was 94 metres thick. Near the contact between the “arenaceous” and “red-weathering” members, the succession was sampled for palynology every metre or less (Fig. 6). Above the contact, sample resolution was every two metres. Samples were prepared for palynological analyses by Leanne Tingley at the Geological

Survey of Canada and by Global GeoLab Ltd. using standard preparation procedures, including acid maceration, oxidative treatment with Schulze’s solution and staining with Safranin O. Liquid bioplastic was used to permanently mount the resulting slurries on glass microscope slides. Spores, pollen grains, and dinoflagellate cysts were enumerated and identified to the lowest taxonomic level with an Olympus BX61 microscope at 400x and 1000x magnification under oil immersion.

A minimum of 300 spores and pollen were counted in unsieved sample slides for statistical analyses (mean 344 ± 28 SD, n=47). A minimum count of 300 specimens has been shown to accurately reflect the constituents of a microfossil assemblage at a 0.95 confidence interval

(Phleger, 1960; Shaw, 1964; Dennison and Hay, 1967; Patterson and Fishbein, 1989; Buzas, 1990;

Revets, 2004). Dinocysts were enumerated alongside terrestrial material to achieve a ratio between terrestrial and marine palynomorphs. Photomicrographs were taken using Stream Motion software and an Olympus DP72 camera.

Figure 6: Lithostratigraphic column of the composite section measured, in metres, at Martin Creek with corresponding sample numbers. Coordinates for the base of the section are 68°09’36”N,

135°40’12”W.

Multivariate Statistical Techniques

Multivariate statistical analyses are based on compositional data of terrestrial palynomorphs preserved in the Husky Formation at the Martin Creek section. Despite uncertainties related to taphonomy that can affect the constituents of the palynoassemblage, the relative abundance of palynomorphs preserved in marine sediment has been shown to accurately reflect the distribution of onshore vegetation (Muller, 1959; Heusser and Balsam, 1977; Mudie, 1982;

Heusser, 1983; Hooghiemstra et al., 1986; Mudie and McCarthy, 1994; Sun et al., 1999; van der

Kars, 2001; van der Kaars and De Deckker, 2003; Hooghiemstra et al., 2006; Montade et al., 2011;

Luo et al., 2014; Zhao et al., 2016). Relative abundance of terrestrial palynomorphs in the Martin

Creek section are thus assumed to be capable of reconstructing the paleoenvironment from which they originated.

The computer program SYSTAT was used for hierarchical cluster analysis (Q- and R- mode), which group similar samples (Q-mode) or taxa (R-mode) in a dataset to provide information for paleoenvironmental interpretations and biostratigraphy. Using Ward’s linkage and

Euclidean distance, these hierarchical cluster analyses form dendrograms that organize sample pairs based on their similarities or differences in palynomorph content, joining similar pairs into a common node, with branch length indicating relative similarities (Paliy and Shankar, 2016). Q- mode cluster analysis groups samples that have similar palynological content while R-mode cluster analysis groups species that commonly occur together within a dataset. These cluster analyses use

Euclidean distance, which assumes a simplified, linear relationship to produce a quick visual of similarities between variables (Paliy and Shankar, 2016). These clusters can therefore be used to make broad interpretations about the inputted variables.

Stratigraphically constrained incremental sum-of-squares cluster analysis (CONISS) with square root data transformation to up-weight rare types was used to define palynological trends through time. Unlike unconstrained cluster analyses, CONISS is capable of defining stratigraphic zones by analyzing palynomorph content in samples in stratigraphically adjacent clusters (Grimm,

1987). Trends seen in clusters that are stratigraphically constrained can therefore be interpreted as trends seen through time, even in assemblages that “repeat”. This is particularly important for biostratigraphic purposes and, subsequently, defining the J/K boundary, as informal assemblages can be created based on the palynological signature in these clusters and compare and potentially correlate them across regions.

The DECORANA function from the R package VEGAN was used with the computer program RStudio for detrended correspondence analysis (DCA), a multivariate statistical technique used to interpret species and sample ordination (Hill and Gauch, 1980). DCA was chosen because biota and their environment typically have a non-monotonic and unimodal relationship, rather than a linear one as assumed by other, widely-used, multivariate statistical techniques (e.g., principal component analysis (PCA)) and can deal with zero data (Paliy and

Shankar, 2016). DCA has similar goal to PCA: selecting two axes or variables that can account for the largest proportion of variance to represent observations in a two-dimensional graph (Paliy and

Shankar, 2016). The eigenvalues associated with each axis depicts the extent to which each axis can influence the observations (Ramette, 2007). DCA is an algorithm created in FORTRAN by

Mark Hill in 1979 to compensate for the “arch” and “edge” effect inherent in other multivariate analyses such as PCA and reciprocal averaging (RA) or correspondence analysis (CA). The “arch” effect, otherwise known as the “horseshoe effect”, is the two-dimensional, mathematical presentation of ecological data as a curve rather than as a straight line due to the second axis being

both uncorrelated with, and dependent on, the first axis (Hill and Gauch, 1980). To make accurate interpretations about these axes, they need to be both uncorrelated and independent of each other

(Hill and Gauch, 1980). The “edge effect” is produced because RA does not preserve compositional distances, wherein two pairs of samples with the same compositional difference will exhibit two different distances in the ordination depending on where they are located on the first axis; compositional differences appear greater in the middle of the gradient than on the ends (Hill and Gauch, 1980). DCA is therefore, effectively, a modified version of RA: DCA eliminates the

“arch effect” by dividing the first axis into segments and then rescaling its axes such that the values on the second axis are equalized to a zero-mean value. This rescaling also ensures that equal compositional differences are represented by equal differences along the gradient, thereby eliminating the “edge effect” (Hill and Gauch, 1980). Through DCA, species distribution and sample variation can be observed. DCA is therefore an invaluable tool for presenting data without the distortion seen in other multivariate techniques and can be easily used in interpreting species and samples ordination for large datasets.

RESULTS

Spores and Pollen

Spores and pollen preserved in samples collected from the Martin Creek section were enumerated to a minimum of 300 counts (mean 344 ± 28 SD, n=47). Spore and pollen ranged from poorly to moderately-well preserved. Specimens were identified to their lowest possible taxonomic division, with 54 different taxa identified (Table 1).

Unknown “palynomorph A” (Plate 19)

Description: trilete(?), laesurae extended ¾ to equator, amb circular, spherical, cingulate, psilate sculpture, 7-10µm in length and width, hilum sometimes appears cracked, can appear as uniplanar tetrads.

Dinoflagellate Cysts

Dinocysts are poorly preserved in the measured section and are characterized by low abundances and low species diversification in contrast to pollen and spores. The average dinocyst count in the Martin Creek section is 5.02 ± 6.26 SD, n=47, for every minimum spore and pollen count of 300. The marine-to-terrestrial ratio varies from zero to a maximum of 13:173 (7.47% dinocyst content), which occurs at 28 m of the measured section. Relatively higher marine content tends to be found in the basal and upper beds of the section; from 0-30 m, dinocyst content ranges from >0-8%; beds located within 32-70 m have a dinocyst percentage consistently below 1.5%; beds located at 72 m and above have a dinocyst percentage ranging from 0-7% (Fig. 8).

Fourteen different dinocyst genera were identified, with eight of those identified to the species level (Table 2). The dinocyst content in the Martin Creek section was primarily composed of what was identified as Paragonyaulacysta? borealis, although the paratabulation in the observed specimens was unclear. Pareodinia ceratophora may be an alternate identification; P. ceratophora is a common and widespread species that is Aalenian (Powell, 1992) to late Callovian- early Oxfordian (Stover et al., 1996) in age. All the observed specimens in the study area tentatively identified as Paragonyaulacysta? borealis have a pareodinioid shape and a suggested intercalary archeopyle, but the number of plates involved are not clear. Identifying this species is made more difficult due to the incomplete description of the holotype of Pareodinia ceratophora, which lacks the archeopyle location, a critical characteristic for identifying dinocysts to the species level (R.A. Fensome, pers. comm., 2018). Other dinocysts identified to the species level were:

Table 1: Biological nomenclature and taxonomic authority of identified spores and pollen grains in the Jurassic-Cretaceous Husky Formation

Selaginellales Selaginellaceae Densoisporites spp. Dettmann 1963

Foveosporites spp. Balme 1957

Neoraistrickia truncata (Cookson) Potonié 1956 incertae sedis Lycospora spp. Schopf et al. 1944

Polypodiopsida Cyatheales Dicksoniaceae Cibotiumspora juncta (Kara-Murza) Zhang 1978 ?Dicksoniaceae, Concavissimisporites (Couper 1958) Döring Cyatheaceae apiverrucatus 1965 Concavissimisporites (Couper) Brenner variverrucatus 1963 Concavissimisporites (Delcourt and crassatus Sprumont 1955) Delcourt et al., 1963 Dicksoniaceae, Cyathidites minor Couper 1953 Cyatheaceae Cyathidites australis Couper 1953

Deltoidospora hallii Miner 1935

Gleicheniales Gleicheniaceae Gleicheniidites Ross 1949 senonicus Matoniaceae Dictyophyllidites harrisii Couper 1958

Matonisporites (Balme 1957) crassiangulatus Dettmann 1963 Osmundales Osmundaceae Baculatisporites (Cookson 1953) comaumensis Potonie 1956 Biretisporites potoniaei Delcourt and Sprumont 1955 Osmundacidites Couper 1953 wellmanii Rugulatisporites spp. Pflug and Thomson 1953 Todisporites (Maljavkina) Pocock rotundiformis 1970 Verrucosisporites spp. Potonié and Kremp 1954 Polypodiales Pteridaceae Contignisporites Dettmann 1963 cooksonii

Schizaeales Schizaeaceae Ischyosporites spp. Balme 1957 Klukisporites spp. Couper 1958

Ruffordiaspora cf. (Cookson 1953) australiensis Dettmann and Clifford 1992 incertae sedis Pilosisporites (Thiergart) Delcourt trichopapillosus and Sprumont 1955 Undulatisporites Brenner 1963 undulapolus Pinopsida Pinales Araucariaceae Araucariacites australis Cookson 1947

Cheirolepidiaceae Classopollis classoides (Pflug) Pocock and Jansonius 1961 Cupressaceae Perinopollenites (Couper) Dettmann elatoides 1963 Cupressaceae, Undifferentiated Taxaceae Pinaceae Cerebropollenites (Couper) Nilsson mesozoicus 1958 Laricoidites magnus (Potonié) Potonié, Thomson and Thiergart 1950 Undifferentiated bisaccate pollen

Cycadopsida, Cycadales, incertae sedis Cycadopites spp. Wodehouse 1933 Ginkgoopsida Ginkgoales Chasmatosporites spp. (Nilsson 1958) Pocock and Jansonius 1968 Cycadopsida Cycadales incertae sedis Entylissa spp. Naumova 1937

Gnetopsida incertae sedis incertae sedis Eucommiidites troedsonii Erdtman 1948 incertae sedis Acanthotriletes Pocock 1962 varispinosus Granulatisporites spp. Potonié and Kremp 1954 Triquitrites spp. (Wilson and Coe 1940) Potonié and Kremp 1954

Table 2: Biological nomenclature and taxonomic authority of identified dinoflagellate cysts

(Division Dinoflagellata) in the Jurassic-Cretaceous Husky Formation

Family Subfamily Genus/Species Authority Gonyaulacaceae Cribroperidinioideae Cribroperidinium cf. jubaris (Davies 1983) Lentin and Williams 1985 Gonyaulacoideae Psaligonyaulax spp. Sarjeant 1966 Spiniferites spp. Mantell 1850 Tubotuberella rhombiformis Vozzhennikova 1967 Leptodinioideae Cymososphaeridium spp. Davey 1982 Dichadogonyaulax spp. Sarjeant 1966 Gonyaulacysta dualis (Brideaux and Fisher 1976) Stover and Evitt 1978 Gonyaulacysta? pectinigera (Gocht 1970) Fensome 1979 Rhynchodiniopsis spp Deflandre 1935 Sirmiodinium grossii Alberti 1961 incertae sedis Hystrichodinium spp. Deflandre 1935 Trichodinium cf. erinaceoides Davies 1983 Pareodiniaceae Pareodinioideae Paragonyaulacysta? borealis (Brideaux and Fisher 1976) Stover and Evitt 1978 Pareodinia ceratophora Deflandre 1947 incertae sedis incertae sedis Epiplosphaera cf. saturnalis (Brideaux and Fisher 1976) Dodekova 1994

Cribroperidinium cf. jubaris, Epiplosphaera cf. saturnalis, Gonyaulacysta dualis,

Gonyaulacysta? pectinigera, Sirmiodinium grossii, Trichodinium cf. erinaceoides and

Tubotuberella rhombiformis. The dinocysts Cribroperidinium cf. jubaris, Epiplosphaera cf. saturnalis, Gonyaulacysta dualis, and Gonyaulacysta? pectinigera are generally found to have higher abundances lower in the section. The dinocysts Trichodinium cf. erinaceoides and

Tubotuburella rhombiformis are generally found to have higher abundances nearer to the top of the section. The dinocysts Paragonyaulacysta? borealis and Sirmiodinium grossii appear sporadically throughout the section.

Multivariate Statistical Analyses: Q- and R-mode Cluster Analysis, CONISS

Q-mode cluster analysis is used to delineate four clusters that correspond to the stratigraphic section (Fig. 7). Clusters 1 and 3 are located generally lower in the section whereas clusters 2 and 4 are generally located higher: cluster 1 contains samples taken from 32 – 40 m; cluster 2 contains samples from 28 – 82 m; cluster 3 contains all samples collected from the

“arenaceous member” and a few samples from the basal beds of the “red-weathering member” (0

– 11 m). The J/K boundary is thought to lie within Cluster 3. Cluster 4 contains samples from 44

– 94 m.

R-mode cluster analysis is used to delineate three clusters that are named after their respective dominant palynomorph group: spore assemblage, pollen assemblage, Cupressaceae-

Taxaceae (CT) pollen assemblage (Fig. 8). The spore assemblage contains a diverse assemblage of mainly Bryophyta (mosses), Polypodiopsida (ferns) and Lycopodiopsida (fern allies) spores and constitutes 9.07% of the total palynomorph count; the most abundant species, on average, are

Leptolepidites verrucatus (0.72% ± 0.86 SD, n=47) and Cingulatisporites spp. (0.72% ± 0.72 SD, n=47) spores, which each represent 8% of this assemblage. They are followed in abundance by the

Figure 7: Q-mode cluster analysis delineated to four different clusters with relatively distinct stratigraphic positions. The J/K boundary is thought to lie within Cluster 3.

Figure 8: CONISS combined with R-mode cluster analysis showing R-mode clusters (Spore,

Pollen, Cupressaceae-Taxaceae (CT) pollen assemblages), % dinocysts per minimum 300 counts of pollen and spores, relative abundance of identified terrestrial palynomorphs, and stratigraphically constrained clusters (A, B, B2, C, D). Red dotted line represents the inferred J/K boundary.

spores Gleicheniidites senonicus (0.68% ± 0.64 SD, n=47) and Todisporites rotundiformis (0.69%

± 0.76 SD, n=47), each representing approximately 7.5% of the spore assemblage.

The pollen assemblage is the largest group in terms of abundance, constituting 56.68% of the total palynomorph content. Palynomorphs are derived mostly from parent plants belonging to the Class Pinopsida (gymnosperms) and with subordinate spores from ferns and allies. The most abundant palynomorphs from the gymnosperms are undifferentiated bisaccate pollen grains

(15.88% ± 6.73 SD, n=47), representing 28% of the pollen assemblage, followed by Classopollis classoides pollen (4.82% ± 5.08 SD, n=47), representing 8.4%. The most abundant spores are

Osmundacidites wellmanii (8.33% ± 4.61 SD, n=47), representing 14.5% of the assemblage, followed by Cyathidites australis (5.50% ± 3.53 SD, n=47), representing 9.6%.

The CT pollen assemblage is composed solely of CT pollen, which constitute 34.25% of the total palynomorph count (33.63% ± 19.35 SD, n=47). All microflora assemblages in the Martin

Creek section are characterized by a prevalence and dominance of CT pollen.

The relative abundances of R-mode defined clusters were stratigraphically constrained using CONISS to observe palynological trends over time (Fig. 8). Clusters defined by CONISS can be divided into four groups: Clusters A-D. Cluster A is a mixed assemblage that is distinguished by a general decrease in abundance of CT pollen and an increase in Araucariacites australis pollen, which peaks at 88 m with a relative abundance of 13.86%, and Unknown

Palynomorph A, which spikes in abundance to 12.68% at 92 m. Cluster B1 is also a mixed assemblage, and can be distinguished by a peak in Laricoidites magnus pollen to 11.85% at 74 m.

Cluster B2 is characterized by relatively high abundances of Classopollis classoides pollen, particularly between 46 and 56 m. Classopollis classoides pollen reach a peak at 52 m with a relative abundance of 25.07%. Cluster C is characterized by a high abundance of CT pollen,

reaching a maximum of 95.97% at 40 m, alongside relatively high abundances of Perinopollenites elatoides pollen, which peaks at 5.46% at 28 m, and Araucariaceae australis pollen, which increases to 9.77% at 28 m. Clusters C and D are separated due to a covered section between 11 and 28 m. Cluster D is characterized by a relatively high abundance of Baculatisporites comaumensis spores, undifferentiated bisaccate pollen, Cerebropollenites mesozoicus pollen, and

Cyathidites australis spores, all of which generally decline in abundance thereafter.

Baculatisporites comaumensis spores reach peak abundance at 10 m, with a relative abundance of

20.39%; undifferentiated bisaccate pollen reaches its maximum in this cluster with a relative abundance of 31.07% at 2 m; Cerebropollenites mesozoicus pollen reach a peak at 0.3 m, with a maximum relative abundance of 6.56%; Cyathidites australis spores reach maximum abundance of 16.26% at 10.8 m. If the J/K boundary lies exactly at the contact between the “arenaceous” and

“red-weathering” members, the location of the J/K boundary must be located within Cluster D.

Detrended Correspondence Analysis (DCA)

Data from the DCA plot (Fig. 9) in the Martin Creek section are affected by changes represented by two axes that have eigenvalues of 0.2 (DCA1) and 0.07 (DCA2). Cupressaceae-

Taxaceae pollen plots near the lowest values on DCA1; Eucommiidites troedsonii pollen plots near the highest values. Classopollis classoides pollen plots near the lowest values on DCA2, whereas

Polypodiopsida, incertae sedis spores and Araucariacites australis pollen plot near the highest values on DCA2. DCA1 is interpreted as representing temperature whereas DCA2 is interpreted to represent moisture based on known ecological tolerances of the parent plants from which the aforementioned palynomorphs are interpreted to have been derived (Table 1).

Figure 9: DCA plot of Martin Creek section with hypothesized factors “moisture” for the axis

DCA2 and “temperature” for the axis DCA1 influencing species ordination. Respective eigenvalues are 0.07 and 0.2.

DISCUSSION

Biostratigraphy

Dinoflagellate Cysts

Dinocysts identified in the Martin Creek section are generally consistent with a Late

Jurassic-Early Cretaceous age (Fig. 10). Trends in the dinocysts described below lends evidence to the J/K boundary existing within the measured section of this study, although low abundance and diversity of marine elements in this section, combined with a lack of taxonomic changes in the assemblage makes this tenuous.

The dinocysts Cribroperidinium cf. jubaris, Epiplosphaera cf. saturnalis, Gonyaulacysta dualis, and Gonyaulacysta? pectinigera, which are restricted to the Upper Jurassic, are generally found to have higher abundances in the lower beds of the Martin Creek section. According to

Davies (1983), Cribroperidinium jubaris has a first occurrence in the early Tithonian and a last occurrence in the late Tithonian. Epiplosphaera saturnalis is thought to occur from late Oxfordian to late Kimmeridgian times (Brideaux and Fisher, 1976). Gonyaulacysta dualis has an age range encompassing all of the Oxfordian and up to the middle Kimmeridgian (Brideaux and Fisher,

1976). The presence of both Epiplosphaera cf. saturnalis and Gonyaulacysta dualis lend evidence to the Husky Formation having a diachronous base with an age ranging from the late Callovian to early Kimmeridgian, or at least evidence of some reworking. Gonyaulacysta? pectinigera was considered to have a first occurrence in early Oxfordian times (Powell, 1992), but may occur from as early as the Bajocian (Wiggan et al., 2017). Paragonyaulacysta? borealis is the longest ranging dinocyst in this study, with a first occurrence in the Toarcian and a last occurrence in the Barremian

(Brideaux, 1977). Sirmiodinium grossii is also relatively long ranging: its first occurrence may be as early as late Bathonian (Powell, 1992) or late Kimmeridigian (Brideaux and Fisher, 1976). The

Figure 10: First and last occurrences of identified dinocyst species observed in the Husky

Formation at the Martin Creek section. See text for age range references.

last occurrence for Sirmiodinium grossii may lie between the Hauterivian and Barremian

(Brideaux, 1977) or even up to early Aptian times (Stover et al., 1996). This is consistent with the trend observed in this study, wherein Paragonyaulacysta? borealis and Sirmiodinium grossii appear throughout the Martin Creek section.

Two dinocyst species identified in the study section are most indicative of an earliest

Berriasian age: Trichodinium cf. erinaceoides and Tubotuburella rhombiformis. Both of these species are typically found in higher abundances in the upper half of the Martin Creek section.

Trichodinium erinaceoides has a first occurrence in the late Tithonian and a last occurrence in the early Valanginian (Davies, 1983). Tubotuburella rhombiformis ranges in age from the Tithonian to the Berriasian (Lebedeva and Nikitenko, 1999). These age ranges are consistent with the age determination based on the presence of Buchia okensis from the Husky Formation in the northern

Richardson Mountains. It is important to note that the specimens of T. cf. erinaceoides observed in the Martin Creek section cannot be confidently identified as the position of the archeopyle is unclear, and may have an alternate identification with a species of Cometodinium.

The dinocyst assemblage in the Martin Creek section is compared to the Oppel-zones defined by Davies (1983; Fig. 5) in his study of Jurassic-Lower Cretaceous dinocysts in the

Sverdrup Basin to see if the J/K boundary can be defined by dinocysts throughout the Canadian

Arctic. Of those Oppel-zones, the Martin Creek dinocyst assemblage most closely resembles

Oppel-zones K-M. The rare species Cribroperidinium jubaris occurs in both the Martin Creek section and in Davies’ (1983) sections in the Sverdrup Basin at Elf Jameson Bay, Elf Wilkins,

Reineer Peninsula, Central Amund Ringnes Dome, northwest Cornwall Island, Jaegar River,

Skaare Anticline, Fosheim Anticline and Cape Ludwig. Cribroperidinium jubaris is used to define

Oppel-Zone K, with an age range of early – late Tithonian, with a possible early Berriasian age in

uppermost sections, although many Kimmeridgian and older specimens also appear in this Zone

(Davies, 1983). It is important to note that the identification of C. cf. jubaris in the Martin Creek section was tentative as the preservation of the few recorded specimens was poor.

Cribroperidinium cf. jubaris appears at the base of the Martin Creek section and in the sample collected at 28 m. It is nowhere else present in the preparations. Assuming the identification of C. cf. jubaris in the Martin Creek section is correct and that C. jubaris has a last occurrence in the latest Tithonian, the Jurassic portion of the Martin Creek section may include basal strata from the

“red-weathering member”, and therefore the J/K boundary is above 28 m. Thus, the J/K boundary may exist at least 22 m above the base of the “red-weathering member” rather than precisely at the contact between the two members. Field notes written by E.W. Mountjoy first records finding the

Berriasian indicator Buchia okensis 36 m above the base of the “red-weathering member”. Placing the J/K boundary at least 22 m above the base of the “red-weathering member” is therefore more consistent with this documented occurrence of B. okensis 36 m above the base of the “red- weathering” member in the Martin Creek area. However, this is in direct contrast to previous studies (Jeletzky, 1967) conducted in the Aklavik Range where B. okensis was found a few feet above the base of the “red-weathering member”. If the J/K boundary is located at or near the base of the “red-weathering member” as originally postulated, the specimens of C. cf. jubaris found in this study may simply be a result of recycling of Upper Jurassic material.

Oppel-zones L “Prolixosphaeridiopsis? spissum – Paragonyaulacysta capillosa” and M

“Trichodinium erinaceoides – Tetrachacysta spinosigiberrosa” have the same range from uppermost Tithonian – Berriasian. Oppel-Zone L is defined by the first occurrence of

Prolixosphaeridiopsis? spissum and the last occurrence of Epiplosphaera saturnalis and/or

Paragonyaulacysta capillosa. The base of this Oppel-Zone therefore corresponds to 5.4 m in the

Martin Creek section based on the last appearance of E. cf. saturnalis. The base of Oppel-Zone M

“Trichodinium erinacoides – Tetrachacysta spinosigiberrosa” is defined by the first occurrence of

Trichodinium erinacoides and Moesiodinium raileanui Antonesçu, 1974. The top of the Zone is defined by the last appearance of T. spinosigiberrosa and M. ehrenbergii. Of the biostratigraphically important dinocysts mentioned in Davies’ (1983) paper, only two could be identified in the Martin Creek section: Epiplosphaera cf. saturnalis and Trichodinium cf. erinaceoides. Epiplosphaera cf. saturnalis has only one recorded appearance in the preparation of the sample collected at 5.4 m. Trichodinium cf. erinacoides appears most abundantly between 0.3

– 2.1 metres, ranging between 2-4 counts (relative abundance of 16-20%), and then sporadically throughout the section thereafter. This is consistent with a latest Tithonian – Berriasian age, but otherwise this zonation scheme is not capable of distinguishing the J/K boundary in the Martin

Creek section.

In comparing the Boreal assemblages of dinocysts outside of North America, many of the dinocysts seen in the Martin Creek section share common, long-ranging species with those encountered in Siberia and Russian Platform sections, specifically Paragonyaulacysta? borealis/Pareodinia ceratophora, Sirmiodinium grossii, and Tubotuberella rhombiformis

(Pestchevitskaya et al., 2011 and references therein). Those species that have a shorter stratigraphic range, and are therefore biostratigraphically significant, are not shared between the

Martin Creek section and sections from Siberia and the Russian Platform.

Spores and Pollen

Most identified spores and pollen were consistent with a Jurassic – Cretaceous age, with the exception of Triquitrites spp. and Lycospora spp., which commonly appear in Carboniferous strata. The presence of these palynomorphs suggest some reworking occurred in the Husky

Formation. Reworking of Carboniferous spores is relatively common in Mesozoic strata (Windle,

1979 and references therein; Guy-Ohlson et al., 1987). Longer ranging palynomorphs such as

Densoisporites spp. and Murospora spp. are also common in Carboniferous sediments, although they appear throughout the Mesozoic as well, whether by reworking or deposition is unclear

(Windle, 1979).

Assemblage zones based on miospores are notoriously difficult to create due to their long- ranging affinities and a limited understanding of their evolutionary history. The absence or presence of a species can only be attributed to resulting evolutionary development when said species appears in the same stratigraphic order at different, geographically separated locations

(Fensome, 1987). The appearance of miospore species in different stratigraphic order suggests a local environmental control rather than a result of evolution (Fensome, 1987). While miospore zonations are most useful locally, they are still limited by the effects of preservation that may obscure the presence of biostratigraphically important taxa. The results of this study are compared to the miospore zonation scheme proposed by Fensome (1987) based on palynomorphs found in the outcrops exposed along Martin Creek north of the section in this study.

The spores Cicatricosisporites, Concavissimisporites, and Pilosisporites have been considered as potentially important biostratigraphic markers for J/K boundary strata in the Aklavik

Range (Fensome, 1983, 1987). The spore Cicatricosisporites abacus is commonly the first of its genus to appear in an area and was often times labelled under different names such as C. australiensis (Norris, 1969) or Plicatella abacus (Wimbledon and Hunt, 1983; Lindstrőm and

Erlstrőm, 2011). Cicatricosisporites abacus spores are thought to have a first occurrence in

Kimmeridgian, early Volgian, or early Tithonian times and persists in Volgian, Tithonian, and

Ryazanian strata (Fensome, 1987). After the first appearance of C. abacus spores,

Cicatrisosiporites spores increase in diversity, with C. australiensis (synonymous to

Ruffordiaspora cf. australiensis seen in this study), C. crassistriatus and C. purbeckensis spores appearing successively thereafter. C. purbeckensis spores first appear at the base of the “red- weathering member” of the Husky Formation, corresponding approximately to the

Preplicomphalus Zone and the boundary between the Volgian and Ryazanian stages. In this study, spores exhibiting cicatricose ornamentation were categorized under the name Ruffordiaspora cf. australiensis, with an alternate identification of Cicatricosisporites aralicus (Bolkhovitina, 1961)

Brenner, 1963. These specimens were mainly canaliculate to isocostate in that their ornamentation had canals that were of approximately equal width or smaller than the accompanying costae. The observed specimens were poorly preserved, often fragmented, and were not found in high abundances, first appearing at the base of the section and disappearing sporadically throughout the succession.

The diversification of schizaealean spores seen in other J/K boundary strata is not as apparent in the Martin Creek section and were not among the more abundant species observed in this study. Schizaealean spores are therefore limited in defining miospore zonations even in a local setting. However, considering the Martin Creek section in this study likely only encompasses latest

Tithonian and earliest Berriasian strata, it is possible that this palynological signature does not become apparent until Valanginian time (e.g., Vakhrameyev and Kofova, 1982).

Miospores from the genus Concavissimisporites in this study appear to follow a trend of increasing diversity as well as increasing relative abundance, comparable to studies done by

Fensome (1983, 1987) and Abbink et al. (2001a). In the Terschelling Basin, the Netherlands, an increase in Concavissimisporites spores was, in part, suggested as a marker within J/K boundary strata, occurring within the Heteroceras kochi Zone (Abbink et al., 2001a). The oldest strata in

which Concavissimisporites crassatus spores has been recorded was within the “lower” member of the Husky Formation (Kimmeridgian) and in the McGuire Formation (Fensome, 1987). The species has also been found in Lower Cretaceous strata in Belgium (Delcourt and Sprumont, 1955) and in Berriasian-Aptian strata in the U.S. (Bebout, 1981). In this study, C. crassatus spores appear primarily within the “red-weathering member” of the Husky Formation, although in very low abundances and only towards the top of the measured section, which is consistent with an early

Berriasian age for this strata. Concavissimisporites apiverrucatus spores appear throughout latest

Jurassic to Albian strata in the subsurface of the Mackenzie Delta-Richardson Mountains area and in the Canadian Arctic Islands (Brideaux and Fisher, 1976). Concavissimisporites apiverrucatus spores are rare in the “lower” and “arenaceous” members of the Husky Formation in the Martin

Creek area (Fensome, 1987). Following Fensome’s (1987) miospore zonation, C. apiverrucatus spores first appear at the base of the Cicatricosisporites abacus Zone. In this study, C. apiverrucatus spores appear only once at 3.2 m; this is part of the meterage of the uppermost beds of the “arenaceous member” of the Husky Formation, which is probably late Tithonian in age. The first and last occurrence of this species is therefore difficult to place as it is unknown whether its presence or absence is due to climatically- or environmentally-controlled factors.

Pilosisporites trichopapillosus spores are mostly known to occur above the J/K boundary, but has been recorded rarely in upper Tithonian and Kimmeridgian strata (Fensome, 1987). In the

Martin Creek section in this study, P. trichopapillosus spores are poorly preserved and appear rarely in the “red-weathering member”, a few times at 30 metres and once at 92 metres, and is entirely absent within the “arenaceous member”.

Based on these palynological comparisons, the J/K boundary is likely preserved in the

Martin Creek section in this study. However, while these results are broadly comparable to those

outlined in Fensome’s 1987 study, the low abundance and generally poor preservation of each biostratigraphically significant miospore in the Martin Creek section makes it difficult to observe palynological trends that may have been apparent in other J/K boundary sections. Even locally, palynological zonation schemes dependent solely on the first occurrences of miospore species proves difficult to use at high resolutions. In place of miospore zonation schemes, informal palynoassemblage zones are created via quantitative multivariate statistical analyses in this study, which can be used to make interpretations about the paleoenvironment and potential changes in the paleoclimate across the J/K boundary. These bioevents may provide more information in defining the J/K boundary than the occurrence of biostratigraphically important taxa, whose presence or absence may fluctuate depending on changes in the paleoenvironment rather than large-scale factors such as climate change.

Paleoecology

Paleoenvironmental interpretations of Jurassic – Early Cretaceous palynomorphs are dependent on understanding the ecology of parent plants from which quantitatively significant spores and pollen are produced (Patterson and Fishbein, 1989). Spores and pollen that are deposited in a given area reflect the composition of the region’s terrestrial plant community. This relationship can be used as a basis for a model to interpret changes in the paleoenvironment, particularly in regards to Cenozoic assemblages for which their parent plants are better known than older time intervals (Abbink et al., 2004). Paleoecological interpretations of major plant groups seen in this study are briefly discussed below.

Bryophytes and Lycophytes

Bryophytes can be divided into three groups: the mosses (Bryophyta), liverworts

(Marchantiophyta) and hornworts (Anthocerotophyta), the latter of which were not observed in

this study. Extant mosses typically grow in moist conditions, although they are also known to withstand droughts in drier environments (Abbink et al., 2004). Extant liverworts commonly thrive in moist, shaded areas, with some capable of living in aquatic habitats (Raven et al., 2013). Plants from the Class Lycopodiopsida (i.e., lycophytes, fern allies) have a diverse ecology, although they more commonly occur in tropical areas. Some species are capable of growing in temperate and polar regions with moist conditions (Galloway et al., 2015). Due to these wide ranges in habitats, most spores derived from lycophytes and bryophytes have a limited use in interpreting paleoenvironmental conditions. One of the exceptions are plants that produced the spores

Densoisporites, which are closely related to the plant Family Pleuromeiaceae. Pleuromeiaceae is related to Isoetes L. (quillworts), a modern heterosporous lycophyte. Based on this affinity,

Densoisporites spores could have been produced by parent plants that lived in tidally-influenced environments as the morphology of the parent plants appears adapted for water dispersal and have been found in tidally-influenced strata (Abbink et al., 2004).

Ferns

The majority of Mesozoic ferns are interpreted to grow in subtropical to tropical environments under moist conditions, similar to extant ferns. These include lowland environments such as swamps, marshes, riparian habitats, forest understories and riverbanks where they formed peat (van Konijnenburg-van Cittert, 2002). Ferns that tend to occur near river banks are from the families Osmundaceae, Schizaceae, Dicksoniaceae, Cyatheaceae, and Pteridiceae (Abbink et al.,

2004). A minority of these plants (e.g., Schizaceae, Gleicheniaceae) can tolerate direct sunlight and grow in open habitats (Abbink et al., 2004; Galloway et al., 2015). Spores from Matoniaceae ferns are known to grow on mountain slopes, which have warm and humid days that are juxtaposed against colder nights. Mesozoic Osmundaceae ferns have been found in floodplain deposits (van

Konijnenburg-van Cittert, 2002). Dicksoniaceae ferns, which are often included with Cyathaceae ferns, first appeared in the Late Triassic and mainly tolerated warm, moist habitats. An exception is the Late Jurassic-Early Cretaceous Onychiopsis psilotoides Stokes and Webb, a fern capable of withstanding high-stress conditions such as high salinity (van Konijnenburg-van Cittert, 2002 and references therein). Extant Dicksoniaceae ferns are mainly limited to tropical and temperate rain forests in the Southern Hemisphere (van Konijnenburg-van Cittert, 2002). Cyathaceae ferns typically grow in tropical and subtropical environments, and are considered to have had similar ecological tolerances during the Mesozoic (van Konijnenburg-van Cittert, 2002). Generally, much of the fern, fern allies and bryophyte spores are thought to occur in lowland and/or river environments, although there are exceptions (Abbink et al., 2004; Galloway et al., 2015).

Conifers

Pollen grains found in the Martin Creek section are primarily derived from conifers from the families Araucariaceae, Cheirolepidiaceae, Cupressaceae and/or Taxaceae, and Pinaceae

(Table 1).

Araucariaceae conifers were likely the parent plants of Araucariacites australis pollen; there are only three extant genera, Agathis Salisbury, Araucaria Jussieu and Wollemia Jones, Hill and Allen, which are now limited to the Southern Hemisphere (Stewart and Rothwell, 1993;

Stockey, 1994; Kunzmann, 2007). Araucariacites is considered a pollen genus for both fossil and extant Araucariaceae as they are virtually indistinguishable from one another (Kunzmann, 2007).

Araucariaceae originated in the Triassic and experienced an increase in abundance and diversity in the Jurassic, where they had a worldwide distribution before continental separation of Pangea.

Climate and environmental change associated with the break-up of Pangea during the Jurassic and late competition associated with the arrival and diversification of angiosperms restricted these

plants to moist, mesothermal environments by the Late Cretaceous (Kershaw and Wagstaff, 2001;

Kunzmann, 2007). The majority of living species of Araucariaceae now live in moist forests at a range of elevations from sea-level or lowland tropics to tropical highlands (de Laubenfels, 1984a), with Agathis being the most tropical out of all extant conifers (Kunzmann, 2007). Generally, the presence of Araucariaceae plants is associated with the presence of rainforest vegetation or environments suitable for rainforests (Kershaw and Wagstaff, 2001).

Cheirolepidiaceae plants were large arborescent conifers that had a high diversity and an advanced reproductive system (Stewart and Rothwell, 1993). This now extinct plant family was known to have thrived in warm, xerophytic and thermophilous habitats (Alvin, 1982; Galloway et al., 2015 and references therein), such as well-drained slopes in upland areas or in lowland coastal environments (Srivastava, 1976). Classopollis classoides pollen are thought to have been produced by cheirolepidiacean conifers. Large amounts of Classopollis pollen are commonly indicators of arid or saline conditions.

Plant representatives from the family Cupressaceae are thought to have originated during the Triassic and diversified during the breakup of Pangea (Mao et al., 2012). They are widely distributed worldwide, occurring on all continents with the exception of Antartica (Mao et al.,

2012), and mainly thrive on the margins of tropical and subtropical highlands (de Laubenfels,

1984b). The Cupressaceae clade has the greatest modern habitat diversity among the conifers, with some extant species occurring in deciduous swamplands (e.g., Taxodium Richard) while others, although rare, occur in deserts, as is the case with Cupressus L. (Pittermann et al., 2012). Mesozoic habitats for Cupressaceae were mainly mesic-hydric, which is reflected in some modern species like Sequoia Endlicher (redwoods). However, the majority of extant Cupressaceae are found in areas with xeric conditions as their ancestral plants began adapting to increasingly arid conditions

exhibited during the Cenozoic (Pittermann et al., 2012). Perinopollenites and CT pollen are both thought to have been derived from Cupressaceae conifers. CT pollen also has affinities with the

Family Taxaceae (yews), which comprises six extant genera that are mainly confined to the

Northern Hemisphere (de Laubenfels, 1984c; Elpe et al., 2018). Extant Taxaceae conifers are known to thrive in moist, temperate conditions or as understory or canopy plants (de Laubenfels,

1984c). In general, modern representatives of Cupressaceae and Taxaceae, which are sister conifer families, are known to grow in more temperate environments such as in upland regions, although these plants may also form stable plant communities in more damp, lowland regions (Galloway et al., 2013).

There exists only ten extant genera in the Family Pinaceae and they are mainly limited to the Northern Hemisphere (Stewart and Rothwell, 1993). These ten genera are Pinus L., Picea

Miller, Abies Miller, Larix Miller, Tsuga Carrière, Keteleeria Carrière, Pseudotsuga Carrière,

Cedrus Trew, Cathaya Chun and Kuang, and Pseudolarix Gordon (Stewart and Rothwell, 1993).

In the Martin Creek section, Cerebropollenites mesozoicus, Laricoidites magnus, and undifferentiated bisaccate pollen were identified as belonging to the Family Pinaceae.

Cerebropollenites pollen are closely related to the extant plant Tsuga, which is shade-tolerant and frequently thrives in mesic, temperate environments (Rogers, 1978; Calcote, 2003; Shang and

Zavada, 2003; Galloway et al., 2013, 2015). The nearest living plant relative of the Pinaceae pollen grain Laricoidites is represented by Larix (Wang et al., 2010). In North America, three species of

Larix exist: L. laricina (Du Roi) K. Koch (tamarack), L. lyallii Parlatore (alpine larch), and L. occidentalis Nuttall (western larch). Larix laricina is the most widely distributed larch across

North America and can grow in a range of climatic conditions, although they most commonly thrive on cold, wet to moist soil (Johnston, 1990). Larix lyallii can thrive in mountainous areas

where conditions are cool-temperate and fairly humid (Arno and Habeck, 1972; Arno, 1990). Larix occidentalis is a deciduous conifer that occurs in mountain valleys or on poorly-drained slopes, therefore requiring cool, moist conditions to thrive (Schmidt et al., 1976; Schmidt and Shearer,

1990). Large relative abundances of Laricoidites magnus pollen is therefore interpreted to reflect cool and humid conditions.

Pinaceae-type bisaccate pollen are thought to have been derived from conifers that likely existed in upland, arid or well-drained communities (Galloway et al., 2015 and references therein).

Bisaccate pollen are named for the two air sacs located on either side of the main body of the pollen grain, which allow for a greater range of dispersal via wind pollination. Bisaccate pollen are also dependent on fluvial transportation and are subject to influence by currents. Upon deposition in water, these pollen grains can act similarly to small sedimentary grains like silt and clay (Heusser, 1988), and can therefore be overrepresented in marine sediment where they are ultimately deposited (Suc and Drivaliari, 1991). The majority of bisaccate pollen found in marine sediment are derived from those conifers located near major drainage systems (Heusser and

Balsam, 1977). Bisaccate pollen can therefore be found in greater abundance further from the source material in comparison to other spores and pollen grains.

Cycads, Ginkgos, and Gnetophytes

Cycads were widespread during the Mesozoic and had a more diverse habitat than extant species, with some plants growing in drier, tropical regions, and others thriving in rain forests

(Abbink et al., 2004). Extant species are generally distributed in the tropics and subtropics (Abbink et al., 2004; Raven et al., 2013). Gingkoales may have a close affinity with Cycadales, although the nature of their relationship is unclear (Zhou, 2009). There exists only one representative species of the ginkgophytes, Ginkgo biloba L., which is morphologically very similar to their Mesozoic

and Cenozoic counterparts (Royer et al., 2003). Ginkgos no longer exist in the wild (Raven et al.,

2013), so knowledge of their paleoecology is limited. Nevertheless, G. biloba is known to have a range of ecological tolerances, capable of being cultivated in Mediterranean to cold temperate climates (Zhou, 2009). Late Cretaceous and Cenozoic Ginkgo were interpreted to have been confined mainly to disturbed streamside and levee environments, likely growing in riparian habitats before being displaced by angiosperms, which began dominating paleoflora from the Late

Cretaceous onward (Royer et al., 2003). Overall, Late Triassic to Early Cretaceous ginkgoaleans are thought to have been highly adaptable and capable of thriving in a diverse range of climates, from hot and dry coastal plains to wet and temperate lowland, riparian habitats, although most ginkgoaleans are found in higher abundances and with greater diversity in areas with mesic, warm temperate climates (Zhou, 2009).

Gnetophytes have angiosperm-like features, although their evolutionary history remains uncertain (Abbink et al., 2004; Raven et al., 2013; Ickert-Bond and Renner, 2015). In the Martin

Creek section, only Eucommiidites troedsonii pollen are interpreted to have an affinity with gnetophytes. Extant gnetophytes are found in a diverse range of habitats, from tropical forests e.g.,

Gnetum L., to deserts, e.g., Ephedra L. and Welwitschia Hooker (Raven et al., 2013; Ickert-Bond and Renner, 2015). Abbink et al. (2004a) placed Eucommiidites troedsonii pollen in their Lowland and/or River sporomorph ecogroup (SEG), and note that the parent plant likely grew in drier and warmer environments.

Paleoenvironment

Q-mode Cluster Analysis and Detrended Correspondence Analysis (DCA)

Q-mode cluster analysis is first used to evaluate whether the palynological clusters at

Martin Creek are stratigraphically distinct (Fig. 7). The resulting clusters from Q-mode cluster

analysis individually correspond to broadly different stratigraphic positions. Stratigraphically- influenced clusters suggest that changes in palynological content are an effect of facies change or paleoenviromental change that influenced the constituents of the plant community. Cluster 3 in Q- mode cluster analysis is of interest as it represents a complete series of consecutive samples that occur between 0 and 10.8 m. The J/K transition is thought to exist 6 m, at which the contact between the “arenaceous” and “red-weathering” members is located. However, the separation of this cluster is likely primarily influenced by a facies change from sandstone- to mudstone- dominated strata. The lithostratigraphy right above the “arenaceous member” is primarily mudstone but is also interspersed with thin sandstone interbeds that occur until 10.8 m. Cluster 3 appears to separate the facies containing sandstone from those facies that are mudstone-dominated.

The palynological variability of the samples is therefore not sufficient to distinguish the J/K transition through Q-mode cluster analysis. Although a facies change is likely the cause of Cluster

3, the other clusters are also somewhat stratigraphically separated, and is therefore interpreted to represent changes in the plant paleocommunity that could have been influenced by changes in climate.

To further understand the major factors that influenced the distribution of palynomorphs identified in this study, DCA was used to interpret principal variables that may have contributed to species ordination (Fig. 9). Based on the ecologies of plant species (Table 1) that plot near the extremes of each axis, DCA1 is interpreted as temperature-influenced whereas DCA2 is interpreted as moisture-influenced. Temperature and moisture are known to heavily influence the distribution of plants, particularly woody plants (Pittermann et al., 2012), and are, in turn, heavily influenced by changes in climate. Moisture was chosen as one of the major factors because of the taxa with known ecological tolerances plotting on each extreme of the DCA2 axis: the group

plotting near the extreme positive values (unclassified fern spores and Araucariacites, Entylissa,

Perinopollenites pollen) on the axis DCA2 is interpreted to represent environments with high effective moisture, whereas the group that plots on the extreme negative values on the axis DCA2

(Classopollis pollen) is interpreted to represent more xerophytic environments. Temperature was interpreted as the other major influence on species ordination based on the distribution of CT pollen and Eucommiidites troedsonii pollen on the DCA plot. The ecological tolerances of the parent plants of the palynomorphs across DCA1 are interpreted to represent a scale of temperate

(CT pollen) to warmer temperatures (Eucommiidites pollen).

While DCA helps make broad interpretations on the main factors that influenced species ordination, changes in palynoassemblages through time cannot be observed. Interpreting paleoenvironmental change through time may be assisted through the use of R-mode cluster analysis and CONISS, the combination of which permit better reconstruction of plant paleocommunities.

R-mode Cluster Analysis

R-mode cluster analysis separates the palynological content preserved in samples collected from the Martin Creek section into three informal groups: a spore assemblage, pollen assemblage, and a CT pollen assemblage (Fig. 8). These groups are interpreted as representing plants that tend to co-occur at a given time in the studied area. The clusters generated from R-mode cluster analysis therefore represents different habitats that reflect the paleoecological tolerances of parent plants from which palynomorphs were derived. Fluctuations in sporomorph assemblages are assumed to be a result of changes in the terrestrial plant community that are controlled mainly by temperature and moisture (Abbink et al., 2004; Galloway et al., 2015) as interpreted from the DCA plot, which are influenced by geography and/or changes in climate.

The “spore assemblage” is the most diverse group, containing the majority of the bryophyte and fern spores. Only those bryophyte spores from the phyla Bryophyta and Marchantiophyta were observed in the Martin Creek section. The “spore assemblage” cluster is therefore interpreted to represent a lowland environment with relatively high moisture conditions.

The “pollen assemblage” cluster is composed of a combination of fern and gymnosperm palynomorphs. Gymnosperms have more specific ecological tolerances in comparison to the wide range of ecological tolerances observed in ferns and bryophytes. Within the Class Pinopsida, occurrences and abundances of Cerebropollenites mesozoicus, Araucariacites australis,

Laricoidites magnus, Classopollis classoides, and undifferentiated bisaccate pollen can provide information on past conditions. This cluster is thought to represent upland plant communities that are characterized by relatively drier or well-drained conditions and cooler temperatures in comparison to lowland environments.

The “CT assemblage” is singularly represented by CT pollen, which are derived from

Cupressaceae and Taxaceae that dominate the palynoassemblage of the Martin Creek section.

Cupressaceae and Taxaceae commonly thrive in more temperate environments where they form upland communities while also being capable of growing in more damp, lowland regions where they may even form climax communities (Galloway et al., 2013, 2015). This cluster is interpreted to represent an intermediate environment between lowland and upland habitats.

R-mode cluster analysis provides information on the paleoecological communities that may have existed during the J/K transition. The “spore assemblage” cluster seen in this study could represent an understory paleocommunity of ferns situated in a lowland area, whereas the “pollen assemblage” cluster could represent more upland paleocommunities. The “CT pollen assemblage” may represent an intermediate region like a slope area between lowland and upland communities.

When combined with CONISS, we are able to observe stratigraphically-constrained taxa grouped via R-mode cluster analysis. The relative abundances of each taxa are extrapolated against section meterage, and we can therefore see changes in palynoassemblages through time. These fluctuations in the palynoassemblage are interpreted as representing a response to changes in temperature and moisture as inferred from the DCA plot, which are influenced by changes in climate.

Stratigraphically Constrained Cluster Analysis (CONISS)

The oldest cluster in the Martin Creek section, Cluster D, contains the putative location of the J/K transition. Cluster D is a mixed assemblage notable for the relatively high abundances of

Baculatisporites comaumensis spores, Pinaceae-type bisaccate pollen, Cerebropollenites mesozoicus pollen, and Cyathidites australis spores. Relative abundances of palynormophs in this cluster remain relatively stable. Baculatisporites and Cyathidites spores are derived from ferns

(Polypodiopsida), with the former assigned to the family Osmundaceae and the latter to either

Dicksoniaceae and/or Cyatheaceae (Table 1). Cluster D is thought to represent a somewhat humid, warm-temperate climate due to the relatively high abundances of Cerebropollenites mesozoicus pollen alongside spores derived from Osmundaceae and Dicksoniaceae/Cyatheaceae ferns.

Possible upland communities may have also existed concurrently as represented by the presence of bisaccate pollen. Cluster D is also characterized by a relatively higher proportion of dinocyst content (>0-6%), which suggests the presence of somewhat more distal facies in a succession of predominantly proximal facies, and therefore may be indicative of transgressive beds.

Predominantly humid conditions are therefore thought to have occcured throughout the latest

Tithonian and early Berriasian, similar to the palynological interpretations of the North Sea

(Abbink et al., 2001b), although the palynoassemblage at Martin Creek lacks a substantial enough shift in content to confidently distinguish the J/K boundary.

Two major palynological events are apparent in the clusters generated by CONISS, the first of which occurs in Cluster C. The lowermost beds in Cluster C are first characterized by a maximum in dinocyst abundance (7.47%); this suggests the presence of the distal-most facies in the succession at 28-30 metres and may represent a flooding surface. Dinocyst content decreases thereafter, which indicate the reoccurrence of regressive beds. Typically, a major transgression can be observed in the uppermost beds of the arenaceous member, which may be indicative of the J/K boundary (Dixon, 1992). This trend is also observed in southern Sweden (Lindstrőm and Erlstrőm,

2011), where a relatively high abundance of dinocysts (55%) is inferred to represent a maximum flooding surface that occurred in the lowermost Berriasian (Subcraspedites preplicomphalus

Zone). Relative abundance of dinocysts decreases to 20% in the lower Berriasian (Subcraspedites lamplughi Zone). Both the S. preplicomphalus and S. lamplughi zones are roughly equivalent to the Buchia okensis Zone (Zakharov, 1987), which marks the J/K boundary in the Canadian Arctic

(Jeletzky, 1960, 1973, 1984; Fig. 5). The peak in dinocyst content in the Martin Creek section is concurrent with a relatively high abundance of Perinopollenites elatoides and Araucariacites australis pollen. Immediately succeeding these trends is a peak in the relative abundance of CT pollen from 44% to 82% at 32 m in this section. This trend continues for several metres between

32-40 m, with a range in relative abundance of CT pollen between 65-96% before returning to

>50% after 40 m (Fig. 8). The increase in Araucariaceae- and Cupressaceae-type pollen and the dominance of CT pollen in Cluster C is therefore interpreted to indicate a period of progressively increasing humidity relative to Cluster D, with prevailing temperate conditions.

The period of relatively high humidity in Cluster C is then followed by the second major palynological event, characterized by a peak in Classopollis classoides pollen between 46-56 m

(Cluster B2; Fig. 8). Relative abundances of Classopollis pollen preserved in the Martin Creek material are generally consistent throughout the Husky Formation with the exception of Cluster

B2, where Classopollis pollen peaks to a relative abundance of ~25%. The original tetrads in which

Classopollis pollen were produced were also preserved in the Husky Formation in the Martin

Creek section. In order for such preservation to occur, minor transportation of Classopollis pollen prior to deposition is likely (Traverse, 2007), which suggests that their parent plants may have grown in coastal areas under dry climates (Pocock and Jansonius, 1961; Galloway et al., 2015).

While relatively high abundances of Classopollis pollen typically indicate a dry event, it is not the dominant pollen type in the Martin Creek section, and is therefore suggestive of seasonally arid conditions rather than a prevailing dry climate (Galloway et al., 2013). Cluster B2 is also characterized by comparatively higher abundances of Stereisporites antiquasporites spores in relation to other clusters. Stereisporites antiquasporites spores are derived from plants from the

Class Sphagnopsida, the bog/peat mosses. Modern peat mosses are categorized into the extant genus, Sphagnum Ochyra, which is closely related to the fossil order Protosphagnales (Raven et al., 2013). Sphagnum is distributed globally, but typically grow in lowland areas with poor drainage (Raven et al., 2013). The increased presence of both Classopollis pollen and

Stereisporites spores, with the former occurring in higher abundances, is therefore thought to represent warm, seasonally arid conditions in Cluster B2.

Cluster B1 is a mixed assemblage that maintains a generally stable relative abundance throughout. Laricoidites magnus pollen reaches a maximum relative abundance in this cluster at

~12%, which corresponds to a decrease in the relative abundance of bisaccate pollen at 74-76 m.

The increase in abundance in Laricoidites pollen and corresponding decrease in bisaccate pollen, which typically represent drier conditions, therefore suggest a brief period of higher humidity alongside a temperate climate within Cluster B1 (74-76 m). The prevalence of CT, bisaccate, and

Laricoidites magnus pollen relative to other ecologically important taxa suggests that Cluster B1 was generally characterized by moderately high humidity and cool-temperate conditions, with periods of slightly drier conditions represented by instances of increased relative abundance of bisaccate pollen.

Cluster A is a mixed assemblage that is characterized by a minimum in CT pollen to ~8% and a peak in Araucariacites australis pollen and Unknown Palynomorph A. At 88 m, there is a concurrent trend of increasing relative abundance of CT pollen and a peak in Araucariacites australis pollen. This trend may indicate a period of increasingly moist conditions and relatively cooler temperatures (Abbink et al., 2001b; Lindstrőm and Erlstrőm, 2011). A secondary peak in dinocyst percentage (up to 6%) occurring simultaneously may indicate the reoccurrence of a secondary flooding surface. A second flooding event has also been observed in southern Sweden during the mid-Berriasian (Subcraspedites lamplughi Zone), where dinocyst percentage reached

70% (Lindstrőm and Erlstrőm, 2011). The last palynological signature that may be significant is the spike in relative abundance of Unknown Palynomorph A, which occurs sporadically throughout the section but appears in highest abundances at 92 metres. These unidentified palynomorphs are spherical and consistently smaller than the other palynomorphs encountered in the Martin Creek section, and do not resemble any of the common forms seen in acritarchs or in palynodebris. Unknown Palynomorph A may be some type of aberrant spore/pollen or an undetermined fungal/algal spore, but it cannot be used for paleoecological interpretations as its identification remains uncertain. Cluster A also experiences slightly higher abundances of

Densoisporites spp. spores, a peak in Entylissa spp. pollen, and a general overall increase in abundance of spores derived from Polypodiopsida, Lycopodiopsida, Bryophyta, and

Marchantiophyta. Together, these increases in relative abundance suggest a paleoenvironment characterized by increased moisture and temperate to warm conditions.

Based on the quantitative multivariate statistical analyses of the palynoassemblage of the

Martin Creek section, the J/K transition in the Canadian Arctic is characterized by a predominantly humid climate, with a seasonally arid phase in the early Berriasian. Temperature is interpreted to have varied between relatively warm to cool-temperate. This interpretation is consistent with the temperate climate that characterized the Arctic Subprovince during the J/K transition (Schneider et al., 2018b).

In comparison, the J/K transition in northwest Europe is often characterized by a climatic shift from semi-arid to humid conditions (e.g., Abbink et al., 2001b; Hunt, 2004; Lindstrőm and

Erlstrőm, 2011; Főllmi, 2012; Schneider et al., 2018b), with fluctuating phases of humidity and aridity in a tropical to subtropical climate (Főllmi, 2012; Schneider et al., 2018b). The earliest

Cretaceous in this region was thought to have been first dominated by Cheirolepidiaceae conifer forests and then by mixed forests with a diverse fern understory, with a paleoclimate paralleling that of the present-day Mediterranean: arid, hot summers juxtaposed against wetter, cooler winters

(Schneider et al., 2018b). In the North Sea, the possible J/K boundary marks the end of an arid phase that occurred during the late Kimmeridgian and Portlandian, with the basal Cretaceous being characterized by an increase in humidity and relatively high temperatures as exemplified by the large diversity of fern spores replacing those palynomorphs derived from drier conifer forests

(Abbink et al., 2001b and references therein). The palynological signature in NW Europe is therefore defined by a decrease in abundance of Classopollis pollen in sedimentary records and an

increase in diversity and abundance of fern spores. The results of the Martin Creek section indicate that the climate during the J/K transition in the Canadian Arctic may have been comparatively less arid and cooler.

CONCLUSION

Palynomorphs provide context for paleoenvironmental interpretations during the Mesozoic but can be particularly difficult to use in terms of biostratigraphic correlation. Defining the J/K boundary in Arctic Canada remains dependent mainly on buchiids. Despite the rarity of dinocysts, a marine-to-terrestrial ratio may be capable of reflecting regressive and transgressive phases and potential flooding surfaces. Quantitative multivariate statistical analyses of palynomorphs yield promising results in interpreting paleoenvironmental changes experienced in one of the regions within the Boreal Realm. The latest Tithonian – earliest Berriasian transition in the Canadian

Arctic is characterized by predominantly humid conditions with a seasonally arid phase.

Palynological trends in NW Europe have documented a gradual decline in Classopollis pollen and an increase in diversity and abundance of spores during the J/K transition and into the Early

Cretaceous. Northwest Europe is therefore thought to have experienced a semi-arid climate in the latest Tithonian to earliest Berriasian before experiencing a transition to predominantly humid conditions that characterized the remainder of the Cretaceous Period. The results of this study indicate that the climate during the J/K transition in the Canadian Arctic may have been comparatively less arid and cooler than Tethyan locations.

100

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CHAPTER 3: LOWER CRETACEOUS PALYNOSTRATIGRAPHY OF THE AKLAVIK

RANGE, NORTHWEST TERRITORIES, ARCTIC CANADA

ABSTRACT

Quantitative palynostratigraphy was conducted for a unit considered either the Husky or

Martin Creek formations in the Aklavik Range, northern Richardson Mountains, Northwest

Territories. Based on the dinoflagellate cyst assemblage containing the potential biostratigraphically significant species Oligosphaeridium cf. tenuiprocessum, a revised late Albian age is herein proposed for the strata, which may be stratigraphically attributed to the Arctic Red

Formation. Multivariate statistical analyses and ordination techniques suggest that a typical coastal wetland paleoenvironment supported a local plant community dominated by parent plants from which Osmundacidites wellmannii and Densoisporites spp. spores, and Cupressaceae-Taxaceae and Pinaceae pollen were derived. Stratigraphically constrained cluster analysis was used to delineate three clusters (A-C) from a small measured section to refine the paleoenvironmental reconstruction over the time interval studied. The increased relative abundance and diversity of fern spores, such as those from the Class Osmundaceae, suggest that high effective moisture conditions prevailed in the Richardson Mountains at this time. Paleotemperatures may have transitioned from cool to warm based on the decrease in relative abundance of Laricoidites magnus and undifferentiated bisaccate pollen and increase of Cerebropollenites mesozoicus, Classopollis classoides, and Cupressaceae-Taxaceae pollen.

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INTRODUCTION

Lower Cretaceous strata outcrop extensively in the Beaufort-Mackenzie area. However, chronostratigraphy of the Cretaceous is limited by difficulties in correlating between newly formed paleogeographic realms following the break-up of Pangea in the Late Jurassic. The majority of

Mesozoic boundary stratotypes are located in Western Europe in the Tethyan Realm (Gradstein et al., 2012). Lower Cretaceous subdivisions are particularly problematic, as only the Albian Stage has been ratified by a Global Stratotype Section and Point (GSSP) out of the six existing subdivisions (Kennedy et al., 2017). While ammonites have traditionally been used to define the stages of the Cretaceous Period, the pervasive biogeographical segregation of this group between the Tethyan Realm and the Boreal Realm have precluded correlation of strata between these two geographical areas. Chronostratigraphic correlation of Cretaceous rock units is therefore difficult to accomplish in polar regions such as the Canadian Arctic as Boreal and Tethyan biota did not share any common species in Late Jurassic – Early Cretaceous times.

Two main marine biotic provinces existed in Canada during the Cretaceous and were defined primarily by endemic ammonite fauna: the North Pacific Biotic Province and the North

American Boreal Province (Jeletzky, 1971). The North Pacific Biotic Province had close associations with the Tethyan Realm, whereas the North American Boreal Province was closely related to the Boreal Realm. These two provinces were geographically isolated from each other throughout the Mesozoic due to the breakup of Pangea (Jeletzky, 1971). During the Berriasian to

Aptian, the North American Boreal Province was restricted within the Canadian Arctic

Archipelago and the Beaufort-Mackenzie region of Yukon and the Northwest Territories (Jeletzky,

1971). Towards the Late Cretaceous, this province occupied parts of the Interior Plains and the

Eastern Cordillera as Boreal seas expanded southward towards the Tethyan seas in the Gulf of

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Mexico (Jeletzky, 1971). While the opening seaway connected the North American Boreal

Province with the seas of the Gulf Coast in the latest middle Albian, provincialism remained strong among marine invertebrates (Jeletzky, 1971). The marine Boreal realm of Canada is also characterized by a paucity of ammonites and a general lack of diversity (Jeletzky, 1971).

Ammonite biostratigraphy must therefore be supplemented by magnetostratigraphy, geochemistry, and/or biostratigraphic correlation using microfossils to help provide a framework in which correlation between Boreal and Tethyan Realms may be accomplished. Much of the biostratigraphic work for the Lower Cretaceous of Arctic Canada was completed by Jeletzky (e.g.,

1958, 1960, 1961, 1967, 1971, 1972), who proposed a zonation scheme primarily based on Buchia

Rouiller, 1845 species. Upper Cretaceous Buchia zones have been identified for the Canadian

Arctic Islands (Jeletzky, 1971), but have not been identified for the Canadian Arctic mainland in the northwest (Dixon, 1993). In the area between the northern Richardson Mountains and

Mackenzie Delta, the Albian biozones remain undetermined (Dixon, 1993). Most palynological zonations have likewise been created for the Lower Cretaceous (Brideaux and Fisher, 1976;

Doerenkamp et al., 1976; Pocock, 1976; McIntyre and Brideaux, 1980; Fensome, 1983, 1987).

Few palynological zonation schemes were created for the Upper Cretaceous (McIntyre, 1974;

Doerenkamp et al., 1976; Andrews, 2012), and only one encompasses the entire Cretaceous Period

(Bujak and Scott, 1984). There is therefore a need for a review on the biostratigraphy in the

Aklavik Range area.

In comparison to ammonites and Buchia fauna, palynomorphs such as fossil spores, pollen grains, and dinoflagellate cysts (i.e., dinocysts) are pervasive and exceptionally well-preserved in both Boreal and Tethyan realms. Due to relatively latitudinal equable climatic conditions (Hallam,

1985; Jenkyns et al., 2004; Hay, 2008), there was much less provincialism of terrestrial vegetation

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as compared to other marine organisms (Fægri and Iversen, 1989). Palynomorphs are deposited in great abundance and can be used to reconstruct past terrestrial and marine ecosystems in which they occurred (Fægri and Iversen 1989; Galloway et al., 2013). The source plants of palynomorphs are often considered the best paleontological evidence for temperature gradients due to their sensitivity to changes in climate (Hallam, 1985). This is particularly useful for the Cretaceous

Period, which was a time of overall extreme global warmth, with some pronounced cold snaps in high northern latitudes (Price and Nunn, 2010; Price and Passey, 2013; Galloway et al., 2015;

Grasby et al., 2017). Although the Mesozoic likely had a more equable climate than the

Quaternary, polar paleotemperatures remain somewhat debatable (Herman and Spicer, 1996;

Price, 1999; Price et al., 2000; Price and Nunn, 2010; Galloway et al., 2012, 2015; Price and

Passey, 2013; Grasby et al., 2017).

Using quantitative palynology, this study aims to improve biostratigraphic correlation of

Lower Cretaceous outcrops in the Beaufort-Mackenzie area. Multivariate statistical analyses and ordination techniques are conducted to investigate paleoenvironmental conditions of northern polar regions. This will provide further insight into the polar paleotemperatures of this time period and allow an improved understanding of the complexities of the greenhouse climate conditions during the Cretaceous. Improved age control of strata of the northern Richardson Mountains is also necessary to refine the timing and history of tectonism in the area resulting in the formation of the Arctic Ocean and Canada Basin. The subsequent establishment of new seaways and linkages can also provide insight into their effects on global climate. An improved understanding of the tectonic history of the area is likewise important for exploration of natural resources within the

Canada Basin.

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GEOLOGIC SETTING AND STUDY AREA

The Beaufort-Mackenzie region was dominated by extensional tectonics resulting in the rifting of the Canada Basin during Late Jurassic to early Aptian time, forming multiple uplifts and grabens in the area that were filled with sediment originating off the craton in the east and southeast

(Dixon, 1992, 1993; Fig. 1). This extensional setting is reflected in the abundance and magnitude of normal faults that proliferated along the northwestern margin of the Eskimo Lakes Arch (Dixon and Dietrich, 1990). A trangressive episode occurred between the Oxfordian to Valanginian and was succeeded by a regressive episode in the late Valanginian to Hauterivian. Shallow marine and deltaic facies dominated the Berriasian to Hauterivian successions (Balkwill et al., 1983). From the Valanginian to Hauterivian, north- and westward delta progradation occurred away from basin margins, allowing for extensive erosion in marginal areas (Balkwill et al., 1983). In the Mackenzie

Delta area, a regional sub- or intra-Hauterivian unconformity marks the onset of rifting of the

Canada Basin that resulted in intense uplift and widespread erosion (Dixon, 1982). This unconformity exists at the base of the Mount Goodenough Formation in the Beaufort-Mackenzie region, at the base of the Isachsen Formation on Banks Island (Miall, 1979), and in the lower part of the Isachsen Formation in the Sverdrup Basin (Galloway et al., 2013). From the Hauterivian to

Cenomanian, the Canadian Arctic was characterized by accelerated subsidence rates and accumulation of thick clastic successions (Balkwill et al., 1983). Extensional tectonics continued into the Albian, when Cordilleran compression began to fill in deep-water troughs with sediment with a western-southwestern provenance (Dixon, 1992). Albian strata therefore consist mainly of marine shelf shales (Dixon and Dietrich, 1990). Rifting of the Canada Basin ended by the beginning of the Cenomanian, and was succeeded by sea-floor spreading contemporaneous to

Cordilleran orogenesis (Dixon, 1992). The Lower and Upper Cretaceous rocks of the study area

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Figure 1: a) General map of structural elements in the Beaufort-Mackenzie region (i.e., Brooks-

Mackenzie Basin) in the northwesternmost area of the Northwest Territories, with area of interest boxed in red (modified from McNeil et al., 2001; Fensome, 1987; Dixon and Dietrich, 1990;

Dixon, 2004). b) Geologic map of the Treeless Creek study area with outcrop location

(67.86473°N, 135.62909°W) represented by the red dot and other localities mentioned in the text

(modified from Norris, 1981). Red square in B is the same area as in Fig. 10.

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125 are separated by a major regional unconformity (Dixon and Dietrich, 1990). During the

Cenomanian and into possibly the late Maastrichtian, the margins of the Canada Basin began their main phase of drifting (Dixon, 1992).

The study area is located at Treeless Creek, an informal term used frequently by Jeletzky for a creek south of Willow River and north of Rat River in the Aklavik Range. Coordinates for the studied outcrop are 67.86473°N, 135.62909°W (Fig. 1). The Treeless Creek area is not well understood and the geologic map by Norris (1981) may be incorrect (T. Poulton, pers. comm.,

2018) due to covered structural features including prevalent faults, most of which have unknown displacement, and regional unconformities (Young and Robertson, 1984; Dixon, 1993).

Husky Formation

The Husky Formation was first defined by Jeletzky (1967) as the “lower shale-siltstone division”. The reference section lies in the Aklavik Range in the northern Richardson Mountains, where it can be divided into four informal divisions: in ascending order they are the lower, arenaceous, red-weathering, and upper members (Dixon, 1982, 1992; Figs. 2, 3). The formation is dominated by mudstone in the Aklavik Range and often contains marine fossils. The formation is interpreted to have been deposited calm waters in a marine setting (Dixon, 1982). The lower and arenaceous members have a higher sand content, whereas the upper and red-weathering members are composed of predominantly dark-coloured shales and siltstones (Dixon, 1982). The red- weathering and upper members represent a regressive phase in a coarsening upward cycle that transitions into the overlying sandstones of the Martin Creek Formation (Dixon, 1982). The lower and arenaceous members are Jurassic in age, whereas the upper and red-weathering members are

Early Cretaceous in age based on the presence of bivalves from the Buchia okensis Pavlow, 1907,

B. uncitoides Pavlow, 1907, and possibly the B. volgensis Lahusen, 1888 zones (Jeletzky, 1967;

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Figure 2: Extensive outcrops of the Husky Formation at Treeless Creek (GSC section 82-10;

Dixon, 1992). 1 - Lower member; 2 - Arenaceous member; 3 - Red-weathering and upper members.

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Figure 3: Lithostratigraphic correlation chart of Lower Cretaceous formations of north and northwestern Canada (modified from Bringué et al., 2018 and references therein).

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Dixon, 1992). The Husky Formation is therefore interpreted to be Oxfordian to early Berriasian in age and spans the Jurassic-Cretaceous boundary (Dixon, 2004).

Martin Creek Formation

The Martin Creek Formation is the sandstone-dominant basal member of the Parsons

Group, which comprises the Martin Creek, McGuire and Kamik formations (Dixon, 1982; Fig. 3).

The Martin Creek formation was previously known as the “buff sandstone member” of the “lower sandstone division” and is thought to be Berriasian to early Valanginian in age based on the presence of bivalves belonging to the Buchia volgensis and B. keyserlingi Lahusen, 1888 zones

(Jeletzky, 1958, 1960; Pocock, 1976; Dixon, 1982, 1992). The type section is located on the south side of Martin Creek in the northeastern Richardson Mountains, with its base and top located at

68°12’11” N, 135°35’46” W and 68°12’53” N, 135°34’55” W, respectively (Dixon and Jeletzky,

1991). The Martin Creek Formation is interpreted to have a nearshore or lower shoreface depositional environment that was dominated by strong wave action (Dixon and Jeletzky, 1991).

The formation can have an abrupt or gradational contact with the underlying Husky Formation and is composed of a series of coarsening-upward cycles on the western slopes of the northern

Richardson Mountains, grading upwards from bioturbated sandy mudstone to bioturbated argillaceous sandstone (Dixon, 1992; 2004). On the eastern slopes of northern Richardson

Mountains and in the subsurface of southern Mackenzie Delta and Tuktoyaktuk, a single coarsening-upward cycle brackets the Husky and Martin Creek succession (Dixon, 1982; 1992).

In the Aklavik Range, the formation is predominantly sandstone and is predicted to have a maximum thickness of 200 metres in the subsurface of the Tuktoyaktuk Peninsula (Dixon, 1992).

Strata tend to be thinner at basin margins due to erosion and depositional trends (Dixon, 2004).

The formation undergoes a facies change basinward to the shale-dominant Kingak Formation

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(Dixon, 2004). Palynomorph and foraminifera aseemblages found in the Martin Creek Formation are similar to those found in the Husky Formation, and suggest a Berriasian age (Brideaux and

Fisher, 1976; Fensome, 1987; Dixon, 1992).

McGuire Formation

The McGuire Formation abruptly overlies the Martin Creek Formation (Dixon, 2004). The formation can sometimes overlie older formations such as the Husky and Kingak formations south of McDougall Pass (Dixon, 1992). The McGuire Formation was initially defined by Jeletzky

(1961) as the “bluish grey shale division” before its name was formally established by Dixon and

Jeletzky (1991). The type section of the McGuire Formation is located on the eastern flank of a ridge located approximately 4.1 km north-northeast of Mount McGuire. Its base occurs at

67°57’54” N, 137°18’14” W (Dixon and Jeletzky, 1991). The McGuire Formation is predominantly composed of mudstone. The lower half of the formation comprises blue-grey to black, fissile shale with abundant ironstone concretions whereas the upper half of the formation has an increased silt content (Dixon and Jeletzky, 1991). The uppermost beds are predominantly composed of argillaceous sandstone with interbeds of mudstone (Dixon and Jeletzky, 1991). The formation is part of an overall coarsening-upward succession that represents the first deposits of progradation (Dixon, 1992). The depositional environment of the McGuire Formation is interpreted to have been a marine shelf with low energy conditions that allowed for mud deposition. Intermittent periods of storm activity caused sand deposition (Dixon and Jeletzky,

1991; Dixon, 1992). The lowermost beds are dated as early to middle Valanginian in age based on fossils belonging to the Buchia keyserlingi and B. inflata Lahusen, 1888 zones (Jeletzky, 1961;

Dixon and Jeletzky, 1991 and references therein; Dixon, 1992). A palynomorph and dinocyst

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assemblage that is distinctly Valanginian in age can be found in the McGuire Formation (McIntyre and Brideaux, 1980).

Kamik Formation

The Kamik Formation is sandstone-dominated, with increasing mudstone interbeds towards the top of the succession (Dixon, 1982). The contact between the Kamik Formation and underlying McGuire Formation is transitional (Dixon and Jeletzky, 1991). The Kamik Formation is the last unit of the coarsening-upward cycle that is represented by the Parsons Group (Dixon,

1982; Fig. 3). Its type section lies within the Gulf Mobil Kamik F-38 well (68°57’25”N,

133°23’54”W) between depths 3006.5 – 3236.4 metres (Dixon, 1982). The formation can be split into a lower and upper member whereby the lower member is differentiated from the upper member by having 70-80% sandstone, whereas the upper member has increasingly interbedded mudstone that constitutes more than 30% of the succession (Dixon, 1982, 1992). The lower beds represent westwardly to northwestwardly progradation of a shoreline while the upper beds represent small-scale transgressive-regressive cycles (Dixon, 1992). A paucity of age-diagnostic macrofossils limits the age constraints of the Kamik Formation (Dixon, 1992). Estimations based on microfossils suggest a Valanginian to Hauterivian age (Dixon, 1982 and references therein). At

Grizzly Gorge, the basal beds of the Kamik Formation contain bivalves identified as Buchia n. sp. aff. inflata. Buchia crassa Pavlow, 1907 and B. n. sp. aff. crassa have been found in the lowermost beds of the Kamik Formation on the eastern slopes of the Richardson Mountains (Dixon, 1992).

Bivalves from the Buchia. ex gr. inflata-sublaevis Keyserling, 1846 Zone have also been reported in stratigraphically equivalent beds (Dixon, 1992). The presence of these bivalve assemblages indicates a mid-late Valanginian age (Dixon, 1992). Age constraints also rely on the overlying

Mount Goodenough Formation, which is late Hauterivian to Barremian in age based primarily on

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ammonite and belemnite fossils. The Kamik Formation is therefore interpreted to have a maximum age range of mid-Valanginian to late Hauterivian (Dixon, 1992).

Mount Goodenough Formation

The Mount Goodenough Formation is dominated by mudstone with ironstone concretions, with thin interbeds of siltstone and very fine sandstone that increase towards the top of the succession where it transitions into the shelf to nearshore sandstones of the overlying Rat River

Formation (Dixon, 2004; Fig. 3). The name Mount Goodenough Formation was given by Dixon and Jeletzky (1991) to replace the informal names “upper shale-siltstone” and “dark-grey siltstone” divisions (Jeletzky, 1958, 1961; Dixon and Jeletzky, 1991). The type section is located on the eastern flank of Mount Goodenough in the Aklavik Range where the formation is well exposed.

The base of the section occurs 67°56’50”N and 135°24’30”W (Jeletzky, 1958; Dixon and Jeletzky,

1991; Dixon, 1992). In this area, the Mount Goodenough Formation rests with an angular unconformity on strata containing fossils from the Buchia n. sp. aff. volgensis and Tollia cf. payeri

Toula, 1874 zones (Jeletzky, 1958), indicating that these strata may belong to either the upper beds of the Husky Formation or the lower beds of the Martin Creek Formation (Pocock, 1976; Dixon and Jeletzky, 1991). This unconformity is pervasive in many of its outcrop localities, where much of the Hauterivian and part of the Barremian has been removed (Pocock, 1976). The Mount

Goodenough Formation can be divided into a lower, mudstone-dominant member and an upper member comprising interbedded mudstone and sandstone (Dixon and Jeletzky, 1991). This succession represents deposition on a mid to outer shelf. As progradation progressed, increasing amounts of sand and silt were deposited as the depositional environment transitioned into the mid- to inner shelf (Dixon, 1992). The presence of the ammonites Simbirskites Pavlow, 1891 and

Crioceratites Léveillé, 1837 and the belemnite Oxyteuthis Stolley, 1911 indicate that the Mount

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Goodenough Formation has an age range extending from the Barremian, possibly late Hauterivian, into the Aptian (Jeletzky, 1958; Pocock, 1976). This age estimation is supported by foraminifera and palynomorphs (Dixon, 1992).

Rat River Formation

The Rat River Formation was originally named “the upper sandstone division” (Jeletzky,

1958) and described from the Mount Goodenough area, although the reference section is located on the lower reaches of Rat River (Dixon, 1992). The formation overlies the Mount Goodenough

Formation via a gradational to abrupt contact and is only present in the northern Richardson

Mountains and beneath the subsurface of the Mackenzie Delta (Dixon, 2004; Fig. 3). The reference section therefore does not include the basal contact with the Mount Goodenough Formation (Dixon and Jeletzky, 1991). The Rat River Formation is predominantly characterized by sandstone with interbeds of mudstone in a series of coarsening upward cycles (Dixon, 1992, 2004). On the western flank of the Richardson Mountains, the formation can be divided into a lower and upper member.

The lower member is composed of interbedded mudstone and sandstone. The upper member is dominated by mudstone with thin sandstone interbedding (Dixon, 1992). The majority of the formation is interpreted as having been deposited on the inner to mid shelf below wave base, with an increase in mudstone towards the west representing a more distal setting (Dixon, 1992). The

Rat River Formation is considered partially equivalent to the Mount Goodenough Formation and interpreted as having an age range spanning the late Barremian to Aptian based on bivalves found in the formation, which include Aucellina aptiensis d’Orbginy, 1850 (Jeletzky, 1958, 1960; Dixon,

1992).

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Rapid Creek Formation and Albian Flysch

The Rapid Creek Formation is a phosphatic iron formation in the northwestern Richardson

Mountains (Young and Robertson, 1984; Fig. 3). The formation was previously referred to as the

“bedded ironstone and shale unit” of the Albian flysch division (Young et al., 1976). The formation encompasses the top 1000 metres of the 4000 metre-thick Albian flysch succession (Young and

Robertson, 1984). The phosphatic beds grade westward from the type section into the non- phophatic beds of the Albian flysch and thins eastward on the Cache Creek Uplift (Dixon, 1992).

At its type section, the Rapid Creek Formation is unconformably overlain by the shales of the

Boundary Creek Formation (Young and Robertson, 1984). Few age-diagnostic fossils are present within the Rapid Creek Formation, although the presence of the pelecypods Pachygrycia Jeletzky,

1981, Inoceramus anglicus Woods, 1911 and Pholadomya Sowerby, 1823 suggests an early

Albian age (Young and Robertson, 1984). The Albian flysch lies within the Blow Trough (Young et al., 1976), and is known to rest unconformably on older strata (Dixon, 1992 and references therein). The Albian flysch is predominantly composed of interbedded mudstone, siltstone and sandstone, with occasional thick successions of conglomerate and sandstone (Dixon, 1992). The flysch sequence is thought to have originated as sediment gravity-flow deposits based on the presence of Bouma-type sequences (Dixon, 1992). The few ammonites that can be found within the flysch sequence suggest an early to middle Albian age (Jeletzky, 1960, 1971).

Boundary Creek Formation

The Boundary Creek Formation is an Upper Cretaceous, mudstone-dominant succession that outcrops extensively in the eastern Yukon Coastal Plain and in the northernmost Richardson

Mountains (Young, 1975; Fig. 3). The type section is a composite section located in the northern bank of Boundary Creek where it abruptly overlies Albian strata (Dixon, 1992). Jeletzky (1958,

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1960) described an Upper Cretaceous “shale division” at Treeless Creek in the eastern Richardson

Mountains, which was considered stratigraphically equivalent to the Boundary Creek Formation

(Young, 1975). At Treeless Creek, the “shale division” is characterized by a predominance of soft, fissile mudstone with abundant concretions that oxidize to produce yellow, orange, grey areas

(Jeletzky, 1960). Ironstone concretions are also common and weather brown-red (Dixon, 1992).

Thin beds of bentonite are also present throughout the succession and weather white or yellow

(Young, 1975). While the type section has a consistent lithology (Young, 1975), the Treeless Creek section of Jeletzky (1960) can be divided into three informal members based on weathering colour: a lower member characterized by dark grey shale, a middle orange-weathering shale member, and an upper bluish-grey weathering member. Age-diagnostic pelecypods are sparse in the succession.

Some bivalve and ammonite fauna that can be found in the “shale division” at Treeless Creek include Inoceramus crippsi Mantell, 1822, Scaphites Parkinson, 1811, and Watinoceras Warren,

1930, which indicate a late Cenomanian to Turonian age (Jeletzky, 1960; Dixon, 1992). This is consistent with the age estimated from palynomorphs and dinocysts (Dixon, 1992 and references therein). In the northern Richardson Mountains, the high organic content in the Boundary Creek

Formation may represent high productivity or low-oxygen conditions, which suggests an outer shelf to slope depositional environment (Dixon, 1992).

METHODS

Palynological Samples

Sixteen organic-rich samples were collected from the Aklavik Range of the northern

Richardson Mountains, Northwest Territories, Canada, for palynological analyses to determine biostratigraphic ages and depositional environments (Fig. 4). The mudstone-dominant succession measured and sampled at Treeless Creek contains numerous red-weathering ironstone concretions.

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Figure 4: Lithostratigraphic column and composite image of Treeless Creek section, outlining a sandstone-dominant basal unit (1), a mudstone-dominant succession with multiple red ironstone concretionary beds (2), and a topmost sandstone-dominant unit (3). Samples collected for palynology are shown adjacent. Scale bar is approximated.

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The top and bottom of the mudstone unit are bordered by sandstone-dominant strata. The measured section is approximately 75 metres thick. Two samples (C-604497-98) were collected from the basal sandstone member (unit 1). Two samples were collected near the base of the upper sandstone member (C-604482-83; unit 3). Sample resolution was every two metres in the mudstone succession (unit 2). Due to time constraints, the mudstone unit was not measured in its entirety and sample resolution changed to every four metres above 60 metres in the stratigraphic section.

Thus, the thickness of the mudstone unit remains unknown.

Samples were prepared by Global GeoLab Ltd. using standard preparation procedures including acid maceration, oxidative treatment with Schulze’s solution and staining with Safranin

O. Palynological residues were then placed on glass microscope slides using liquid bioplastic. All samples are curated in the permanent collections of the Geological Survey of Canada. Spores and pollen grains were enumerated to a minimum of 300 specimens (316 ± 27, n=16) where possible and identified using an Olympus BX61 microscope at 400x and 1000x magnification. A minimum count of 300 has been shown to accurately represent the proportion of species in a microfossil assemblage at a confidence interval of 0.95 (Phleger, 1960; Shaw, 1964; Dennison and Hay, 1967;

Patterson and Fishbein, 1989; Buzas, 1990; Revets, 2004). Counts were conducted on unsieved sample slides. Dinocysts were enumerated and identified to the lowest taxonomic level possible alongside the minimum count of spores and pollen to calculate a terrestrial-to-marine ratio.

Photomicrographs were procured through use of Stream Motion software and an Olympus DP72 camera.

Multivariate Statistical Techniques

Relative abundance of terrestrial palynomorphs preserved in marine sediment has been shown to correlate well with terrestrial vegetation on the nearby continent despite taphonomic

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uncertainties (Muller, 1959; Heusser and Balsam, 1977; Mudie, 1982; Heusser, 1983;

Hooghiemstra et al., 1986; Mudie and McCarthy, 1994; Sun et al., 1999; van der Kars, 2001; van der Kaars and De Deckker, 2003; Hooghiemstra et al., 2006; Montade et al., 2011; Luo et al.,

2014; Zhao et al., 2016). Relative abundance of spores and pollen is therefore assumed to closely reflect the paleoenvironment from which they originated. Compositional data of terrestrial palynomorphs is herein used as a basis for the multivariate statistical techniques in this study.

Q- and R-mode hierarchical cluster trees were generated using Ward’s method using the computer program RStudio. Q-mode cluster analysis was used to create groups of samples that have similar palynological content while R-mode cluster analysis was used to group taxa that tend to co-occur in a sample. Ward’s method is a type of agglomerative clustering in multivariate

Euclidean space, wherein the distance between clusters is based on the increase in sum-of-squares

(variance) in the merged cluster (Everitt et al., 2011). Ward’s method was created to minimize the increase in total within-cluster variance that typically results when forming clusters for data intepretation (Ward, 1963; Murtagh and Legendre, 2014; de Amorim, 2015). The algorithm starts by having single data points representing their own cluster and then merging the pair that minimizes the increase in the total within-cluster variance; this is then repeated until only one cluster remains (Wilks, 2011; Murtagh and Legendre, 2014). Hierarchical cluster analyses using

Ward’s method therefore separate variables into groups based on minimum variance and merging cost, the relationship of which can be visualized through dendrograms (Ward, 1963; Wilks, 2011).

Ward’s method was used in this study as it takes into account the degree of similarity with respect to multiple variables (Ward, 1963) and can be represented in Euclidean space unlike some other agglomerative clustering methods (Everitt et al., 2011). This Euclidean space is the same reference

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space in other multivariate ordination methods such as principal component analysis (PCA) or detrended correspondence analysis (DCA) (Murtagh and Legendre, 2014).

Stratigraphically constrained incremental sum-of-squares (i.e., Ward’s method) cluster analysis (CONISS) using square root data transformation (to up-weigh rare types) was used to show trends through time in the relative abundance of pollen and spores in the stratigraphic section.

In comparison to unconstrained cluster analysis mentioned above, CONISS forms clusters by analyzing stratigraphically adjacent samples (Grimm, 1987). The resulting dendogram can be delineated at any height to define clusters and, consequently, stratigraphic zones (Grimm, 1987).

The palynological signature within these stratigraphic zones can then be used to infer paleoenvironmental conditions over time.

The computer program RStudio was used in conjunction with the DECORANA function in the R package VEGAN to observe and interpret site and species ordination in DCA. DCA is a multivariate statistical technique used for large datasets to help interpret major environmental factors, expressed through eigenvalues, which may have affected ecological communities (Hill and Gauch, 1980). Real world ecological data has a heterogeneous nature that causes many multivariate methods such as PCA to produce an “arch” and/or “edge” effect (Hill and Gauch,

1980; Jackson and Somers, 1991). The “arch” effect occurs when the second axis is curved and distorted relative to the first axis, presenting datasets as a curve rather than as a straight line (Hill and Gauch, 1980). This limits interpretation of possible environmental gradients represented by the second axis. The “edge” effect is produced in reciprocal averaging (RA) because compositional distances are not preserved, so pairs of samples with equal compositional differences will appear to have a greater distance between them in the middle of the gradient compared to at the edges

(Hill and Gauch, 1980). To circumvent these problems, DCA was created. DCA uses RA to

142 eliminate the “arch” effect by separating the first axis into sections and detrending the second axis so that it is equalized to a zero-mean value (Hill and Gauch, 1980). The “edge” effect is eliminated by rescaling the axes under the assumption that species have a non-monotonic and unimodal relationship with their environment (Jackson and Somers, 1991; Paliy and Shankar, 2016). DCA is therefore the preferred tool for interpreting environmental gradients and has been used previously in palynological datasets to interpret Cretaceous paleoenvironmental conditions

(Stukins et al., 2013; Galloway et al., 2015).

RESULTS

Dinoflagellate Cysts

Dinocysts are poorly preserved in the Treeless Creek section and are occasionally absent from samples (Fig. 5). Due to their poor preservation and relatively low abundance, species-level identification was difficult to determine (Table 1). Dinocysts have an average count of 32 ± 38 SD

(n=16) for every minimum spore and pollen count of 300 in unsieved sample preparations. The marine-to-terrestrial ratio based on spores, pollen grains, and dinocysts varied between a minimum of zero and a maximum of 69:172 (40.1% dinocysts) at 75 m. Dinocysts show a general trend of increasing relative abundance proportional to spores and pollen towards the top of the measured mudstone section (unit 2) at 75 m. Dinocysts were generally found in higher abundances (1-40%) in mudstone samples in comparison to those collected from the sandstone-dominant strata (0-6%).

No dinocysts were preserved in the basal sandstone-dominant unit (unit 1), and only two species were identified in the overlying sandstone-dominant unit (unit 3): Oligosphaeridium cf. tenuiprocessum and Paragonyaulacysta? borealis.

The following dinocyst taxa reach peak relative abundance within the mudstone-dominant succession (unit 2): Epiplosphaera cf. saturnalis (22.22% at 55 m), Gonyaulacysta dualis (9.42%

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Figure 5: Marine-to-terrestrial ratio of the Treeless Creek section juxtaposed against relative abundance of dinocysts identified to the species level. Marine-to-terrestial ratio is based on the percentage of dinocysts relative to a minimum count of 300 of spores and pollen.

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Table 1: Biological nomenclature and taxonomic authority of identified dinoflagellate cysts

(Division Dinoflagellata) in the Treeless Creek section.

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Family Subfamily Genus/Species Authority Gonyaulacaceae Gonyaulacoideae Spiniferites spp. Mantell 1850 Gonyaulacaceae incertae sedis Cometodinium spp. Deflandre and Courteville 1939 Hystrichodinium spp. Deflandre 1935 incertae sedis Epiplosphaera cf. saturnalis (Brideaux and Fisher 1976) Dodekova 1994 Gonyaulacaceae Leptodinioideae Gonyaulacysta dualis (Brideaux and Fisher 1976) Stover and Evitt 1978 Gonyaulacysta spp. Deflandre 1964 Stiphrosphaeridium anthophorum (Cookson and Eisenack 1958) Lentin and Williams 1985 Oligosphaeridium pulcherrimum (Deflandre and Cookson 1955) Davey and Williams 1966 Oligosphaeridium cf. tenuiprocessum Singh 1983

Pareodiniaceae Pareodinioideae Paragonyaulacysta? borealis (Brideaux and Fisher 1976) Stover and Evitt 1978 Pareodinia ceratophora Deflandre 1947 Ceratiaceae Muderongia tetracantha Gocht 1957

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at 75 m), G.? pectinigera (0.72% at 75 m), Muderongia tetracantha (29.41% at 51 m),

Stiphrosphaeridium anthophorum (17.65% at 51 m), Oligosphaeridium cf. tenuiprocessum

(55.32% at 63 m), and Paragonyaulacysta? borealis (18.18% at 47 m). Oligosphaeridium cf. tenuiprocessum is the most abundant species found in the Treeless Creek section in terms of relative abundance (21.68% ± 16.34 SD, n=16).

Oligosphaeridium cf. tenuiprocessum specimens all exhibit the long processes, narrow stems, and recurved spines that characterize O. tenuiprocessum (Singh, 1983), although the process endings on the observed specimens in the Treeless Creek section are more complex than the holotype (R.A. Fensome, pers. comm., 2018).

Identification of Paragonyaulacysta? borealis is uncertain as the presence of paratabulation is unclear. The possible absence of paratabulation suggests an alternative identification of Pareodinia ceratophora, which is a common species in the Canadian Arctic, known to be Aalenian to late Callovian or early Oxfordian in age (Powell, 1992; Stover et al.,

1996). Correct identification of Pareodinia ceratophora is uncertain due to an incomplete description of its holotype that lacks a known archeopyle location, which is a diagnostic feature used to identify dinocysts to the species level (R.A. Fensome, pers. comm., 2018).

Spores and Pollen

Spores and pollen grains at Treeless Creek were enumerated to a minimum of 300 specimens per unsieved sample preparation (316 ± 27 SD, n=16) with the exception of sample C-

604498, which had a count of 227. Forty-five different taxa were identified to the genus or species level (Table 2). Specimens were poorly- to moderately-well preserved. No angiosperm pollen were observed.

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Table 2: Biological nomenclature and taxonomic authority of identified spores and pollen grains from the Treeless Creek section.

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Division Class Order Family Genus/Species Authority Marchantiophyta Marchantiopsida Sphaerocarpales Sphaerocarpaceae Aequitriradites spp. (Delcourt and Sprumont 1955) Cookson and Dettmann 1961 Bryophyta Sphagnopsida Sphagnales Sphagnaceae Cingulatisporites spp. Thomson and Pflug 1953 Cingutriletes clavus (Balme) Dettmann 1971 Stereisporites (Wilson and antiquasporites Webster) Dettmann Tracheophyta Lycopodiopsida Lycopodiales Lycopodiaceae Camarozonotriletes Norris 1967 insignis Coronatispora (Couper) Dettmann valdensis 1963 Lycopodiumsporites Hedlund 1966 crassimacerius Lycopodiumsporites Singh 1971 expansus Retitriletes (Cookson 1953) austroclavatidites Doring et al. in Krutzsch 1963 Sestrosporites (Couper) Dettmann pseudoalveolatus 1963 incertae sedis Leptolepidites Couper 1953 verrucatus Selaginellales Selaginellaceae Densoisporites spp. Dettmann 1963 Neoraistrickia (Cookson) Potonié truncata 1956 incertae sedis Lycospora spp. Schopf et al. 1944 Polypodiopsida Cyatheales Cyatheaceae Kuylisporites lunaris Cookson and Dettmann 1958 Dicksoniaceae, Cyathidites australis Couper 1953 Cyatheaceae Deltoidospora hallii Miner 1935 Gleicheniales Gleicheniaceae Gleicheniidites Ross 1949 senonicus Matoniaceae Dictyophyllidites Couper 1958 harrisii Matonisporites (Balme 1957) crassiangulatus Dettmann 1963 Osmundales Osmundaceae Baculatisporites (Cookson 1953) comaumensis Potonie 1956 Biretisporites Delcourt and potoniaei Sprumont 1955 Osmundacidites Couper 1953 wellmanii Todisporites (Maljavkina) rotundiformis Pocock 1970 Verrucosisporites spp. Potonié and Kremp 1954 Schizaeales Schizaeaceae Distaltriangulisporites Singh 1971 spp. Klukisporites spp. Couper 1958 Ruffordiaspora cf. (Cookson 1953) australiensis Dettmann and Clifford 1992 incertae sedis Apiculatisporites spp. Potonié and Kremp 1956 Leiotriletes directus Balme and Hennelly 1956 Cycadopsida, Cycadales, incertae sedis Entylissa spp. Naumova 1937 Ginkgoopsida Ginkgoales Cycadopites spp. Wodehouse 1933

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Cycadopsida Cycadales incertae sedis Chasmatosporites (Nilsson 1958) spp. Pocock and Jansonius 1968 Pinopsida Pinales Araucariaceae Araucariacites Cookson 1947 australis Cheirolepidiaceae Classopollis (Pflug) Pocock and classoides Jansonius 1961 Cupressaceae Perinopollenites (Couper) Dettmann elatoides 1963 Cupressaceae, Undifferentiated Taxaceae Pinaceae Cerebropollenites (Couper) Nilsson mesozoicus 1958 Laricoidites magnus (Potonié) Potonié, Thomson and Thiergart 1950 Undifferentiated bisaccate pollen Gnetopsida incertae sedis incertae sedis Eucommiidites Erdtman 1948 troedsonii incertae sedis Acanthotriletes Pocock 1962 varispinosus Granulatisporites spp. Potonié and Kremp 1954 Psilatriletes radiatus Brenner 1963 Triquitrites spp. (Wilson and Coe 1940) Potonié and Kremp 1954

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Unknown “palynomorph A” (Plate 19)

Cluster Analyses: Q-, R-mode and CONISS

Q-mode cluster analysis is used to delineate three groups within the stratigraphic section

(Fig. 6). Cluster 1 includes two samples collected from below the mudstone succession and two samples within the upper beds of the mudstone succession (unit 2). Cluster 2 groups four samples collected from within the mudstone succession and one sample above. Cluster 3 contains the largest portion of consecutive samples within the mudstone succession and a few samples collected from above and below the mudstone unit. The clusters do not follow any definitive trends that correspond to distinct heights in the stratigraphy.

R-mode cluster analysis is used to delineate four informal clusters: 1) “spore assemblage”,

2) “spore-pollen assemblage”, 3) “Osmundacidites-Densoisporites (O-D) assemblage” and 4)

“bisaccate-CT pollen” assemblage (Fig. 7). The spore assemblage is the smallest assemblage in terms of relative abundance but the largest in terms of diversity. The spore assemblage constitutes

6% of the total palynomorph count. The assemblage is mainly composed of a mix of spores derived from the classes Sphagnopsida (mosses), Lycopodiopsida (fern allies), and Polypodiopsida (ferns), with some pollen from Cycadopsida (cycads), Gnetopsida (gnetophytes) and Pinopsida (conifers).

On average, the most abundant species in this assemblage are Perinopollenites elatoides pollen

(0.53% ± 0.49 SD %, n=16) and Deltoidospora hallii spores (0.45% ± 0.62 SD %, n=16).

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Figure 6: Q-mode cluster analysis of the Treeless Creek section showing three clusters with no apparent stratigraphic significance. Meterages 95 and 97 are used to designate those samples found in unit 3 (C-604482-83).

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Figure 7: CONISS combined with R-mode cluster analysis showing R-mode clusters (Spore,

Spore-Pollen, Osmundacidites-Densoisporites (O-D) spore, Bisaccate – Cupressaceae-Taxaceae

(CT) pollen assemblages) and juxtaposed against percentage of dinocysts per minimum 300 counts of pollen and spores, relative abundance of identified terrestrial palynomorphs, and stratigraphically constrained clusters (A, B, C). Grab samples are assigned arbitrary heights to view stratigraphically constrained palynological trends.

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In comparison to the spore assemblage, the spore-pollen assemblage contains less species from the spore-producing plants and more from the pollen-producing plants. The palynomorphs in this assemblage are derived from plants from the classes Lycopodiopsida, Polypodiopsida,

Sphagnopsida, Marchantiopsida (liverworts), and Pinopsida. This assemblage constitutes 22% of the total palynomorph count. The most abundant species from this assemblage, on average, are

Laricoidites magnus (4.31% ± 3.29 SD %, n=16) and Cerebropollenites mesozoicus pollen (2.85%

± 1.50 SD %, n=16).

The O-D assemblage contains only two species: Osmundacidites wellmannii (7.91% ± 3.64

SD %, n=16) and Densoisporites spp. (5.65% ± 3.29 SD %, n=16) spores. Together, they constitute

24% of the total palynomorph count.

The bisaccate-CT pollen assemblage is the largest group in terms of relative abundance, constituting 48% of the total palynomorph count. The assemblage only contains two taxa: undifferentiated Pinaceae-type bisaccate (21.39% ± 8.94 SD %, n=16) and Cupressaceae-Taxaceae

(CT) pollen (27.53% ± 6.64 SD %, n=16). Bisaccate and CT pollen are both prevalent throughout the studied section.

Clusters delineated from R-mode cluster analysis are stratigraphically constrained using

CONISS to interpret palynological trends through time. CONISS reveals three informal clusters within the studied section (Fig. 7). Cluster C groups the samples from the underlying sandstone- dominant unit (unit 1). This cluster is characterized by a maximum in Laricoidites magnus and bisaccate pollen: Laricoidites magnus pollen peaks to a relative abundance of 14.51% while bisaccate pollen peaks to a relative abundance of 40.06%. Osmundacidites spores decrease to a minimum relative abundance of 2.52% in this cluster. Cluster B spans the lower half of the mudstone unit (unit 2). Palynological trends reveal an increase in the spores Densoisporites spp.,

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Osmundacidites wellmannii, and multiple other spore species derived mainly from ferns.

Simultaneously, bisaccate and CT pollen show a decrease in relative abundance, with CT pollen reaching a minimum of 17.85% at 51 m. Cluster A is another mixed assemblage that is characterized by the appearance of several species that were previously absent, such as:

Apiculatisporites spp., Coronatisporites valdensis, Distaltriangularisporites spp., and

Klukisporites spp. spores. Towards the top of Cluster A, CT, Cerebropollenites mesozoicus, and

Classopollis classoides pollen reach their respective maximum of 39.40%, 6.00%, and 6.62% in terms of relative abundance.

Detrended Correspondence Analysis

Detrended correspondence analysis was generated using the DECORANA function in the

R package VEGAN. The resulting DCA plot reveals two axis labelled DCA1, which has an eigenvalue of 0.09, and DCA2, which has an eigenvalue of 0.04 (Fig. 8). Araucariacites australis and Pinaceae pollen plot near the extreme negative values on DCA1; Aequitriradites spp.,

Schizaceae, and unclassified spores from the Class Polypodiopsida plot near the extreme positive values. Classopollis classoides, Chasmatosporites spp., and Araucariacites australis pollen plot near the extreme negative values on DCA2. Leptolepidites verrucatus spores and Eucommiidites troedsonii pollen plot near the extreme positive values on DCA2.

DISCUSSION

Biostratigraphy

Dinocyst Zonations

There exists four proposed dinocyst zonations for the Cretaceous in Arctic Canada

(McIntyre, 1974; Doerenkamp et al., 1976; Bujak and Scott, 1984; Andrews, 2012). Proposed palynological zonation schemes for the Cretaceous are typically missing upper Albian to lower

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Figure 8: DCA plot of Treeless Creek section with axes representing temperature (DCA1) and habitat (DCA2) along a typical coastal wetland profile (modified from Environment Canada and

Wilcox, 2002; Wilcox et al., 2007).

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Santonian data due to a sub-Santonian unconformity present throughout the Mackenzie Delta,

Banks Island, and Anderson Plain areas. Comparisons of late Albian-early Cenomanian dinocyst assemblages are therefore limited within large areas of the Canadian Arctic. The first zonation scheme was created by McIntyre (1974) to ascertain the ages of an Upper Cretaceous section from

Horton River, which had three microfloral divisions (H1-H3) that were Santonian-Maastrichtian in age. The Horton River dinocyst assemblage did not share any commonalities with the Treeless

Creek section.

The second zonation scheme was created by Doerenkamp et al. (1976) based on the palynology from Banks, Eglinton, Prince Patrick and Mackenzie King Islands and the Anderson

Plain. Palynology from Cretaceous and Paleogene rocks revealed a series of zones labelled CI-III and CV-VII for Cretaceous zones and TI-TII for Paleogene zones. The C1 Zone, Trilobosporites-

Classopollis zone, is characterized by the presence of Trilobosporites Potonié,

Concavissimisporites Delcourt and Sprumont, Cicatricosisporites Potonié and Gelletich, and

Appendicisporites Weyland and Greifeld spores and Cerebropollenites pollen. The C1 Zone ranges from the upper Valanginian(?) to Aptian. The CII Zone, Gardodinium trabeculosum zone, is divided into three subzones: CIIa, CIIb, CIIc. CIIa is characterized by the rare appearances of

Pseudoceratium cf. pelliferum Gocht, Odontochitina operculata, Batioladinium jaegeri,

Stiphrosphaeridium anthophorum, Oligosphaeridium pulcherrimum, Palaeoperidinium cretacea and Circulodinium distinctum (Deflandre and Cookson) Jansonius. CIIb is characterized by the disappearance of Pseudoceratium cf. pelliferum and an increase in abundance in Batioladinium micropodum (Eisenack and Cookson) Brideaux, Apteodinium maculatum Eisenack and Cookson, and Oligosphaeridium complex. CIIc is characterized by an overall increase in dinocyst abundance and diversity. In terms of terrestrial palynomorphs, the CII Zone contains an abundance of

161 bisaccate and Classopollis classoides pollen. The CII Zone corresponds to the Christopher

Formation (upper Aptian – middle Albian) on Banks Island, the Crossley Lakes Member (upper

Aptian – middle Albian) and the Horton River Formation (lower – upper Albian) in the Horton-

Anderson River area (Brideaux and McIntyre, 1975; Doerenkamp et al., 1976; Nøhr-Hansen and

McIntyre, 1998).

The CIII Zone, ?Palaeostomocystis sp. Deflandre, Apteodinium sp. aff. reticulatum Singh zone, is characterized by the appearance of ?Palaeostomocystis sp., Fromea fragilis (Cookson and

Eisenack) Stover and Evitt and Apteodinium sp. aff. reticulatum, and the disappearance of various dinocysts including Gardodinium trabeculosum, Impagidinium verrucosum (Brideaux and

McIntyre) Stover and Evitt, Pterodinium aliferum Eisenack, Trichodinium sp. aff. spinosum Singh, and Ellipsoidictyum imperfectum (Brideaux and McIntyre) Lentin and Williams. In terms of terrestrial microflora, the CIII Zone is characterized by the rare appearance of tricolpate pollen and the disappearance of Concavissimisporites, Pilosisporites Delcourt and Sprumont, and

Trilobosporites spores. The CIII Zone corresponds to the Hassel Formation on Ellef Ringnes and

Banks Island. The CV-VII zones range Santonian or earlier in age. Comparisons of the dinocyst assemblage at Treeless Creek with the dinocyst zonation scheme of Doerenkamp et al. (1976) reveals a similarity with CIIa Zone (upper Aptian), where Stiphrosphaeridium anthophorum and

Oligosphaeridium pulcherrimum both appear. Stiphrosphaeridium anthophorum and

Oligosphaeridium pulcherrimum are species that commonly occur in upper Lower Cretaceous strata globally (Brideaux and McIntyre, 1975), but are difficult to use in biostratigraphy due to their relatively long age ranges (Fig. 9).

The palynological zonation scheme proposed by Bujak and Scott (1984) examined palynological data compiled on Cretaceous and Cenozoic rocks of the Beaufort-Mackenzie region.

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Figure 9: First and last occurrences of identified dinocyst species observed in the Treeless Creek section. See text for age range references.

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The Berriasian to Cenomanian zones are: Tetrachacysta spinosigibberosa (Brideaux and Fisher)

Backhouse (lower to upper Berriasian), Gochteodinia judilentiniae McIntyre and Brideaux (lower to middle Valanginian), Gochteodinia villosa (Vozzhennikova) Norris (Hauterivian), Muderongia asymmetrica (Aptian to lower Albian), Gardodinium trabeculosum (middle Albian), and

Isabelidinium magnum (Davey) Stover and Evitt (upper Cenomanian to Turonian) zones. The CIIa

Zone of Doerenkamp et al. (1976) correlates with the Muderongia asymmetrica Zone. However, none of the diagnostic dinocysts observed in Bujak and Scott’s (1984) M. asymmetrica Zone are observed in the Treeless Creek study. Morever, the dinocyst assemblages reported by Bujak and

Scott (1984) are entirely dissimilar to those found in the Treeless Creek section.

The most recent dinocyst zonation scheme was proposed by Andrews (2012) based on the

Kanguk Formation on Ellef Ringnes Island, Sverdrup Basin. This Upper Cretaceous zonation scheme was created by collecting data from various studies throughout the Arctic (Århus, 1991;

McIntyre in Dixon, 1995; Nøhr-Hansen, 1993, 1996; Núñez-Betelu, 1994; MacRae, 1996). The dinocysts observed in Andrews’ (2012) dissertation were absent in the Treeless Creek section.

Age Estimation

The Treeless Creek area mainly depends on microfossil assemblages for chronostratigraphic purposes due to the paucity of age-diagnostic macrofossils. The area is also difficult to map due to the prevalence of faults and the presence of multiple unconformities.

Dinocysts identified to the species level in the Treeless Creek section can be divided into three groups with relatively distinct ages: Late Jurassic, Valanginian-Barremian, and Albian-

Cenomanian (Fig. 9). The Upper Jurassic dinocysts are: Epiplosphaera cf. saturnalis (Brideaux and Fisher, 1976), Gonyaulacysta dualis (Brideaux and Fisher, 1976), and Gonyaulacysta? pectinigera (Powell, 1992; Wiggan et al., 2017). Muderongia tetracantha has an age range

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exclusive to the Valanginian-Barremian (Powell, 1992). Stiphrosphaeridium anthophorum has an age range that extends from the Hauterivian/Barremian to the Albian (Powell, 1992; Stover et al.,

1996). Oligosphaeridium cf. tenuiprocessum has the youngest age range of the identified dinocysts, and appears between 45 m and the top of the section. The age range of O. cf.tenuiprocessum is considered as latest Albian – early Cenomanian (Singh, 1983), and has been found in the Fish Scale Marker Bed, which marks the Albian-Cenomanian boundary in northern

Alberta (Leckie et al., 1992). Due to their larger age ranges, Oligosphaeridium pulcherrimum (Cox et al., 1987; Al-Ameri and Batten, 1997) and Paragonyaulacysta? borealis (Brideaux, 1977) overlap with the other age groups. Based on the age ranges of these dinocysts, the section at

Treeless Creek section is late Albian in age and may extend into the Cenomanian. Reworking of dinocysts from Jurassic and Lower Cretaceous sediment into younger sediment likely resulted in the presence of dinocysts with an older age range in the Treeless Creek section.

Dinocyst species identified in the Treeless Creek section can also be found in the Upper

Cretaceous Kanguk Formation in Glacier Fiord, Axel Heiberg Island, and Slidre Fiord, Ellesmere

Island, Sverdrup Basin (Núñez-Betelu, 1991, 1994; Froude, 2018). These include:

Stiphrosphaeridium anthophorum and Oligosphaeridium pulcherrimum. In the Horton-Anderson

Plains area, Stiphrosphaeridium anthophorum and Oligosphaeridium pulcherrimum are restricted to the Crossley Lakes Member, which is late Aptian to middle Albian in age (Brideaux and

McIntyre, 1975). In the Hume River section in the Northwest Territories, Oligosphaeridium pulcherrimum and Stiphrosphaeridium anthophorum occur within the Slater River Formation

(Fensome, 2016). Stiphrosphaeridium anthophorum is also present in the Arctic Red Formation in the same area (Fensome, 2016). The Arctic Red Formation is early to middle Albian in age (Dixon,

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1992; Fensome, 2016) whereas the Slater River Formation is late Albian to Cenomanian age

(Fensome, 2016).

Comparisons between the dinocyst assemblages and age estimations of the formations in the Treeless Creek area are summarized in Table 3.

GSC locality 88085 lies approximately 2 km north of the Treeless Creek section (GSC

Report no. Km-7-1974-JAJ; Fig. 1, mapped as Husky Formation). The age estimation of this section is based on the presence of Buchia concentrica Sowerby, 1827, which is late Oxfordian in age (Jeletzky, 1967). The age estimations on the original map may have been mistakenly labelled as Kimmeridgian-Portlandian. It is important to note that the recorded latitude and longitude for

GSC locality 88085 were rounded to the nearest minute, and the section may have originally been located approximately 4 km to the south (T. Poulton and L. Lane, pers. comm., 2018).

Approximately 16 km to the southeast of the Treeless Creek section is the grab sample C-

604499 (Appendix 2), collected from the Rat River Formation on the banks of Rat River at

67.7407°N, 135.45938°W. The age estimation based on the dinocyst assemblage in this sample is inconsistent with the currently accepted age range for the Rat River Formation. The dinocyst assemblage includes the species Kiokansium williamsii, which has a first appearance datum in the middle-late Albian and a last appearance datum in the middle Cenomanian (Stover et al., 1996;

Fensome et al., 2009; Fensome; 2016), suggesting that the age of the Rat River Formation may extend into the middle Albian. This interpretation is supported by the presence of K. cf. unituberculatum, which has an age extending from the Barremian-Cenomanian (Stover et al.,

1996). In the Horton-Anderson Plains area, Kiokansium unituberculatum is also present and is restricted to the Horton River Formation, which is early to late Albian in age (Brideaux and

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Table 3: Paleontological reports and age estimations based on fossil assemblages collected from various localities in the Treeless Creek area (see Fig. 1).

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Location Formation Accepted Fossils Age according to age for estimation Norris’ (1981) formation based on map fossils GSC loc. Martin Creek Berriasian Lunatadinium dissolutum Brideaux and McIntyre Barremian- C-27094 Formation Muderongia mcwhaei Cookson and Eisenack Aptian M. tetracantha (Brideaux, Odontochitina nuda (Gocht) Dörhöfer and Davies 1974) Oligosphaeridium albertense (Pocock) Davey and Williams Oligosphaeridium? asterigerum (Gocht) Davey and Williams O. complex (White) Davey and Williams Spiniferites ramosus (Ehrenberg) Mantell Tenua hystrix Eisenack Wrevittia helicoidea (Eisenack and Cookson) Sargeant

GSC loc. Husky Formation Oxfordian- Buchia concentrica Sowerby, 1827 late 88085 early Oxfordian Berriasian

GSC section Rat River late Batioladinium longicornutum (Alberti) Brideaux Barremian 72-WB-12 Formation Barremian- B. micropodum (Eisenack and Cookson) Brideaux (McIntyre, Aptian Callaiosphaeridium asymmetricum (Deflandre and 1990) Courteville) Davey and Williams Dingodinium cerviculum (Cookson and Eisenack) Mehrotra and Sarjeant Fromea amphora Cookson and Eisenack Gardodinium trabeculosum (Gocht) Alberti Heslertonia heslertonensis (Neale and Sarjeant) Sarjeant Kleithriasphaeridium cooksoniae (Singh) Fensome Muderongia asymmetrica Brideaux Odontochitina nuda O. operculata (Wetzel) Deflandre and Cookson Oligosphaeridium albertense (Pocock) Davey and Williams O. complex O. pulcherrimum Pseudoceratium gochtii Neale and Sarjeant P. retusum Brideaux Stiphrophaeridium anthophorum Subtilisphaera perlucida (Alberti) Jain and Millepied Tubotuberella uncinata (Brideaux) Davies Wallodinium luna (Cookson and Eisenack) Lentin and Williams

GSC loc. Rat River late Batioladinium jaegeri (Alberti) Brideaux Aptian- C-27085 Formation Barremian- Cribroperidinium? muderongense (Cookson and Eisenack) Albian Aptian Davey (Brideaux, Gardodinium trabeculosum 1974) Hystrichosphaeridum irregulare (Pocock) Davey and Williams Impletosphaeridium? clavulum (Davey) Islam Kiokansium unituberculatum Leptodinium? hyalodermopse (Cookson and Eisenack) Stover and Evitt Lunatidinium dissolutum Odontochitian operculata Oligosphaeridium complex Palaeoperidinium cretaceum (Pocock ex Davey) Lentin and Williams Subtilisphaera perlucida (Alberti) Jain and Millepied

Sample Rat River late Epiplosphaera cf. saturnalis middle C-604499 Formation Barremian- Kiokansium cf. unituberculatum (Tasch) Stover and Evitt Albian – Aptian K. williamsii Singh middle Oligosphaeridium pulcherrimum Cenomanian Sirmiodinium grossii Alberti (see text; Appendix 2)

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McIntyre, 1975; Dixon, 1992). However, K. unituberculatum may have a modified age range extending from the late Berriasian to the late Cenomanian (Fensome, 2016), which limits its use for biostratigraphic correlation. The paleontological report for GSC locality C-27085 and the dinocyst assemblage for C-604499 suggest that the age range for the Rat River Formation may extend into the lower Albian.

According to Norris’ (1981) map, the outcrops shared between GSC locality 88085, C-

27094 and this study’s section are mapped as the Husky Formation (Oxfordian-Berriasian), Martin

Creek Formation (Berriasian), and Mount Goodenough Formation (Barremian/Hauterivian –

Aptian) respectively. Comparisons between the dinocyst assemblages across Arctic Canada and from the Treeless Creek area (Table 3) suggest that the outcrops at Treeless Creek have a younger age than originally postulated. The outcrops currently labelled as the Martin Creek Formation (C-

27094) may have a Barremian-Aptian age which is more consistent with these strata being assigned to the Rat River Formation. Overyling this formation would be upper Albian strata of the

Treeless Creek section in this study. Currently, there are no known upper Albian-lower

Cenomanian formations in the Aklavik Range due to a regional unconformity (Dixon, 1993).

Jeletzky (1960) described an “Albian shale-siltstone division” that occurs throughout the eastern Richardson Mountains between Willow River and Stony Creek (Norris, 1981). The presence of Beudanticeras affine Whiteaves, 1892, Lemuroceras(?) sp. indet. Casey, 1954 and

Cleoniceras(?) sp. indet. Parona and Bonnarelli, 1895 within the lower beds of the division suggest a middle Albian age (Jeletzky, 1960). The upper beds of the Christopher Formation and the lower beds of the Hassel Formation in the Canadian Arctic Archipelago are thought to be equivalent to the Albian shale-siltstone division (Heywood, 1957; Jeletzky, 1960). Mountjoy and Chamney

(1969) recognized that this division correlated with the Arctic Red Formation identified in the Peel

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Plateau. Norris’ (1981) map reflects this correlation by plotting the “Albian shale-siltstone division” as the Arctic Red Formation. Dixon (1992, 1993) likewise referred to the “Albian shale- siltstone division” as the Arctic Red Formation for simplicity. A formal name for the “Albian shale-silstone division” does not yet exist. The “Albian shale-siltstone division” is described as consisting of predominantly mudstones and siltstones with a prevalence of interbedded ironstone concretions (Jeletzky, 1960), which is similar to the main lithology of the Arctic Red Formation

(Mountjoy and Chamney, 1969; Dixon, 1992). The lower beds of the “Albian shale-siltstone division” are composed of siltstone and sandstone, with an approximate 2 m thick bed of “pebbly shale” interbedded with “pebbly grit” near the base (Jeletzky, 1960). This lithology is consistent with observations at Treeless Creek.

The microfossil assemblages throughout the Treeless Creek area suggest that exposures of

J/K boundary strata (Husky Formation) are unlikely to exist in the immediate area north of Mount

Lang, with the exception of GSC section DFA-82-10 (Figs. 1, 2). At this location, the Husky

Formation appears at the surface extensively, with the contact between the arenaceous and red- weathering members clearly visible (Dixon, 1992; Fig. 2). The Husky Formation may be largely missing in the Treeless Creek area as a result of its proximity to the Tuktoyaktuk Fault Flexure

Zone (Fig. 1) or because of the sub-Hauterivian unconformity. The unconformity at the base of the Mount Goodenough Formation may have removed the Husky and Martin Creek formations entirely. Adjacent strata to the Bug Creek Group could therefore belong to the Mount Goodenough,

Rat River, and Arctic Red formations (or upper Albian equivalents) in place of Husky, Martin

Creek, Mount Goodenough formations, respectively (Fig. 10).

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Figure 10: Possible revision of the geology in the Treeless Creek area (modified from Norris,

1981). Mount Goodenough Formation replaces Husky Formation. Rat River Formation replaces

Martin Creek Formation. Arctic Red Formation (or equivalent) replaces Mount Goodenough. See

Fig. 1 for comparison.

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Paleoecology

Bryophytes, Lycophytes and Monilophytes

Bryophyte spores in the Treeless Creek section include Sphaerocarpaceae spores, which were derived from liverworts (Marchantiophyta), and Sphagnaceae spores, derived from mosses

(Bryophyta). Extant liverworts predominantly thrive in moist, shaded areas; some species are aquatic (Raven et al., 2013). Mosses typically thrive under moist conditions, although they can also withstand droughts in arid environments (Raven et al., 2013). Some extant mosses are also dominant in areas within the Arctic Circle, can withstand extreme cold temperatures in the

Antarctic, and can grow on mountainous slopes above the tree line (Raven et al., 2013). Some species are aquatic, although none are entirely marine (Raven et al., 2013). Sphagnaceae spores are derived from the peat mosses (Sphagnopsida), which are an early-diverging lineage of mosses

(Shaw et al., 2010). Sphagnaceae has a single genus, Sphagnum L. Sphagnum species are xerophytic hydrophytes, growing globally in wetlands such as bogs while being capable of surviving periodic droughts (Andrus, 1986; Raven et al., 2013).

In the Treeless Creek section, two known lycophyte (i.e., fern allies) families are present:

Lycopodiaceae and Selaginellaceae. Like bryophytes, Lycopodiaceae ferns have a diverse range of habitats that limits their use in interpreting paleoenvironmental conditions, although they most often thrive in tropical areas with relatively high humidity. However, some species are capable of thriving in temperate or even polar habitats, although they are rarely the prominent plant in these communities (Raven et al., 2013; Galloway et al., 2015). Selaginellaceae spores in the Treeless

Creek section include Densoisporites spp. and Neoraistrickia truncata, but only Densoisporites spores can be used for ecological interpretations. Densoisporites spores are closely associated with the Pleuromiaceae family, which is related to the “living fossil” Isoetes L. (Raine et al., 1988;

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Abbink et al., 2004). Mesozoic Isoetes were aquatic plants that gradually shifted to more terrestrial, xeric environments following the Cretaceous (Taylor and Hickey, 1992). Some

Pleuromiaceae plants may have been salt-tolerant and well-adapted for coastal environments, although most occur in inland basins where marine influence was minimal (Retallack, 1997 and references therein).

Monilophyte (i.e., fern) spores in the Treeless Creek section include those from the families

Cyatheaceae, Dicksoniaceae, Gleicheniaceae, Matoniaceae, Osmundaceae, and Schizaceae. The majority of extant ferns thrive under warm and humid conditions that are characteristic in tropical and subtropical habitats. Lowland environments such as swamps, marshes, forest understories, and riverbanks are ideal habitats for most ferns (van Konijnenburg-van Cittert, 2002). Those ferns that typically grow near river banks include Cyatheaceae, Dicksoniaceae, Osmundaceae, and

Schizaceae (Abbink et al., 2004). Both extant and Mesozoic Cyatheaceae ferns prefer tropical and subtropical environments (van Konijnenburg-van Cittert, 2002). Cyatheaceae ferns are often grouped with Dicksoniaceae ferns, which are limited to tropical and temperate rainforests in the

Southern Hemisphere (van Konijnenburg-van Cittert, 2002). Mesozoic Dicksoniaceae ferns are likewise considered to have mainly thrived in warm, moist habitats (van Konijnenburg-van Cittert,

2002 and references therein). Mesozoic Osmundaceae ferns are known to occur in floodplain deposits (van Konijnenburg-van Cittert, 2002). Gleicheniaceae and Schizaceae ferns can tolerate full sunlight and grow in open habitats (Abbink et al., 2004; Galloway et al., 2015). Matoniaceae ferns are limited to the mountain ranges of the Malaysian Archipelago, although Matoniaceae ferns were distributed worldwide in the Mesozoic (Stewart and Rothwell, 1993; van Konijnenburg-van

Cittert, 1993; Abbink et al., 2004). Matoniaceae ferns are divided into two extant genera: Matonia

R.Br. ex Wall. and Phanerosorus Copeland (van Konijnenburg-van Cittert, 1993). They are unique

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in having resistant cuticles and a thick epidermis that allow Matoniaceae ferns to propagate in humid and warm areas, while being capable of tolerating harsher, cold conditions at night (van

Konijnenburg-van Cittert, 1993; Abbink et al., 2004).

Cycads and Ginkgos

Extant cycads mainly thrive in the tropics or subtropics, although their Mesozoic counterparts had a more diverse range of habitats that included rainforests, deltaic areas, slopes, and upland regions (Abbink et al., 2004 and references therein). Gingkos only have one representative extant species, Ginkgo biloba L., which is morphologically similar to its Mesozoic ancestors (Royer et al., 2003). Upper Cretaceous and Cenozoic Ginkgo were likely capable of tolerating a diverse range of climates, including hot and dry coastal plains, and wet and temperate lowlands. However, they likely mainly thrived in riparian habitats such as disturbed streamside and levee environments in humid and warm climates (Royer et al., 2003; Zhou, 2009). Cycads and/or gingkos are represented by Entylissa spp. and Cycadopites spp. pollen in the Treeless Creek section, although their precise botanical affinities remain uncertain (Zavialova et al., 2011).

Conifers

Pollen from five conifer families are identified in rocks sampled from the Treeless Creek section: Araucariaceae, Cheirolepidiaceae, Cupressaceae, Taxaceae, and Pinaceae. Araurariaceae conifers reached peak diversity during the Jurassic and declined towards the Late Cretaceous, where they were mainly restricted to moist, mesothermal habitats (Miller, 1977; Stockey, 1994;

Kershaw and Wagstaff, 2001; Kunzmann, 2007). Extant Araucariaceae conifers are now limited to the Southern Hemisphere where they mainly thrive in rainforests (Kershaw and Wagstaff, 2001).

Araucariaceae conifers are represented in the Treeless Creek section by Araucariacites australis pollen.

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Cheirolepidiaceae is a family of extinct conifers interpreted to have thrived in warm, xerophytic and thermophilous habitats (Alvin, 1982; Galloway et al., 2015 and references therein).

Cheirolepidiaceae conifers likely grew in upland areas with well-drained slopes or in lowland, coastal regions (Srivastava, 1976; Alvin, 1982; Galloway et al., 2015 and references therein). In the Treeless Creek section, Cheirolepidiaceae conifers are represented by Classopollis classoides pollen.

Cupressaceae and Taxaceae are sister conifer families that are generally known to thrive in temperate environments in upland regions, although they can grow in lowland regions where they can form stable plant communities (Galloway et al., 2013). Cupressaceae conifers occur globally and have a diverse ecological tolerance, including swamplands and deserts (Pittermann et al.,

2012). However, most Cupressaceae thrive in the margins of tropical and subtropical highlands

(de Laubenfels, 1984a). Mesozoic Cupressaceae are interpreted to have grown mainly in mesic- hydric habitats (Pittermann et al., 2012). Extant Taxaceae species mainly grow as shrubs or small trees, with a few exceptions that grow as large trees, e.g., Amentotaxus Pilger, and some species of Taxus L. and Torreya Walker-Arnott (Miller, 1977; Stewart and Rothwell, 1993). Taxaceae conifers today are predominantly limited to the Northern Hemisphere where they grow in moist habitats and temperate conditions or as part of the understory or canopy (de Laubenfels, 1984b).

In the Treeless Creek section, Cupressaceae and Taxaceae conifers are represented by

Perinopollenites elatoides and CT pollen.

Pollen with Pinaceae affinities preserved in samples collected from the Treeless Creek section include Cerebropollenites mesozoicus, Laricoidites magnus, and undifferentiated bisaccate pollen. Cerebropollenites pollen may be related to the living Tsuga Carrière (Shang and Zavada,

2003; Galloway et al., 2013, 2015). Tsuga is commonly found in mesic, temperate environments

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(Rogers, 1978; Calcote, 2003). Laricoidites pollen may be related to the extant Larix Miller (Wang et al., 2010), the living genera of which commonly thrive in cooler and moist environments (Arno,

1990; Johnston, 1990; Schmidt and Shearer, 1990). Bisaccate pollen were likely derived from conifers that grew in upland, arid or well-drained habitats (Galloway et al., 2015). These pollen grains have two air sacs located on each side of the central body and are adapted to be transported long distances via wind. They are also notoriously over-represented in sedimentary deposits.

Gnetophytes

Eucommiidites pollen have been found in situ from Early Jurassic and Early Cretaceous fossils seeds (Hughes, 1961; Brenner, 1963, 1967a; Reymanówna, 1968; Pedersen et al., 1989).

These seeds can be divided into three fossil genera: Erdtmanispermum Pedersen, Crane and Friis,

Spermatites Miner, and Allicospermum Harris (Mendes et al., 2008). Initially, Eucommiidites pollen were thought to have had an affinity with Gnetales (Doyle et al., 1975; Trevisan, 1980;

Crane, 1985), however, investigation into the pollen organs Erdmanitheca and Eucommiitheca revealed a relationship with Erdtmanithecales (Pedersen et al., 1989; Friss and Pedersen, 1996;

Mendes et al., 2008; 2010). Friis and Pedersen (1996) used Erdtmanithecales to accomodate those pollen organs and seeds with Eucommiidites pollen found in situ. Erdtmanithecales first appeared in the Early Jurassic fossil record, evidenced by dispersed pollen grains, and proliferated during the Early Cretaceous. Mesofossils, seeds, and in situ pollen grains are only known from the

Cretaceous fossil record (Rydin et al., 2006; Mendes et al., 2008). Erdtmanithecales is thought to be phylogenetically related to Bennettitales and Gnetales (Rydin et al., 2006 and references therein; Friis et al., 2007, 2009), although this relationship is only partly understood (Rydin and

Friis, 2010). Together, they form the Bennettitales-Erdtmanithecales-Gnetales (BEG) group.

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Only plants from the order Gnetales still exist, of which there are three living genera whose evolutionary history is poorly understood: Gnetum L., Ephedra L., and Welwitschia Hooker

(Raven et al., 2013). Gnetum are mainly large canopy climbers that predominantly occur in tropical environments where conditions are warm and moist (Raven et al., 2013; Ickert-Bond and Renner,

2015). Ephedra are shrubs that typically grow in arid environments such as deserts or seasonally dry habitats such as deciduous woodlands in the Northern Hemisphere and South America (Rydin et al., 2006; Raven et al., 2013; Ickert-Bond and Renner, 2015). Ephedra seeds are comparable to those of the extinct Erdtmanithecales (e.g., Rydin et al., 2006). Welwitschia is unique in that the majority of the plant body grows buried in sandy soil; only a woody, concave disk is exposed to the air from which two leaves grow. Welwitschia is known to only grow in the coastal deserts of southwest Africa (Raven et al., 2013). Mesozoic Welwitschia relatives were more widespread, occurring in Africa, South America, northeastern China, Central Asia, Europe, and eastern North

America (Friis et al., 2014). Mesozoic Gnetales likely had a much greater diversity and were more widespread than their living counterparts (Rydin et al., 2006). Eucommiidites pollen have a limited use in paleoenvironmental reconstruction as the evolutionary history, basic ecology, and morphology of the Gnetales remains an enigma (Ickert-Bond and Renner, 2015). In the Treeless

Creek section, the occurrence of Eucommiidites troedsonii pollen is interpreted to represent warm temperatures.

Paleoenvironment

Q-mode Cluster Analysis

Q-mode cluster analysis groups samples containing similar palynomorph taxa. In the

Treeless Creek section, Q-mode cluster analysis delineates three clusters that do not exhibit stratigraphic trends (Fig. 6). Clusters are not separated through changes in facies nor do they

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correspond to distinct stratigraphic positions. The three clusters generated from Q-mode cluster analysis are therefore interpreted as not stratigraphically unique. Differences between palynological samples, which are interpreted as representative of the paleo-plant community, are unlikely to have been directly caused by large-scale factors such as climate change that would be expected to result in marked changes in the assemblage over time. Detrended correspondence analysis (DCA) is conducted to determine principal variables that could have affected changes in the palynoassemblages in the Treeless Creek section.

Detrended Correspondence and R-mode Cluster Analyses

DCA was used to interpret two possible variables (DCA1, DCA2) that could have contributed to species ordination (Fig. 8). These interpretations are based on the palynomorphs that plot near the maximum variance of each axis and the inferred ecologies of their parent plants.

Araucariacites australis and Pinaceae pollen plot near the extreme negative values of DCA1.

Extant Araucariaceae conifers are capable of tolerating dry and cold conditions due to their leathery leaves and thick cuticles (Abbink et al., 2004). Pinaceae pollen in the Treeless Creek section include Cerebropollenites mesozoicus, Laricoidites magnus, and undifferentiated bisaccate pollen, all of which have parent plants that are thought to have been capable of thriving in cool or temperate conditions. This group is therefore interpreted to represent environments characterized by cooler temperatures. Species that plot near the maximum variance on DCA1 include

Aequitriradites spp., Schizaceae, and unclassified spores from the Class Polypodiopsida. Most spores identified in the Treeless Creek section plot within the positive variance on the DCA1 axis.

Together, this group is interpreted to represent warmer temperatures. The first principal factor

(DCA1) that influenced species ordination may therefore represent temperature variation.

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The second variable (DCA2) that influenced species ordination in the Treeless Creek section is interpreted as habitat. Abbink et al.’s (2004) sporomorph ecogroup (SEG) model outlined six SEGs for the Late Jurassic to Early Cretaceous: 1) Upland communities that reside above groundwater level; 2) Lowland regions such as plains, marshes, or swamps; 3) River vegetation that is periodically submerged in water; 4) Pioneer areas that are unstable and could have been recently affected by disturbances such as floods; 5) Coastal areas, which are not submerged by the sea but are affected by the proximity to salt spray; and 6) Tidally-influenced vegetation that is periodically submerged by the sea. In comparison, the Treeless Creek section is observed to have up to five different habitats reflected along the DCA2 axis (Fig. 8), with Abbink et al.’s (2004) River and Lowland SEG re-interpreted as two different habitats: peatlands and marshlands. Peatlands (i.e., swamps and bogs) are commonly found near bodies of water such as rivers. Peatlands are characterized by a prevalence of trees and shrubs, limited drainage, and periodic standing water (Environment Canada and Wilcox, 2002). Marshes are characterized by frequent flooding and a mixed assemblage of aquatic plants (Environment Canada and Wilcox,

2002). While both peatlands and marshlands are part of what is known as the wetlands, they are generally distinguished from one another by the presence of Sphagnum moss and the accumulation of peat in the former (Larsen, 1982). These habitats are often intermingled and grade into the other

(Larsen, 1982). A typical coastal wetland would be characterized by having an upland community adjacent to a swamp, with a marsh and a large body of water (Environment Canada and Wilcox,

2002; Fig. 8).

Palynomorphs that plot within the positive values on the DCA2 axis include Sphagnaceae,

Leptolepidites verrucatus, Matoniaceae, Gleicheniidites senonicus, and Osmundaceae spores, and

Entylissa spp., Eucommiidites troedsonii, and Cycadopites spp. pollen. Lycopodiales

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(Leptolepidites verrucatus) and Sphagnaceae spores are attributed to the River SEG of Abbink et al. (2004), although they can also be attributed to the Lowland SEG. In the Treeless Creek section, this group is attributed to swamps/peatlands due to the presence of Sphagnaceae spores and their close affinity with Sphagnum moss.

The group that includes Selaginellaceae and Lycopodiaceae spores is interpreted to possibly represent a Tidally-influenced SEG in Abbink et al.’s (2004) model, although they may be more representative of a Lowland and/or River SEG. Lycopodiaceae spores include Retitriletes, which Abbink et al. (2004) questionably attributed to the Tidally-influenced SEG, and

Sestrosporites pseudoalveolatus, which is attributed to the River SEG. Selaginellaceae spores preserved in rocks collected from the Treeless Creek section include Densoisporites spp. and

Neoraistrickia truncata. While Densoisporites spp. spores are associated with Pleuromiaceae family, some species of which are salt-tolerant, their presence may indicate a wetland habitat more so than a tidally-influenced habitat as most species are thought to grow inland with minimal marine influence (Retallack, 1997). Due to the broad ecological tolerances of Lycopodiaceae and

Selaginellaceae plants, the habitat of this group may belong to swamp/peatland and/or marshland habitats.

The group that includes Pinaceae pollen may represent a lowland conifer forest characteristic of swamps, with hinterland components. This group includes undifferentiated bisaccate, Cerebropollenites mesozoicus and Laricoidites magnus pollen. Consideration of the ecological tolerances of the parent plants of Cerebropollenites and Laricoidites pollen indicate a preference for moist habitats. In North America, conifer swamps are often populated by Larix

(Johnston, 1990), and hemlock (Tsuga) swamps are also common (e.g., Paratley and Fahey, 1986).

However, undifferentiated bisaccate pollen are more indicative of drier, upland communities,

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although their ordination with Cerebropollenites and Laricoidites pollen may be explained by long-distance transport of bisaccate pollen (Mudie, 1982; Chmura et al., 1999). In Quaternary research, wind-pollinated bisaccate pollen are the most resistant palynomorphs studied, are produced in the largest quantities, and tend to be transported the furthest distances (Traverse,

2007). Bisaccate pollen is therefore often overrepresented in marine sediment. However, site ordination of the Treeless Creek section shows neither a similarity nor a dissimilarity of bisaccate pollen to those samples with higher marine content. The presence of bisaccate pollen in this DCA group is thus inferred as representing deposition in the peatland rather than deposition in an open marine setting through aerial transportation. Combined, the positive values on the DCA2 axis are interpreted to represent a swamp/peatland habitat due to the presence of Sphagnaceae spores with undifferentiated bisaccate pollen sourced from the hinterland.

The group located nearest to the zero-value in Euclidean space on the DCA2 axis includes most of the fern spores and the Cupressaceae conifers. Cyatheaceae, Dicksoniaceae,

Osmundaceae, and Schizaceae ferns commonly grow near riverbanks (Abbink et al., 2004) and are therefore more indicative of marshland and riparian habitats. The presence of CT and

Perinopollenites elatoides pollen within this group are representative of the Cupressaceae conifers, which can grow in lowland regions such as swamps (Pittermann et al., 2012; Galloway et al.,

2013). Cupressaceae-Taxaceae pollen also represents the Taxaceae family, which includes shrubs or small trees that are common in swamp areas. Due to the absence of Sphagnaceae spores and the abundance of fern spores, this group is interpreted to represent a predominantly marshland habitat, with some swamp components expressed through CT and Perinopollenites elatoides pollen.

The group that plots nearest to the maximum negative variance of DCA2 includes

Classopollis classoides, Chasmatosporites spp., and Araucariacites australis pollen.

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Araucariaceae conifers often grew near coastal environments in Jurassic and Cretaceous times due to their high salt-tolerance (Abbink et al., 2004 and references therein). Cheirolepidiaceae conifers could have also thrived in lowland, coastal regions (Galloway et al., 2015 and references therein).

Paleoenvironmental interpretations based on Chasmastosporites spp. pollen are more enigmatic.

These pollen are questionably attributed to the Cycadales, though they may belong to the Gnetales or Ginkgoales (Boulter and Windle, 1993; Balme, 1995; Slater and Wellmann, 2015). Extant cycads typically have a low salt tolerance, with the exception of Zamia roezlii Linden, which thrive in coastal marshes (Goel and Khuraijam, 2015). Pollen grains of Zamia L. are closely comparable with those from the Androstrobus Schimper species (Couper, 1958; van Konijnenburg-van Cittert,

1971). Androstrobus pollen found in situ are considered equivalent to Chasmatosporites pollen

(Balme, 1995; Slater and Wellmann, 2015). This group is therefore interpreted to represent a coastal or near-coastal habitat. In summary, DCA2 is interpreted to represent habitats along a typical coastal wetland profile. These habitats are likewise reflected in R-mode cluster analysis

(Fig. 7).

R-mode cluster analysis delineates four groups that are interpreted as representing common co-occurrences between palynomorphs in the samples (Fig. 7). The spore assemblage is the largest group in terms of diversity, containing 32 out of the 46 taxa identified in the Treeless Creek section.

This assemblage is dominated by spores derived from ferns, fern allies, and mosses, with minor cycad and conifer pollen. The relatively large diversity of fern spores in this assemblage indicate a habitat with relatively high moisture and temperature. The spore assemblage is therefore interpreted to represent a lowland environment. In terms of the coastal wetland profile derived from DCA, this assemblage could represent the marshland component. The spore-pollen assemblage is predominantly composed of a mix of fern spores and conifer pollen. The increased

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diversity of pollen derived from conifers suggests the presence of an open forest. The O-D assemblage contains the species Densoisporites spp. and Osmundacidites wellmannii. The association of Osmundaceae spores with floodplain deposits and Densoisporites spores with

Isoetes suggests a transitional habitat between a lowland and coastal environment in which marine waters could have habitually influenced. The bisaccate-CT pollen assemblage is composed of undifferentiated Pinaceae bisaccate and undifferentiated Cupressaceae-Taxaceae pollen. The bisaccate-CT pollen assemblage may therefore represent a hinterland or upland forest dominated by Pinaceae and Cupressaceae-Taxaceae conifers.

The paleoenvironment for Treeless Creek is thus interpreted to be predominantly characterized by marsh and swampland vegetation, with components from an upland forest and vegetation from nearby coastlines. This conclusion is reflected in a study on Albian – lower

Cenomanian strata on the North Slope of Alaska (Shane, 1984). The Alaskan paleoenvironment was separated into delta front colonizers, mixed swamp forest vegetation and highland forest vegetation. Mixed swamp forest vegetation was considered to be the primary contributor of the spores and pollen, and was thought to have been likely transported by leveed distributaries and delta channels. Highland forest vegetation likely existed as a coniferous forests in the Brooks

Range.

Together, the Treeless Creek assemblages delineated from R-mode cluster analysis and

DCA suggest the existence of a typical coastal wetland and an upland, coniferous forest in the area during the mid-Cretaceous. Palynomorphs were likely transported fluvially through distributaries and delta channels similar to those suggested for the North Slope of Alaska, evidenced by the presence of a coal sample (C-604497) from the sandstone-dominated beds in the Treeless Creek section. This indicates the presence of local back swamp environments that would have been a

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proximal source of palynomorphs. The mode of transportation implies that the palynomorphs at

Treeless Creek were likely predominantly derived from local vegetation and are not representative of the vegetation on a regional scale, such as input into a drainage basin (Chmura et al., 1999).

Stratigraphically Constrained Cluster Analysis (CONISS)

Changes in the Treeless Creek palynoassemblage through time are expressed through three different clusters delineated by CONISS (Fig. 7). These changes are assumed to reflect changes in the paleoenvironment in the area and may give insight into mid-Cretaceous paleoclimate. Cluster

C is characterized by a peak in Laricoidites magnus and undifferentiated bisaccate pollen. The high relative abundances of these conifer pollen suggest possibly cooler temperatures during this time interval and an increased input from well-drained, upland communities. Cluster C is also characterized by the complete absence of dinocysts and may therefore represent the most regressive facies within the Treeless Creek section (McCarthy and Mudie, 1998; McCarthy et al.,

2003).

Cluster B is characterized by maximum fern diversity in the Treeless Creek section, a general increase in the relative abundance of fern spores, and a decrease in the relative abundance of bisaccate pollen. Spore species that reach their maximum relative abundance include

Granulatisporites spp., Stereisporites antiquasporites, Deltoidospora hallii, Kuylisporites lunaris,

Gleicheniidites senonicus, Ruffordiaspora cf. australiensis, Verrucosisporites spp.,

Lycopodiumsporites expansus, Cingutriletes clavus, Neoraistrickia truncata, Lycopodiumsporites crassimaceras, Sestrosporites pseudoalveolatus, Todisporites rotundiformis, Leiotriletes directus,

Cingulatisporites spp., Baculatisporites comaumensis, and Densoisporites spp. The increased diversity and relative abundance of fern and bryophyte spores and decrease in the relative abundance of bisaccate and Laricoidites pollen are interpreted as a transition into warmer and

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wetter conditions during this time interval, associated with climate change and/or relative sea-level rise. Increased development of vegetation from wetland areas may have contributed to this increased relative abundance and diversity (Shane, 1984). The marine-to-terrestrial ratio increases to 3-13% dinocysts, suggesting increasing marine influence and the presence of more distal facies.

Cluster A is characterized by a maximum in the marine-to-terrestrial ratio, reaching 40% dinocysts at 75 m. Dinocyst percentage decreases back to 0% at the base of the uppermost sandstone unit before reaching 6% at the top of the section. The maximum in dinocyst percentage therefore represents the distal-most facies in the Treeless Creek section and may reflect a possible flooding surface. A maximum in the relative abundance of Cerebropollenites mesozoicus,

Classopollis classoides, and CT pollen is likewise apparent towards the top of this cluster. While

Classopollis pollen is commonly a xeric indicator, the presence of this pollen has also been suggested to indicate forests nearby coastlines due to putative salt tolerance (Wall, 1965; Hughes and Moody-Stuart, 1966; Hughes, 1973; Srivastava, 1976; Alvin, 1982; Vakhrameev, 1987, 1991;

Galloway et al., 2015). Spore diversity decreases in comparison to Cluster B but remains higher than in Cluster C. Alongside the increase in the relative abundance of Cerebropollenites mesozoicus, Classopollis classoides, and CT pollen, Cluster A may represent predominantly humid and warm-temperate conditions.

The paleoenvironment for the Treeless Creek section is therefore interpreted as having predominantly wet conditions while experiencing a transition from cool to warm temperatures towards the Late Cretaceous. The area also likely experienced sea-level changes that are represented by the percentage of dinocysts in the observed marine-to-terrestrial ratio. Inferences about the paleoclimate are more difficult to determine. Understanding the paleoclimate requires determining polar paleotemperatures, which remains a subject of debate for the Cretaceous Period

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(e.g., Herman and Spicer, 1996; Jenkyns et al., 2004). Paleobotanical data for the North Slope of

Alaska suggests a humid and warm-temperate climate during the Albian-early Cenomanian

(Smiley, 1967; Shane, 1984). Smiley (1967) suggested a warm-temperate climate based on the presence of cycadophyte foliage at high-latitudes. However, Spicer and Parrish (1986) argued that

Mesozoic cycads had broader ecological tolerances in contrast to extant cycads, and suggested a cool-temperate climate for this time interval instead. This conclusion was based on reports of fossils of deciduous plants, toothed-leafed angiosperms and large-leafed conifers in Albian-

Cenomanian strata on the North Slope of Alaska. However, physiognomic analysis of fossil floras from the Russian high arctic, near the location of the mid-Cretaceous North Pole, suggest a climate similar to the modern temperate or warm-temperate zones (Herman and Spicer, 2010).

Interpretations of cooler paleotemperatures conflict with oxygen isotope data used to create

Cretaceous paleotemperature profiles for the Arctic Ocean (Jenkyns et al., 2004). These paleotemperature profiles indicate a warming period that lasted from the early Albian to the middle

Cenomanian, peaking during the early to middle Turonian. This trend is also reflected in foraminiferal stable isotope data at southern high-latitude regions (Huber et al., 2018). Inferences about the paleoclimate based on the Treeless Creek section are thus qualitative but it may have been temperate, perhaps even warm-temperate, and moist.

Angiosperms

No angiosperm pollen were observed in the Treeless Creek section. This paucity in angiosperm pollen is reflected in Galloway et al.’s (2012) study based on the upper Albian-

Cenomanian Hassel Formation in the Sverdrup Basin. Galloway et al. (2012) noted a low abundance and diversity of angiosperm pollen in the Hassel Formation in comparison to low- latitude regions in North America and suggested angiosperms may not have appeared in the

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Canadian Arctic until the late Albian. Other palynological studies based in the Arctic (Smiley,

1969; Shane, 1984; Retallack and Dilcher, 1986) also suggest that angiosperms may not have appeared in high-latitude regions in North America until the late Albian, with angiosperm pollen typically found in scarce abundances in upper Albian strata. In the North Slope of Alaska and in the Kome Formation of central West Greenland, angiosperm fossil flora were reported in strata of late Albian-Cenomanian age or younger, with fossils being extremely rare in uppermost Albian strata and then diversifying in strata of Late Cretaceous age or younger (Smiley, 1969; Dam et al.,

2009; Friis et al., 2011). In contrast, Mesozoic angiosperms were dominant in plant communities at lower latitudes, e.g., northeastern Asia or the USA, by the late Albian-Cenomanian (Brenner,

1967b; Herman, 2002; Spicer et al., 2002). Angiosperm diversity is also reportedly higher at lower latitudes. For example, middle Albian-Cenomanian strata from Alberta have a greater dicotyledonous angiosperm diversity than those from the upper Albian Christopher Formation

(Hopkins, 1974), Hassel Formation and equivalents in Arctic Canada (Hopkins and Balkwill,

1973; Doerenkamp et al., 1976; Plauchut and Jutard, 1976; Galloway et al., 2012).

Angiosperms appeared in Canada in the late Aptian (Axelrod, 1959) and did not proliferate into the Arctic until the late Albian. The paucity or absence of flowering plants in the Canadian

Arctic during this time interval suggests restricted angiosperm migration and/or limited angiosperm competition in the area. Poleward migration of angiosperms may have been hindered by geographical barriers resulting from the spread of a southern embayment of the Arctic Ocean and subsequent opening of the Western Interior Seaway during the mid-Cretaceous (Friis et al.,

2011; Galloway et al., 2012).

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CONCLUSION

Investigation of the dinocyst assemblage at Treeless Creek in the Aklavik Range, Beaufort-

Mackenzie area, reveals a need for biostratigraphic revision in the area. Comparisons between the dinocyst assemblages in the immediate area suggest that J/K boundary strata have limited exposures at Treeless Creek and may be entirely missing due to unconformities or unrecognized structural features. Based on the occurrence of the dinocyst species Oligosphaeridium cf. tenuiprocessum, the succession previously thought to be Berriasian in age is proposed instead to be late Albian in age and may extend into the early Cenomanian. Therefore, a possible stratigraphic attribution for this unit is the Arctic Red Formation. Quantitative multivariate statistical analyses of terrestrial palynomorphs demonstrate: 1) the existence of a coastal wetland community during this time period, and 2) a predominantly humid paleoenvironment that experienced a transition from cool to warm temperatures towards the Late Cretaceous.

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CHAPTER 4: SUMMARY AND FUTURE WORK

SUMMARY

The Jurassic-Cretaceous (J/K) boundary is the last Phanerozoic systems boundary to be defined in the Geologic Time Scale. The boundary is difficult to define globally due, in part, to biogeographical provincialism, which led to pervasive endemism of fauna and floras during the

Mesozoic. Three broad biogeographical realms developed as a result: Austral, Tethyan, and Boreal realms.

The primary marker for the J/K boundary in the Tethyan Realm is currently the base of the

Calpionella alpina Lorenz, 1902 Subzone (Wimbledon, 2017). The base of the Alpina Subzone is characterized by a widespread and consistent turnover in the rock record from large Calpionella and Crassicollaria to smaller, globular Calpionella alpina (Wimbledon et al., 2013; Schnabl et al.,

2015; Wimbledon, 2017). However, no such marker exists in the Boreal Realm, and many other potential markers such as calcareous nannofossils are absent in Boreal strata (Harding et al., 2011).

The Boreal Realm is confined to the Northern Hemisphere and includes the Arctic and parts of

Western Canada (Wimbledon et al., 2011). In Arctic Canada, Lower Cretaceous stages are defined primarily through the use of Buchia Rouillier, 1845 zones, with the J/K boundary at the base of the Buchia okensis Pavlow, 1907 Zone (Jeletzky, 1973, 1984; Zakharov, 1987; Urman et al.,

2014). In the Beaufort-Mackenzie area, the Husky Formation purportedly preserves the J/K boundary and is extensively exposed in the Aklavik Range in the northern Richardson Mountains.

The formation can be divided into four informal members: upper, red-weathering, arenaceous, and lower members. The J/K boundary is hypothesized to exist at or near the contact between the red- weathering and arenaceous members based on the occurrence of Buchia okensis near this contact

(Jeletzky, 1958; Brideaux, 1976). However, there is a paucity of age-diagnostic macrofossils in

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Arctic Canada and a general lack of diversity relative to those found in the Tethyan Realm.

Comparatively, palynomorphs are useful for biostratigraphic correlation as they are produced in high abundances and are exceptionally well-preserved. Palynomorphs are found in both Tethyan and Boreal realms as there were less endemism in terrestrial vegetation compared to marine organisms during the Mesozoic. Quantitative palynostratigraphy was thus conducted for two successions reported as early Berriasian in age: 1) the Husky Formation at the Martin Creek section and 2) the Martin Creek Formation at the Treeless Creek section.

Forty-seven mudstone samples were collected at the Martin Creek section from the arenaceous and red-weathering members of the Husky Formation. The samples yielded 53 genera and species of spores and pollen, and 14 genera and species of dinoflagellate cysts (i.e., dinocysts) that were consistent with the late Tithonian – early Berriasian age previously reported for the formation. Comparisons between existing palynological zonations reveal Concavissimisporites spores and the dinocyst Cribroperidinium cf. jubaris as potentially important taxa for biostratigraphic correlation of J/K boundary strata. However, these palynological zonations have a limited use as many biostratigraphically important taxa previously identified in J/K boundary strata (e.g., Cicatricosisporites purbeckensis spores, Tetrachacysta spinosigibberosa) are absent in this study. Definition of the J/K boundary in Arctic Canada therefore remains dependent mainly on the Buchia zones established by Jeletzky (1965, 1966, 1969, 1970, 1971, 1973, 1979, 1984).

Multivariate statistical analyses such as hierarchical cluster analyses are used to delineate stratigraphically unique terrestrial palynoassemblages in the Martin Creek section. Ordination techniques such as detrended correspondence analysis (DCA) reveals two principal variables, moisture and temperature, that affected the composition of the palynoassemblage. Combined, stratigraphically unique palynoassemblages and these two principal variables suggest that large

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scale factors such as climate likely influenced changes in the palynoassemblage. Stratigraphically constrained cluster analysis (CONISS) reveals these changes expressed through time. A palynological signature can therefore be identified for the J/K boundary, and the paleoenvironment during the Early Cretaceous can likewise be elucidated. In the Martin Creek section, two major palynological events are observed: 1) a period when palynoassemblages were dominated by 44-

82% of Cupressace-Taxaceae (CT) pollen, interpreted to represent an increase in humidity; followed by 2) a time characterized by increased relative abundance of Classopollis classoides pollen, peaking at 25%, which is interpreted to represent a seasonally arid phase. These results for the Martin Creek section suggest that the paleoclimate during the J/K transition in Arctic Canada may have been comparatively less arid than those regions in the Tethyan Realm.

Sixteen samples were collected at the Treeless Creek section from the purported Martin

Creek Formation. The samples yielded 46 genera and species of spores and pollen, and 11 species and genera of dinocysts. The dinocyst species identified at the Treeless Creek section can be divided into three distinct ages: Late Jurassic, Valanginian-Barremian, and Albian-Cenomanian.

The dinocyst with the narrowest age range is the species Oligosphaeridium cf. tenuiprocessum, which is limited to upper Albian to lower Cenomanian strata. Comparisons to other paleontological reports of the Geological Survey of Canada reveal consistently younger ages in the formations in the Treeless Creek area. Thus, J/K boundary strata likely have minimal exposures in the area. The late Albian age indicated by the dinocyst assemblage at the Treeless Creek section suggests equivalence with the Arctic Red Formation or Jeletzky’s (1960) Albian shale-siltstone division, both of which have only been reported in the Peel Plateau region (Mountjoy and Chamney, 1969) and between the Richardson Mountains and Mackenzie River (Norris, 1981; Dixon, 1992), respectively. While the samples at Treeless Creek cannot provide information on the J/K transition,

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quantitative palynostratigraphy is useful for reconstructing the paleoenvironment in polar regions.

This could provide further insight into mid-Cretaceous paleoclimate in high-latitude regions, as polar paleoclimates are still a subject of some controversy (Herman and Spicer, 1996; Price, 1999;

Price et al., 2000; Price and Nunn, 2010; Galloway et al., 2012. 2015; Price and Passey, 2013;

Grasby et al., 2017). Hierarchical cluster analyses and DCA reveal the existence of a coastal wetland habitat where most palynomorphs preserved in the Treeless Creek section were likely deposited. CONISS delineated three clusters that revealed high moisture conditions likely prevailed in the area during the mid-Cretaceous. Trends in the relative abundance of gymnosperm pollen such as Laricoidites magnus, Cerebropollenites mesozoicus, Classopollis classoides,

Cupressaceae-Taxaceae, and undifferentiated bisaccate pollen suggest that paleotemperatures transitioned from cool to warm towards the Late Cretaceous.

FUTURE WORK

A palynological signature revealed through use of multivariate statistical analyses from the

Jurassic-Cretaceous Deer Bay Formation, Sverdrup Basin, could provide further insight into the possible changes in climate interpreted from the Husky Formation in the Martin Creek section. A consistent palynological signature could potentially be used for biostratigraphic correlations of J/K boundary strata in place of Buchia zones where these macrofossils are sparse or absent.

Magnetostratigraphic studies on J/K strata in Arctic Canada could also supplement chronostratigraphic correlation between biogeographic realms as seen in Houša et al. (2007), and may be particularly useful considering the recent revisement of magnetozone M19n.2n as one of the Tethyan markers for the J/K boundary (Wimbledon, 2017).

There is a need to revisit the Treeless Creek area to identify any unrecognized structures or unconformities. Norris’ (1981) map indicates extensive exposures of the Husky Formation in

217

the area; many of these may have been misidentified and may require review. Comparisons between the dinocyst assemblages from this study and Jeletzky’s Albian shale-siltstone division

(plotted as Arctic Red Formation in Norris’ map) would be useful for more precise age estimations.

The terrestrial palynoasemblage for this formation could lead to further insight into mid-

Cretaceous polar paleoclimates, thus helping improve our understanding of Cretaceous climates.

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PLATES

Slide information includes: GSC Calgary curation number (C-#), preparation number (P-#), sieve size, and England Finder coordinates.

224

SPORES AND POLLEN

Plate 1

Division Marchantiophyta

Figure 1: Aequitriradites sp. (C-604459, P5366-25B, UN, M25-4)

Figure 2: Aequitriradites sp. (C-604466, P5366-32B, UN, P26-3)

Figure 3: Aequitriradites sp. (C-604472, P5366-38B, UN, O27-1)

Figure 4: Aequitriradites sp. (C-604455, P5366-21B, C-604455, UN, H34-2)

Figure 5: Aequitriradites sp. (C-604488, P5366-54B, UN, U20-2)

Division Bryophyta, Classes Sphagnopsida, incertae sedis

Figure 6: Cingulatisporites sp. (C-604488, P5366-54B, UN, S13-2)

Figure 7: Cingulatisporites sp. (C-604499, P5366-64B, UN, H26-2)

Figure 8: Cingulatisporites sp. (C-604470, P5366-36B, UN, K15-2,4)

Figure 9: Stereisporites antiquasporites (C-604447, P5366-13B, UN, L24-4)

Figure 10: Stereisporites antiquasporites (C-604469, P5366-35B, UN, N47-4)

Figure 11: Cingutriletes clavus (C-604471, P5366-37B, UN, N46-2)

Figure 12: Annulispora folliculosa (C-604476, P5366-42B, UN, L27-3)

Figure 13: Annulispora folliculosa (C-604449, P5366-15B, UN, K7-2)

Figure 14: Rogalskaisporites cicatricosus (C-604470, P5366-36B, UN, L42-1)

225

226

Plate 2

Division Tracheophyta, Class Lycopodiopsida, Order Lycopodiales

Figure 15-16: Camarozonotriletes insignis (C-604404, P5366-1B, UN, Q16-2); 15 – proximal;

16 – distal.

Figure 17: Coronatispora valdensis (C-604475, P5366-41B, UN, S26-4)

Figure 18: Lycopodiumsporites crassimacerius (C-604491, P5366-57B, UN, F49-2)

Figure 19: Lycopodiumsporites expansus (C-604439, P5366-5B, UN, P20-4)

Figure 20: L. expansus (C-604480, P5366-46B, UN, S33-2)

Figure 21: Lycopodiumsporites marginatus (C-604474, P5366-40B, UN, N47-4)

Figure 22: L. marginatus (C-604473, P5366-39B, UN, P45-2)

Figure 23: Retitriletes austroclavatidites (C-604443, P5366-9B, UN, J25-2,4)

Figure 24: R. austroclavatidites (C-604446, P5366-12B, UN, G10-4)

Figure 25: Sestrosporites pseudoalveolatus (C-604439, P5366-5B, UN, W17-2)

Figure 26-27: Leptolepidites verrucatus (C-604452, P5366-18B, UN, P23-2); 26 – proximal; 27

– distal.

227

228

Plate 3

Division Tracheophyta, Class Lycopodiopsida, Orders Selaginellales, incertae sedis

Figure 28: Densoisporites sp. (C-604453, P5366-19B, UN, N37-1)

Figure 29: Densoisporites sp. (C-604489, P5366-55B, UN, T21-2)

Figure 30: Densoisporites sp. (C-604439, P5366-5B, UN, R25-3)

Figure 31: Foveosporites sp. (C-604468, P5366-34B, UN, L36-2)

Figure 32: Neoraistrickia truncata (C-604479, P5366-45B, UN, K29-3)

Figure 33: N. truncata (C-604445, P5366-11B, UN, N20-3)

Figure 34: Lycospora sp. (C-604471, P5366-37B, UN, N38-2)

Figure 35: Lycospora sp. (C-604471, P5366-37B, UN, J29-4)

229

230

Plate 4

Division Tracheophyta, Class Polypodiopsida, Order Cyatheales

Figure 36: Cibotiumspora juncta (C-604438, P5366-4B, UN, U29-3)

Figure 37: Concavissimisporites apiverrucatus (C-604477, P5366-43B, UN, L15-1)

Figure 38: Concavissimisporites variverrucatus (C-604438, P5366-4B, UN, F38-1)

Figure 39: C. variverrucatus (C-604439, P5366-5B, UN, V38-3)

Figure 40: Concavissimisporites crassatus (C-604439, P5366-5B, UN, U19-3)

Figure 41: Cyathidites australis (C-604468, P5366-34B, UN, K20-3)

Figure 42: C. australis (C-604453, P5366-19B, UN, O34-1)

Figure 43: C. australis (C-604435, P5366-2B, UN, G26-4)

Figure 44: C. australis (C-604455, P5366-21B, UN, H36-2)

Figure 45: Cyathidites minor (C-604446, P5366-12B, UN, F29-1)

Figure 46: Deltoidospora hallii (C-604439, P5366-5B, UN, T14-1)

Figure 47: D. hallii (C-604457, P5366-23B, UN, K13-3)

Figure 48-49: Kuylisporites lunaris (C-604499, P5366-64B, UN, E38-2); 48 – proximal; 49 –

distal.

Figure 50: K. lunaris (C-604489, P5366-55B, UN, S38-3)

231

232

Plate 5

Division Tracheophyta, Class Polypodiopsida, Orders Gleicheniales, Osmundales

Figure 51: Gleicheniidites senonicus (C-604450, P5366-16B, UN, K41-2)

Figure 52: G. senonicus (C-604441, P5366-7B, UN, E15-4)

Figure 53: Dictyophyllidites harrisii (C-604446, P5366-12B, UN, G36-3)

Figure 54-55: Matonisporites crassiangulatus (C-604404, P5366-1B, UN, P30-2); 54 – distal; 55

– proximal.

Figure 56: Baculatisporites comaumensis (C-604456, P5366-22B, UN, H35-4)

Figure 57: B. comaumensis (C-604467, P5366-33B, UN, G50-1)

Figure 58: B. comaumensis (C-604452, P5366-18B, UN, L35-1)

Figure 59: Biretisporites potoniaei (C-604494, P5366-60B, UN, J29-4)

Figure 60: B. potoniaei (C-604469, P5366-35B, UN, S38-3)

Figure 61: Osmundacidites wellmanii (C-604437, P5366-3B, UN, G22-3)

Figure 62: O. wellmanii (C-604499, P5366-64B, UN, F48-4)

Figure 63: Rugulatisporites sp. (C-604467, P5366-33B, UN, H40-1)

Figure 64: Todisporites rotundiformis (C-604456, P5366-22B, UN, K31-1)

Figure 65: T. rotundiformis (C-604446, P5366-12B, UN, G36-3)

233

Figure 66: Verrucosisporites sp. (C-604474, P5366-40B, UN, R52-3)

Figure 67: Verrucosisporites sp. (C-604499, P5366-64B, UN, J31-1)

234

Plate 6

Division Tracheophyta, Class Polypodiopsida, Orders Polypodiales, Schizaeales, incertae sedis

Figure 68: Contignisporites cooksonii (C-604446, P5366-12B, UN, F35-3)

Figure 69: Distaltriangularisporites sp. (C-604486, P5366-52B, UN, K21-1)

Figure 70-71: Klukisporites sp. (C-604439, P5366-5B, UN, O38-3); 70 – proximal view; 71 –

distal view

Figure 72: Ruffordiaspora cf. australiensis (C-604481, P5366-47B, UN, G32-3)

Figure 73: R. cf. australiensis (C-604457, P5366-23B, UN, K34-1)

Figure 74: R. cf. australiensis (C-604492, P5366-58B, UN, H48-1)

Figure 75: Apiculatisporites sp. (C-604499, P5366-64B, UN, K31-2)

Figure 76: Leiotriletes directus (C-604492, P5366-58B, UN, M23-4)

Figure 77: L. directus (C-604488, P5366-54B, UN, R35-4)

Figure 78: Undulatisporites undulapolus (C-604447, P5366-13B, UN, K16-4)

Figure 79: U. undulapolus (C-604442, P5366-8B, UN, J38-4)

235

236

Plate 7

Division Tracheophyta, Class incertae sedis

Figure 80: Acanthotriletes varispinosus (C-604474, P5366-40B, UN, O48-2)

Figure 81: Granulatisporites sp. (C-604466, P5366-32B, C-604466, UN, Q38-2)

Figure 82: Granulatisporites sp. (C-604484, P5366-50B, UN, L39-1)

Figure 83: Psilatriletes radiatus (C-604492, P5366-58B, UN, G36-4)

Figure 84: Triquitrites sp. (C-604456, P5366-22B, UN, K20-2)

237

238

Plate 8

Division Tracheophyta, Classes Cycadopsida, Ginkgoopsida, Gnetopsida

Figure 85: Chasmatosporites sp. (C-604474, P5366-40B, UN, M26-4)

Figure 86: Chasmatosporites sp. (C-604404, P5366-1B, UN, T10-4)

Figure 87: Cycadopites sp. (C-604442, P5366-8B, UN, K32-2)

Figure 88: Cycadopites sp. (C-604464, P5366-30B, UN, J15-1)

Figure 89: Cycadopites sp. (C-604452, P5366-18B, UN, P26-2)

Figure 90: Cycadopites sp. (C-604440, P5366-6B, UN, H19-4)

Figure 91: Cycadopites sp. (C-604446, P5366-12B, UN, E24-3)

Figure 92: Entylissa sp. (C-604491, P5366-57B, UN, G39-2)

Figure 93: Eucommiidites troedsonii (C-604491, P5366-57B, UN, L32-4)

239

240

Plate 9

Division Tracheophyta, Class Pinopsida

Figure 94: Araucariacites australis (C-604451, P5366-17B, UN, J19-3)

Figure 95: Classopollis tetrad (C-604457, P5366-23B, UN, K10-4)

Figure 96: Classopollis tetrad (C-604456, P5366-22B, UN, J13-4)

Figure 97: Classopollis classoides (C-604455, P5366-21B, UN, H39-4)

Figure 98: Perinopollenites elatoides (C-604468, P5366-34B, UN, J36-2)

Figure 99: P. elatoides (C-604447, P5366-13B, UN, J34-2)

Figure 100: Undifferentiated Cupressaceae-Taxaceae pollen (C-604439, P5366-5B, UN, U20-4)

Figure 101: Undifferentiated Cupressaceae-Taxaceae pollen (C-604478, P5366-44B, UN, H31-

Figure 102: Cerebropollenites mesozoicus (C-604472, P5366-38, UN, R38-1)

Figure 103: Laricoidites magnus (C-604447, P5366-13B, UN, K21-2)

Figure 104: Undifferentiated bisaccate pollen (C-604447, P5366-13B, UN, K43-1)

241

242

DINOFLAGELLATE CYSTS

Plate 10

Division Dinoflagellata, Class Dinophyceae, Order Gonyaulacales, Family Ceratiaceae

Figure 105: Muderongia tetracantha (C-604491, P5366-57B, UN, E42-4)

Figure 106: M. tetracantha (C-604489, P5366-55B, UN, U22-3)

Figure 107: M. tetracantha (C-604492, P5366-58B, UN, L25-1)

Figure 108: M. tetracantha (C-604488, P5366-54B, UN, N40-3)

243

244

Plate 11

Division Dinoflagellata, Class Dinophyceae, Order Gonyaulacales, Family Gonyaulacaceae,

Subfamilies Cribroperidinioideae and Gonyaulacoideae

Figure 109: Cribroperidinium cf. jubaris (C-604468, P5366-34B, UN, J38-2,4)

Figure 110: C. cf. jubaris (C-604468, P5366-34B, UN, K42-4)

Figure 111: Spiniferites sp. (C-604482, P5366-48B, UN, M13-2)

Figure 112: Spiniferites sp. (C-604484, P5366-50B, UN, J42-3)

Figure 113: Spiniferites sp. (C-604489, P5366-55B, UN, U23-3)

Figure 114: Spiniferites sp. (C-604486, P5366-52B, UN, Q44-3); possibly Achomosphaera cf.

neptuni (Eisenack) Davey and Williams

Figure 115-116: Spiniferites sp. (C-604493, P5366-59B, UN, S47-3); 115 – anterior focus; 116 –

posterior focus

245

246

Plate 12

Division Dinoflagellata, Class Dinophyceae, Order Gonyaulacales, Family Gonyaulacaceae,

Subfamily Gonyaulacoideae

Figure 117: Tubotuberella rhombiformis (C-604448, P5366-14B, UN, H12-4)

Figure 118: T. rhombiformis (C-604448, P5366-12B, UN, E27-4)

Figure 119: T. rhombiformis (C-604468, P5366-34B, UN, K44-2)

Figure 120: Psaligonyaulax sp. (C-604452, P5366-18B, UN, L43-4)

247

248

Plate 13

Division Dinoflagellata, Class Dinophyceae, Order Gonyaulacales, Family Gonyaulacaceae,

Subfamily Leptodinioideae

Figure 121: Cymososphaeridium sp. (C-604480, P5366-46B, UN, J26-1)

Figure 122: Dichadogonyaulax sp. (C-604468, P5366-34B, UN, J29-3)

Figure 123: Gonyaulacysta dualis (C-604489, P5366-55B, UN, U37-3)

Figure 124: G. dualis (C-604495, P5366-61B, UN, O32-1)

Figure 125: Gonyaulacysta? pectinigera (C-604484, P5366-50B, UN, G21-3)

Figure 126: Gonyaulacysta? pectinigera (C-604476, P5366-42B, UN, R38-2)

249

250

Plate 14

Division Dinoflagellata, Class Dinophyceae, Order Gonyaulacales, Family Gonyaulacaceae,

Subfamily Leptodinioideae

Figure 127: Stiphrosphaeridium anthophorum (C-604492, P5366-58B, UN, L48-2)

Figure 128: Oligosphaeridium pulcherrimum (C-604499, P5366-64B, UN, K24-3)

Figure 129: O. pulcherrimum (C-604495, P5366-61B, UN, P44-4)

Figure 130: O. pulcherrimum (C-604486, P5366-52B, UN, Q46-1)

251

252

Plate 15

Division Dinoflagellata, Class Dinophyceae, Order Gonyaulacales, Family Gonyaulacaceae,

Subfamily Leptodinioideae

Figure 131: Oligosphaeridium cf. tenuiprocessum (C-604495, P5366-61B, UN, T32-2)

Figure 132: O. cf. tenuiprocessum (C-604482, P5366-48B, UN, O15-1)

Figure 133: O. cf. tenuiprocessum (C-604487, P5366-53B, UN, O45-1)

Figure 134: O. cf. tenuiprocessum (C-604486, P5366-52B, UN, Q17-4)

253

254

Plate 16

Division Dinoflagellata, Class Dinophyceae, Order Gonyaulacales, Family Gonyaulacaceae,

Subfamily Leptodinioideae

Figure 135: Rhynchodiniopsis sp. (C-604446, P5366-12B, UN, E28-3)

Figure 136: Sirmiodinium grossii (C-604481, P5366-47B, UN, G21-4)

Figure 137: S. grossii (C-604404, P5366-1B, UN, Q42-2)

Figure 138: Hystrichodinium sp. (C-604489, P5366-55B, UN, V35-1)

Figure 139: Kiokansium williamsii (C-604499, P5366-64B, UN, J37-1)

Figure 140: Kiokansium cf. unituberculatum (C-604499, P5366-64B, UN, J48-2)

Figure 141: Trichodinium cf. erinaceoides (possibly Cometodinium) (C-604455, P5366-21B,

UN, H30-1)

Figure 142: T. cf. erinaceoides (C-604461, P5366-27B, UN, L16-1)

255

256

Plate 17

Division Dinoflagellata, Class Dinophyceae, Order Gonyaulacales, Family Pareodiniaceae

Figure 143: Paragonyaulacysta? borealis (alternatively identified as Pareodinia ceratophora)

(C-604447, P5366-13B, C-604447, UN, J26-3)

Figure 144: P. borealis (C-604447, P5366-13B, UN, J33-1)

Figure 145: P. borealis (C-604486, P5366-52B, UN, M45-2)

Figure 146: P. borealis (C-604460, P5366-26B, UN, L19-1)

Figure 147: P. borealis (C-604468, P5366-34B, UN, J20-3)

Figure 148: P. borealis (C-604446, P5366-12B, UN, G27-4)

257

258

Plate 18

Division Dinoflagellata, Class Dinophyceae, Order Gonyaulacales, Family incertae sedis

Figure 149: Epiplosphaera cf. saturnalis (C-604490, P5366-56B, UN, N14-2)

Figure 150: E. cf. saturnalis (C-604484, P5366-50B, UN, N19-1)

Figure 151: E. cf. saturnalis (C-604499, P5366-64B, UN, F39-1,3)

259

260

UNKNOWN PALYNOMORPHS

Plate 19

Figure 152: Unknown (C-604435, P5366-2E, -20, M6-4)

Figure 153: Unknown (C-604435, P5366-2E, -20, N31-4)

Figure 154: Unknown (C-604435, P5366-2E, -20, N29-2)

Figure 155: Unknown (C-604435, P5366-2E, -20, N19-2)

Figure 156-157: Unknown (C-604435, P5366-2E, -20, N22-2); 156 – proximal; 157 – distal

Figure 158-159: Unknown (C-604435, P5366-2E, -20, K16-2); 158 – proximal; 159 – distal

Figure 160: Unknown (C-604435, P5366-2E, -20, N32-3)

Figure 161: Unknown (C-604435, P5366-2E, -20, J37-1)

Figure 162: Unknown (C-604435, P5366-2E, -20, N17-1)

Figure 163: Unknown (C-604435, P5366-2E, -20, K19-1)

Figure 164: Unknown (C-604435, P5366-2E, -20, L41-2)

Figure 165: Unknown (C-604435, P5366-2E, -20, O38-1)

Figure 166: Unknown (C-604435, P5366-2E, -20, M11-3)

Figure 167: Unknown (C-604435, P5366-2E, -20, J37-1)

Figure 168: Unknown uniplanar tetrad (C-604435, P5366-2B, UN, E18-2)

Figure 169: Unknown uniplanar tetrad (C-604435, P5366-2E, -20, M28-3)

261

Figure 168: Unknown palynomorph triad (C-604435, P5366-2B, UN, E18-2)

Figure 169: Unknown palynomorph triad (C-604435, P5366-2E, -20, M28-3)

262

APPENDICES

263

A p p e n d i x 1 : R a w c o u n t d a t a o f s p o r e s a n d p o l l e n a n d d i n o f l a g e l l a t e c y s t s f r o m t h e H u s k y F o r m a t i o n a t t h e M a r t i n C r e e k s e c t i o n . A l l s a m p l e s a r e s t o r e d a t t h e G e o l o g i c a l S u r v e y o f C a n a d a , C a l g a r y O f f i c e . S l i d e i n f o r m a t i o n i n c l u d e s c u r a t i o n ( C - # ) a n d p r e p a r a t i o n ( P - # ) n u m b e r s .

SPORES AND POLLEN

P# P5366-1B P5366-2B P5366-3B P5366-4B P5366-5B P5366-6B P5366-7B P5366-8B P5366-9B P5366-10B P5366-11B P5366-12B P5366-13B P5366-14B P5366-15B P5366-16B P5366-17B P5366-18B P5366-19B P5366-20B P5366-21B P5366-22B P5366-23B P5366-24B P5366-25B P5366-26B P5366-27B P5366-28B P5366-29B P5366-30B P5366-31B P5366-32B P5366-33B P5366-34B P5366-35B P5366-36B P5366-37B P5366-38B P5366-39B P5366-40B P5366-41B P5366-42B P5366-43B P5366-44B P5366-45B P5366-46B P5366-47B C# C-604404 C-604435 C-604437 C-604438 C-604439 C-604440 C-604441 C-604442 C-604443 C-604444 C-604445 C-604446 C-604447 C-604448 C-604449 C-604450 C-604451 C-604452 C-604453 C-604454 C-604455 C-604456 C-604457 C-604458 C-604459 C-604460 C-604461 C-604462 C-604463 C-604464 C-604465 C-604466 C-604467 C-604468 C-604469 C-604470 C-604471 C-604472 C-604473 C-604474 C-604475 C-604476 C-604477 C-604478 C-604479 C-604480 C-604481 Thickness (m) 94 92 90 88 86 84 82 80 78 76 74 72 70 68 66 64 62 60 58 56 54 52 50 48 46 44 42 40 38 36 34 32 30 28 10.8 10.7 10.4 10 9 7 6 5.4 3.2 2.1 2 0.3 0.25 Acanthrotriletes varispinosus 3 0 0 0 0 0 1 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 2 0 1 0 2 0 0 1 0 0 Aequitriradites spp. 6 3 3 0 8 4 0 1 3 1 0 2 1 1 0 0 3 3 0 1 1 1 1 1 2 3 1 0 0 0 0 1 1 1 0 3 1 1 1 6 3 2 0 0 0 0 0 Annulispora folliculosa 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 1 1 0 0 0 0 0 0 0 0 0 Araucariacites australis 11 10 21 42 26 10 8 7 8 9 3 10 10 14 6 5 18 1 4 14 9 6 5 5 3 2 5 0 1 6 2 0 11 34 0 0 0 1 0 2 0 0 2 0 0 1 0 Baculatisporites comaumensis 1 3 4 2 13 5 6 7 7 4 3 6 10 8 7 2 6 16 8 8 5 7 7 3 4 9 10 0 3 2 0 0 18 18 31 27 44 63 35 43 32 36 19 28 17 22 21 Biretisporites potoniaei 3 0 0 0 7 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 2 0 0 0 2 0 0 0 0 0 1 1 0 Camarazonotriletes insignis 4 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 2 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Cerebropollenites mesozoicus 3 4 2 0 0 7 7 5 3 4 0 9 1 3 5 2 5 2 3 5 3 2 2 2 3 1 0 1 1 3 1 2 1 0 11 11 14 2 11 17 18 18 17 20 18 21 14 Chasmatosporites spp. 2 5 2 1 2 5 4 3 3 0 1 2 4 0 2 2 2 0 0 0 9 0 2 1 2 2 1 0 1 0 2 1 0 1 0 4 0 0 1 1 1 2 3 9 7 0 3 Cibotiumspora juncta 0 0 0 1 2 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 2 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Cingulatisporites spp. 10 8 1 1 3 9 4 7 1 0 0 1 4 4 1 1 1 2 4 1 0 2 0 4 4 0 1 0 0 1 3 2 4 5 1 4 0 1 2 4 3 1 1 4 1 5 0 Cingutriletes clavus 0 2 0 0 6 2 0 0 4 1 2 0 2 1 1 2 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 3 0 0 0 1 0 4 3 1 0 0 2 2 1 1 Classopollis classoides 8 4 3 2 9 16 10 10 6 10 16 29 16 15 15 9 15 18 17 14 54 91 59 65 22 20 11 1 4 6 0 1 0 0 18 27 3 6 11 15 10 16 18 13 19 21 23 Concavissimisporites apiverrucatus 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 Concavissimisporites crassatus 7 0 0 0 5 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 Concavissimisporites variverrucosus 0 1 0 1 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Contignisporites cooksonii 0 0 0 0 0 1 0 0 0 3 0 1 0 1 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Coronatispora valdensis 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 2 0 0 1 0 0 Cupressaceae-Taxaceae pollen 46 29 46 69 40 71 145 68 96 128 126 103 78 91 125 127 108 98 175 114 91 114 100 112 180 71 122 333 351 254 307 320 141 159 53 86 76 67 85 80 86 75 93 76 67 90 72 Cyathidites australis 27 18 10 11 33 14 19 23 13 16 10 22 15 16 15 7 12 13 12 12 18 6 0 10 11 26 20 0 1 10 2 2 22 18 53 31 37 40 27 29 35 31 16 18 26 32 34 Cyathidites minor 0 0 0 0 0 0 0 2 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Cycadopites spp. 2 2 1 5 1 3 4 1 3 0 2 9 0 0 0 0 4 4 0 0 3 0 1 0 2 0 1 0 1 1 0 0 8 3 0 4 11 2 2 7 0 1 0 2 3 3 1 Deltoidospora hallii 1 9 18 10 15 27 3 10 3 11 7 9 7 1 8 5 20 3 0 4 2 4 8 1 7 11 14 0 9 4 2 0 3 4 7 0 12 2 2 0 3 0 1 1 1 2 3 Densoisporites spp. 14 14 9 20 16 14 12 12 16 5 6 1 9 7 2 9 8 6 9 5 2 3 1 0 2 6 9 0 2 2 6 1 3 1 4 1 2 4 3 0 4 1 0 3 3 1 1 Dictyophyllidites harrisii 2 0 0 0 0 1 0 1 0 0 2 5 1 0 1 1 1 6 0 4 0 1 0 0 0 4 2 0 1 3 0 1 4 1 0 0 3 1 0 0 0 0 1 0 0 0 0 Entylissa spp. 1 1 4 1 0 0 0 2 3 1 0 1 0 0 0 0 0 0 1 0 0 1 0 0 0 0 0 0 0 1 0 0 0 2 0 0 0 0 0 0 0 0 0 0 0 0 0 Eucommiidites troedsonii 0 0 0 0 2 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Foveosporites spp. 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 1 2 0 0 0 0 0 0 0 0 0 Gleicheniidites senonicus 3 9 4 4 5 2 3 8 6 2 3 0 2 6 0 3 3 2 3 1 4 3 3 2 1 1 1 0 0 0 1 0 0 1 0 1 7 0 2 1 2 4 0 2 0 1 3 Granulatisporites spp. 6 3 0 3 14 2 0 0 0 0 0 0 0 0 0 0 0 2 0 0 0 1 0 0 0 0 2 0 0 0 0 2 5 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Ischyosporites spp. 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 2 0 1 0 0 0 0 0 0 0 0 0 0 1 1 Klukisporites spp. 0 1 0 0 3 0 0 0 0 0 0 0 0 0 0 1 1 0 0 0 0 2 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Laricoidites magnus 26 11 27 17 9 14 10 17 14 19 39 25 25 21 12 16 20 13 14 9 17 14 14 19 13 9 11 3 2 7 8 17 6 22 1 9 4 2 8 7 2 10 11 5 3 6 11 Leptolepidites verrucatus 0 10 13 2 8 5 5 0 5 3 3 6 1 3 1 0 1 4 0 1 4 1 0 5 0 2 4 0 2 0 2 1 8 0 2 2 0 2 4 4 0 0 2 1 0 0 0 Lycopodiumsporites expansus 5 2 6 2 2 1 0 3 2 0 0 0 3 1 1 0 1 1 0 1 0 1 0 0 0 0 0 0 0 0 0 0 0 2 4 12 8 6 5 3 0 4 2 0 0 4 5 Lycopodiumsporites marginatus 4 1 1 1 2 3 0 0 1 2 2 0 1 1 0 1 3 1 0 0 0 0 0 1 3 0 2 0 0 0 0 0 0 0 0 0 0 0 0 2 4 1 0 1 0 0 0 Lycospora spp. 4 0 2 1 4 5 0 0 3 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 2 0 0 0 0 0 0 0 1 2 1 1 3 2 5 3 3 1 0 1 1 2 1 Matonisporites crassiangulatus 3 3 2 6 8 0 3 2 5 4 0 2 1 4 4 2 3 0 0 0 1 0 2 1 0 1 0 0 1 1 0 0 0 0 0 0 0 1 0 0 0 0 0 0 7 0 0 Neoraistrickia truncata 1 0 0 0 1 0 0 1 1 0 1 0 2 2 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 1 4 2 1 3 1 1 1 0 0 5 4 1 1 0 Osmundacidites wellmannii 32 42 35 34 27 9 30 40 49 48 43 38 50 53 35 38 47 53 46 40 36 31 35 32 52 50 49 3 20 26 14 12 0 3 25 29 4 9 13 14 18 8 14 9 14 21 18 Perinopollenites elatoides 3 0 0 0 2 0 0 0 0 0 0 0 2 0 0 0 1 0 0 0 0 0 0 2 0 0 1 0 0 1 0 0 3 19 0 0 0 0 3 3 0 1 2 2 3 1 4 Pilosisporites trichopapillosus 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 3 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Retitriletes austroclavatidites 9 5 2 5 16 4 1 4 5 5 3 8 6 6 1 2 1 12 5 8 2 5 5 10 2 10 11 0 1 1 1 0 7 11 8 11 13 7 12 7 12 6 5 9 7 6 4 Rogalskaisporites cicatricosus 1 0 0 0 1 0 0 0 0 0 0 0 0 0 1 2 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 2 0 0 0 0 0 0 0 0 Ruffordiaspora australiensis 0 0 0 0 0 0 0 1 1 2 0 0 0 1 0 0 0 1 1 0 1 0 3 0 1 0 0 0 0 0 0 0 1 0 1 0 0 0 0 0 0 0 0 0 0 0 1 Rugulatisporites spp. 0 2 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 2 0 2 0 0 0 0 0 0 0 0 0 0 0 Sestrosporites pseudoalveolatus 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Stereisporites antiquasporites 10 12 3 3 14 6 1 3 5 10 5 5 11 7 8 3 8 4 7 20 16 9 13 10 15 20 11 0 2 6 3 0 3 4 6 4 3 4 15 8 16 15 2 10 12 6 8 Todisporites rotundiformis 9 4 7 0 9 6 0 1 3 2 2 1 1 2 1 0 4 6 3 0 2 4 0 3 2 6 10 1 1 0 3 0 7 0 0 4 1 2 1 2 1 0 0 0 1 0 2 Triquitrites spp. 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 1 1 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 0 Undifferentiated bisaccate pollen 49 66 62 28 100 65 54 73 63 23 44 62 64 66 74 73 54 65 43 42 58 39 55 54 37 60 60 4 14 43 11 18 22 14 69 53 65 44 53 43 74 66 89 99 105 53 75 Undulatisporites undulapolus 1 1 2 2 0 0 0 0 0 1 0 0 1 0 0 2 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Unknown 22 20 16 29 28 17 10 14 10 5 6 7 9 7 3 6 11 14 11 14 9 11 3 7 4 15 6 1 6 8 6 9 31 16 26 10 19 26 13 15 4 11 25 18 17 17 15 Unknown Palynomorph A 0 45 4 0 2 6 0 4 5 3 0 0 0 0 0 0 0 7 0 3 1 1 0 0 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0 3 0 0 0 0 0 0 0 Verrucosisporites spp. 1 4 0 0 1 3 0 3 1 0 0 0 3 7 0 0 0 3 0 0 0 0 0 0 0 1 2 0 0 0 0 0 0 0 0 1 2 8 1 1 0 1 0 0 0 0 0

DINOFLAGELLATE CYSTS

264 A p p e n d i x 2 : R a w c o u n t d a t a f o r s p o r e s a n d p o l l e n a n d d i n o f l a g e l l a t e c y s t s f r o m t h e T r e e l e s s C r e e k s e c t i o n . A l l s a m p l e s a r e s t o r e d a t t h e G e o l o g i c a l S u r v e y o f C a n a d a , C a l g a r y O f f i c e . S l i d e i n f o r m a t i o n i n c l u d e s c u r a t i o n ( C - # ) a n d p r e p a r a t i o n ( P - # ) n u m b e r s .

SPORES AND POLLEN P# P5366-48B P5366-49B P5366-50B P5366-51B P5366-52B P5366-53B P5366-54B P5366-55B P5366-56B P5366-57B P5366-58B P5366-59B P5366-60B P5366-61B P5366-62B P5366-63B P5366-64B C# C-604482 C-604483 C-604484 C-604485 C-604486 C-604487 C-604488 C-604489 C-604490 C-604491 C-604492 C-604493 C-604494 C-604495 C-604497 C-604498 C-604499 Thickness (m) 97 95 75 71 67 63 59 57 55 53 51 49 47 45 35 18 16 Acanthrotriletes varispinosus 2 0 0 0 0 0 0 0 0 0 0 2 2 3 0 0 4 Aequitriradites spp. 0 1 13 1 4 0 9 8 7 2 8 5 5 2 1 0 9 Apiculatisporites spp. 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Araucariacites australis 0 0 0 0 0 0 0 0 0 0 0 2 3 4 2 1 24 Baculatisporites comaumensis 3 9 3 5 6 0 3 21 0 10 6 4 4 11 2 2 7 Biretisporites potoniaei 0 0 1 0 0 0 0 0 0 0 0 0 1 0 0 0 0 Camarazonotriletes insignis 0 0 0 0 0 0 0 4 0 1 0 0 2 0 0 0 4 Cerebropollenites mesozoicus 4 18 19 14 7 8 10 7 16 9 5 8 8 6 3 4 1 Chasmatosporites spp. 0 1 0 0 4 0 0 1 0 0 1 0 5 0 2 0 0 Cingulatisporites spp. 1 3 10 6 6 8 9 11 11 9 13 8 1 3 0 0 9 Cingutriletes clavus 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 Classopollis classoides 20 10 10 1 2 1 12 9 8 5 9 7 8 5 1 3 2 Coronatispora valdensis 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 Cupressaceae-Taxaceae pollen 119 57 88 82 96 84 104 78 66 80 58 68 109 119 90 83 73 Cyathidites australis 3 3 10 6 8 5 8 9 7 4 3 6 6 8 2 1 2 Cycadopites spp. 0 1 1 0 2 0 0 0 0 3 2 2 0 1 0 0 1 Deltoidospora hallii 2 2 1 0 1 0 1 0 0 0 5 7 3 0 1 0 3 Densoisporites spp. 20 14 27 14 13 10 7 31 36 29 25 23 26 17 0 0 19 Dictyophyllidites harrisii 0 1 1 0 0 0 0 0 0 0 0 0 0 0 0 0 1 Distaltriangulatisporites spp. 0 0 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 Entylissa spp. 0 2 0 2 0 0 0 0 3 3 0 0 0 1 1 0 0 Eucommiidites troedsonii 0 0 0 0 0 0 0 0 0 3 0 0 0 1 0 0 0 Foveosporites spp. 0 0 0 0 0 0 0 2 1 2 0 0 0 0 0 0 0 Gleicheniidites senonicus 0 2 0 2 0 0 1 3 1 0 3 1 0 0 0 1 1 Granulatisporites spp. 0 0 1 1 2 0 1 0 2 0 0 6 2 1 0 0 19 Ischyosporites spp. 0 0 0 1 0 1 1 0 0 0 0 0 0 0 0 0 0 Kuylisporites lunaris 0 0 0 0 0 0 2 2 3 6 0 1 4 0 2 0 5 Laricoidites magnus 7 18 7 8 13 26 8 5 20 6 12 12 9 15 46 6 9 Leiotriletes directus 0 0 0 0 0 0 11 9 5 5 7 2 7 0 0 0 7 Leptolepidites verrucatus 0 5 1 1 2 0 0 1 0 1 0 5 0 0 0 0 7 Lycopodiumsporites crassimacerius 0 0 0 0 0 0 0 0 0 2 2 0 0 0 1 0 0 Lycopodiumsporites expansus 1 0 0 1 1 0 0 2 0 0 0 0 1 0 0 1 0 Lycospora spp. 13 5 8 6 8 4 1 14 4 5 14 1 5 9 2 0 1 Matonisporites crassiangulatus 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 Neoraistrickia truncata 0 0 0 0 0 0 0 1 0 0 1 0 0 0 0 0 0 Osmundacidites wellmannii 16 45 47 32 22 20 21 16 25 17 39 33 21 37 8 7 64 Perinopollenites elatoides 4 0 2 1 0 1 3 1 3 4 0 4 3 0 0 1 2 Psilatriletes radiatus 0 0 0 0 0 0 0 0 0 0 1 1 0 0 1 0 0 Retitriletes austroclavatidites 1 3 0 0 1 1 1 0 1 0 2 3 0 2 0 1 0 Ruffordiaspora australiensis 0 1 1 1 2 0 0 6 0 0 1 0 0 0 0 0 0 Stereisporites antiquasporites 1 2 0 3 2 0 0 0 0 1 1 4 0 0 2 0 0 Todisporites rotundiformis 3 2 8 7 8 5 3 4 13 18 23 16 11 3 5 0 12 Triquitrites spp. 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 1 Undifferentiated bisaccate pollen 54 64 35 123 64 94 57 33 59 64 49 71 50 48 127 76 17 Unknown 28 31 49 24 46 39 48 32 41 23 32 27 27 32 16 38 24 Unknown Palynomorph A 0 0 0 0 0 0 0 0 0 0 3 0 0 0 0 0 1 Verrucosisporites spp. 0 0 0 0 0 0 2 6 0 0 0 0 2 1 2 2 0

265 DINOFLAGELLATE CYSTS P# P5366-48B P5366-49B P5366-50B P5366-51B P5366-52B P5366-53B P5366-54B P5366-55B P5366-56B P5366-57B P5366-58B P5366-59B P5366-60B P5366-61B P5366-62B P5366-63B P5366-64B C# C-604482 C-604483 C-604484 C-604485 C-604486 C-604487 C-604488 C-604489 C-604490 C-604491 C-604492 C-604493 C-604494 C-604495 C-604497 C-604498 C-604499 Thickness (m) 97 95 75 71 67 63 59 57 55 53 51 49 47 45 35 18 16 Cometodinium spp. 3 0 56 0 5 5 3 3 3 6 2 6 0 7 0 0 0 Epiplosphaera cf. saturnalis 0 0 2 0 0 0 0 0 2 5 0 2 1 0 0 0 1 Gonyaulacysta dualis 0 0 13 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Gonyaulacysta spp. 0 0 0 0 0 0 0 7 0 1 0 1 0 0 0 0 0 Gonyaulacysta ? pectinigera 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 0 0 Hystrichodinium spp. 0 0 0 0 9 1 4 6 0 0 0 0 0 0 0 0 0 Kiokansium cf. unituberculatum 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 Kiokansium williamsii 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 Muderongia tetracantha 0 0 0 0 0 1 3 4 1 2 5 1 1 3 0 0 0 Oligosphaeridium anthophorum 0 0 1 1 5 8 2 0 0 0 3 0 0 0 0 0 0 Oligosphaeridium pulcherrimum 0 0 0 0 1 0 0 0 0 0 0 0 0 1 0 0 1 Oligosphaeridium cf. tenuiprocessum 7 0 30 2 19 52 4 6 0 4 6 7 3 11 0 0 0 Paragonyaulacysta? borealis 1 0 5 0 1 0 0 0 0 2 0 1 2 0 0 0 0 Sirmiodinium grossii 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 1 Spiniferites spp. 3 0 15 0 8 13 1 5 1 5 0 13 2 5 0 0 0 unknown 3 0 15 3 9 14 1 7 2 2 1 5 2 13 0 0 10

266 Appendix 3: RStudio scripts used for quantitative statistical analyses.

>install.packages("vegan") >library(vegan)

#DCA >data_frame<-read.table(file="/filelocation/file_name.csv", header=TRUE, sep=",") >data_frame.dca<-decorana(data_frame) >summary(data_frame.dca) >plot(data_frame.dca)

#For DCA plot aesthetics >plot(data_frame.dca, display=c("none"), cols=c(1,2)) >points(data_frame.dca, display=c("sites"), choices=1:2, pch=3, col="green") >points(data_frame.dca.dca, display=c("species"), choices=1:2, pch=4, cex=0.7) >legend("topleft",inset=.02, legend=c("sites", "species"), col=c("green", "black"), pch=c(3,4), cex=0.8) >text(data_frame.dca.dca, display=c("species"), choices=1:2) >text(data_frame.dca.dca, display=c("sites"), choices=1:2)

#Qmode >data_frame<-read.table(file="/filelocation/file_name.csv", header=TRUE, sep=",") >dist.mat<-dist(data_frame,method="euclidean") >clust.res<-hclust(dist.mat,method="ward.D2") >plot(clust.res,hang=-1)

#Rmode >transposed.frame<-t(data_frame) >dist.mat<-dist(transposed.frame,method="euclidean") >clust.res<-hclust(dist.mat,method="ward.D2") >plot(clust.res,hang=-1)

267

Appendix 4: DCA of Treeless Creek section showing site ordination.

268