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Pliensbachian–Toarcian (Early Jurassic) extinction in western North America Caruthers, Andrew Harry 2013

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Pliensbachian–Toarcian (Early Jurassic) Extinction in Western North America by Andrew Harry Caruthers M.S., The University of Montana, 2005 B.S., The University of Puget Sound, 2002  A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF  DOCTOR OF PHILOSOPHY in THE FACULTY OF GRADUATE STUDIES (Geological Sciences) THE UNIVERSITY OF BRITISH COLUMBIA (Vancouver) April 2013 © Andrew Harry Caruthers, 2013  Abstract The Pliensbachian–Toarcian marine extinction is observable at the species and generic levels. Ammonite diversity data from Europe and parts of the Arctic suggest a multi-phased event with diversity declining over six separate intervals. The main-phase of decline begins at the Pliensbachian–Toarcian boundary and extends into the Early Toarcian to a level correlative with the Tenuicostatum / Serpentinum Zone boundary. To date only this main-phase has been demonstrated as being global in extent, affecting multiple taxonomic groups. The entire Pliensbachian-Toarcian extinction has been attributed to regional and global controlling mechanisms associated with the Volcanic Greenhouse Scenario, an hypothesis linking eruption of the Karoo–Ferrar large igneous province (LIP) to global warming and mass extinction, specifically involving the release of methane hydrate from shelf reservoirs and a global marine anoxic event in the Early Toarcian (the T–OAE). The study presented herein uses paleontology and isotope geochemistry to investigate the duration and potential controlling mechanisms of this protracted extinction. A primary objective is to compare new data from western North America with previously established records in Europe, testing: 1) the multiphased nature of this extinction, 2) its magnitude within two taxonomic groups (ammonoids and foraminifera) in western North America and 3) its controlling mechanisms, relating to methane hydrate release and geographic extent of the T-OAE. Results show that all six phases of species decline are recognizable in western North America, even the oldest episode which was previously thought to be an event restricted to the Tethys Ocean area of Europe. This research strongly supports a correlation between the timing of the entire multi-phased extinction and formation of the Karoo igneous province. The study also provides one of the first records of the Early Toarcian ‘negative carbon-isotope excursion’ outside the Tethys Ocean area (concomitant with the main-phase of extinction) which implicates global methane hydrate release. Lastly, geochemical results do not support the presence of an anoxic water mass in the northeast paleo Pacific Ocean at the time of the so-called global Toarcian event (T-OAE).  ii  Preface Data presented in this study is the result of collaboration between UBC’s paleontology laboratory and the Stable-Isotope Biogeochemistry Laboratory at Durham University (UK), led by Dr. Paul Smith, Dr. Darren Gröcke, and Andrew Caruthers. I was responsible for all of the geochemical sample collection / preparation of all of the samples, and for analyzing the Yakoun River 2010, Whiteaves Bay 2010 and South Barrow #3 datasets (a total of 1243 samples). I was also responsible for compiling the stratigraphic range charts, performing the extinction analysis, and was the primary writer in all published material. Dr. Smith supervised the sample collection, compilation of stratigraphic range charts, extinction analysis, and the writing process. Dr. Gröcke supervised the sample collection and isotope analysis (in Durham). Dr. Gröcke was also responsible for analyzing the Yakoun River 2008 and Whiteaves Bay 2008 datasets (a total of 353 samples). This thesis is composed of published and unpublished material. Published material is spread-out throughout all chapters of this work and does not constitute the entirety of any one chapter. In particular, modified versions of Figures 1.9, 1.11, 2.1, 3.5, 4.1, 4.4–4.6, 5.7, 5.8, 6.3, Plate 1 and the related text within each chapter have been published in the following contributions: Caruthers, Andrew H., Gröcke, D.R., and Smith, P.L., 2011. The significance of an Early Jurassic (Toarcian) carbon-isotope excursion in Haida Gwaii (Queen Charlotte Islands), British Columbia, Canada. Earth and Planetary Science Letters, 307, 19–26. I collected and prepared the samples for isotope geochemistry analysis and wrote the manuscript. Dr. Gröcke supervised sample collection and performed the isotopic analysis of these samples from the Yakoun River section. Dr. Smith supervised the sample collection and the writing process. Caruthers, Andrew H., and Smith, P.L., 2012. Pliensbachian ammonoids from the Talkeetna Mountains (Peninsular Terrane) of Southern Alaska. Review de Paléobiologie Volume spécial 11, 365–378. I collected the ammonites and wrote the manuscript. Dr. Smith supervised the fossil identification and writing process.  iii  Table of Contents Abstract ..................................................................................................................ii Preface .................................................................................................................. iii Table of Contents ..................................................................................................iv List of Figures ........................................................................................................vi List of Plates ..........................................................................................................ix Acknowledgements ............................................................................................... x Dedication ............................................................................................................ xii Chapter 1 Introduction ........................................................................................... 1 1.1 Mass extinctions ........................................................................................ 1 1.2 Mass extinction and the eruption of large igneous provinces.................... 9 1.3 The Pliensbachian–Toarcian extinction event ......................................... 14 1.3.1 Geochemical data from northwest Europe and the Mediterranean Tethys .......................................................................................................... 20 1.3.2 Controlling mechanisms .................................................................... 26 1.4 Study objectives ...................................................................................... 30 Chapter 2 The Pliensbachian–Toarcian time scale ............................................. 32 2.1 Pliensbachian zonal scheme .................................................................. 35 2.2 Toarcian zonal scheme .......................................................................... 40 2.3 Calibration with U–Pb ages .................................................................... 44 2.4 Zonal standards for the high-Arctic ......................................................... 45 2.4.1 Foraminiferal zone scheme for the high Arctic .................................. 47 Chapter 3 Methods: extinction & geochemical analyses..................................... 51 3.1 Extinction analysis .................................................................................. 51 3.1.1 Diversity measurements ................................................................... 61 3.2 Geochemical analysis ............................................................................. 63 Chapter 4 Regional geology ................................................................................ 67 4.1 Haida Gwaii ............................................................................................ 69 4.1.1 Lithostratigraphy ............................................................................... 71 4.2 Northern Alaska ...................................................................................... 75 4.2.1 Lithostratigraphy ............................................................................... 78 4.3 Southern Alaska ..................................................................................... 80 4.3.1 Lithostratigraphy ............................................................................... 81 4.3.1.1 Biostratigraphy of the Hicks and Camp Creek sections ............... 83 4.3.1.1.1 Camp Creek section ............................................................... 85 4.3.1.1.2 Hicks Creek section ............................................................... 89 Chapter 5 Results ............................................................................................... 91 5.1 Extinction data and patterns of diversity ................................................. 91 5.1.1 Extinction and Origination patterns ................................................... 98  iv  5.2 Geochemical data ................................................................................. 101 5.2.1 Whiteaves Bay section ................................................................... 101 5.2.2 Yakoun River section ...................................................................... 107 5.2.2.2 Higher resolution sampling ........................................................ 115 5.2.3 South Barrow #3 core ..................................................................... 117 3.2.3.1 Temporal constraints ................................................................. 120 Chapter 6 Discussion ........................................................................................ 123 6.1 A multi-phased Pliensbachian–Toarcian mass extinction .................... 123 6.1.1 Correlation with the Karoo–Ferrar magmatism ............................... 128 6.2 The negative CIE interval & the long-term carbon-isotope record........ 133 6.2.1 The long-term carbon-isotope record ............................................. 137 6.3 Global vs. regional marine anoxia ........................................................ 147 6.3.1 Total organic carbon ....................................................................... 149 6.3.2 Nitrogen-isotope data ..................................................................... 153 6.3.2.1 Regional denitrification during the Early Pliensbachian ............. 157 Chapter 7 Systematic paleontology .................................................................. 162 Chapter 8 Conclusions ...................................................................................... 207 References ........................................................................................................ 215 Appendices ....................................................................................................... 245 Appendix A: Diversity and Rate Metrics ........................................................ 245 Appendix B: Geochemical Data ..................................................................... 249  	
    v  List of Figures Figure 1.1  Extinction rate and potential causes of major extinction events throughout the Phanerozoic .......................................................... 3 Figure 1.2 The Volcanic Greenhouse Scenario ............................................. 5 Figure 1.3 Continental and oceanic large igneous provinces ...................... 10 Figure 1.4 A comparison in timing of mass extinction events and eruption ages of large igneous provinces ................................................. 12 Figure 1.5 Discrepancy in timing of events related to the Pliensbachian– Toarcian extinction ...................................................................... 14 Figure 1.6 Family-level extinction data for the Pliensbachian–Toarcian extinction event ........................................................................... 16 Figure 1.7 Species-level extinction data for the Pliensbachian–Toarcian extinction event from northwest Europe ...................................... 18 Figure 1.8 Multi-phased extinction data for the Pliensbachian–Toarcian extinction event ........................................................................... 19 Figure 1.9 Negative carbon-isotope excursion data from northwest Europe and parts of the Mediterranean ....................................... 22 Figure 1.10 Correlative chemostratigraphy of northwest Europe and parts of the Mediterranean .................................................................... 25 Figure 1.11 Global paleogeographic map for the Early Toarcian showing locations for localities discussed ................................................ 27 Figure 1.12 Schematic diagram of the restricted basin hypothesis ................ 29 Figure 2.1 Figure 2.2 Figure 2.3 Figure 2.4  Figure 3.1 Figure 3.2 Figure 3.3 Figure 3.4 Figure 3.5 Figure 3.6  Correlative Pliensbachian–Toarcian time scale .......................... 34 Stratigraphic ranges of ammonite species in western North America ....................................................................................... 37 Stratigraphic ranges of ammonite species in western North America ....................................................................................... 39 Stratigraphic ranges of ammonite species in western North America ....................................................................................... 42 Stratigraphic ranges of foraminiferal species in western North America ............................................................................. 52 Stratigraphic ranges of foraminiferal species in western North America ............................................................................. 54 Stratigraphic ranges of foraminiferal species in western North America ............................................................................. 55 Stratigraphic ranges of foraminiferal species in western North America ............................................................................. 56 Map of western North America showing localities in this Study .......................................................................................... 57 Correlation chart of previously published stratigraphic  vi  Figure 3.7 Figure 4.1 Figure 4.2 Figure 4.3 Figure 4.4 Figure 4.5 Figure 4.6  sections ...................................................................................... 58 Parameters and metrics used to analyze extinction ................... 63 Map showing major tectonic elements of the North American Cordillera .......................................................... 68 Pliensbachian–Toarcian stratigraphy of Haida Gwaii ................. 71 Summary of the depositional sequence of the Kingak Shale, Northern Alaska ............................................................... 78 Litho- and biostratigraphy of the Hicks Creek and Camp Creek sections, Talkeetna Mountains Alaska ........... 84 Distribution of Pliensbachian–Toarcian ammonite taxa identified in this study ......................................................... 86, 163 U–Pb concordia diagram for Ash Sample CCA-1, Camp Creek section .............................................................................. 88  Figure 5.1 Figure 5.2  Ammonite and foraminiferal species diversity data .................... 92 Foraminiferal species diversity data, combined and without long-ranging taxa ........................................................... 96 Figure 5.3 Extinction rate data for ammonites and foraminifera without longer stratigraphic ranges ........................................... 100 Figure 5.4 Litho- and biostratigraphy of the Whiteaves Bay section, Haida Gwaii ................................................................. 102 Figure 5.5 Geochemistry of the Whiteaves Bay section, Haida Gwaii .............................................................................. 105 Figure 5.6 Litho- and biostratigraphy of the Yakoun River section, Haida Gwaii ................................................................. 108 Figure 5.7 Geochemistry of the Yakoun River section, Haida Gwaii .............................................................................. 111 Figure 5.8 Seawater 87Sr/86Sr for the Late Pliensbachian– Early Toarcian of Yorkshire (UK) .............................................. 114 Figure 5.9 Higher-resolution carbon-isotope data for the Yakoun River section, Haida Gwaii .......................................... 117 Figure 5.10 Geochemical data for the South Barrow #3 core, northern Alaska ........................................................................ 119 Figure 5.11 Litho- and biostratigraphy for the South Barrow #3 core, northern Alaska ............................................................... 122 Figure 6.1 Figure 6.2 Figure 6.3 Figure 6.4  Correlative multi-phased species level extinction of western North America, Europe and the Arctic ........................ 124 Comparison of timing between the eruptions of the Karoo–Ferrar province and the multi-phased extinction ........... 131 Correlative global negative carbon-isotope excursion data .......................................................................... 135 Late Triassic–Early Jurassic carbon-isotope data  vii  from Haida Gwaii ...................................................................... 139 Figure 6.5 Combined Pliensbachian–Toarcian carbon-isotope data from Haida Gwaii and northern Alaska ............................. 142 Figure 6.6 A comparison of Late Pliensbachian carbon-isotope data between western North America and Portugal ................. 143 Figure 6.7 A composite Late Permian–Middle Triassic carbonisotope record from southern China ......................................... 146 Figure 6.8 Combined Pliensbachian–Toarcian organic carbon data from Haida Gwaii and northern Alaska ............................. 152 Figure 6.9 Combined Early Pliensbachian and Toarcian nitrogenisotope data from Haida Gwaii and northern Alaska ................ 156 Figure 6.10 Adapted flow chart for the Volcanic Greenhouse Scenario showing the potential effects of global warming in restricted vs. unrestricted basins .............................................................. 157 Figure 6.11 Combined Early Pliensbachian and Toarcian total nitrogen data from Haida Gwaii and northern Alaska ............... 160  viii  List of Plates Plate 7.1 Plate 7.2 Plate 7.3 Plate 7.4 Plate 7.5 Plate 7.6 Plate 7.7 Plate 7.8  Late Pliensbachian ammonoids from the Talkeetna Mountains ................................................................................. 189 Toarcian ammonoids from Haida Gwaii .................................... 192 Specimens of the Toarcian genus Cleviceras from the Yakoun River section ................................................................ 195 Toarcian ammonoids from Haida Gwaii .................................... 197 A specimen from the genus Phymatoceras from the Yakoun River section ................................................................ 199 Toarcian ammonoids from Haida Gwaii .................................... 201 A specimen from the genus Rarenodia, Yakoun River section ...................................................................................... 203 Toarcian ammonoids from the Yakoun River section ............... 205  ix  Acknowledgements This research could not have been accomplished without support and guidance from countless individuals. My supervisors, Paul Smith and Darren Gröcke, were instrumental in this undertaking! I would like to express my sincerest gratitude for their unwavering intellectual support, generous funding, and continuous enthusiasm for this research topic. I would also like to thank my committee, Stuart Sutherland, Jim Mortensen and Mike Orchard, for their assistance and support. Stuart and Mike, for the many insightful tips and conversations during my committee meetings and Jim (as well as Rich Friedman at the Pacific Centre for Isotopic and Geochemical Research, UBC), for conducting the U–Pb TIMS isotope analyses. Funding for this work was generously provided by NSERC research grants to Paul Smith (#8493) and Darren Gröcke (#288321), a NERC grant to Darren Gröcke (NE/H021868/1), financial support to the Stable-Isotope Biogeochemistry Laboratory (SIBL) at Durham University was provided by TOTAL UK, a grant from the Alaska Geological Society (to AHC), and tuition support from UBC. I would like to express my sincerest gratitude to those who helped me complete the various portions of this research. In Alaska: Robert B. Blodgett for pointing me in the direction of the Talkeetna Mountains back in 2004, for access to previously collected material from the Camp Creek section, and for his guidance in potential contacts for the South Barrow #3 Core; Curvin Metzler was of critical assistance while in the field, and provided much logistical support as well as access to previously collected material from both sections at Hicks Creek and Camp Creek; Ken Papp, Kurt Johnson and Jean Riordan at the Geologic Materials Center in Eagle River Alaska for allowing me to collect samples from the South Barrow #3 Core, and for their patience!; Ken Bird (USGS) and John Reeder (DNR) provided much assistance in selecting the South Barrow #3 Core, and allowed access to published (and unpublished) material; and Boris Nikitenko generously provided an unpublished version of the foraminiferal zone scheme for the South Barrow #3 core, which greatly aided the level of biostratigraphic constraint for geochemical results. József Pálfy and Giselle Jakobs helped locate the starting point of their sections along the Yakoun River, which ultimately led to the understanding of new exposures in the lower part of the section. Joanne Peterkin at the Stable-Isotope Biogeochemistry Laboratory (SIBL) at Durham University provided analytical assistance during sample analysis. Marc Bustin is thanked for access to his centrifuge and Wayne Vogl (at the Life Sciences Centre, UBC) for ‘ultrapure’ Milli-Q water during sample preparation and George Stanley Jr. at the University of Montana Paleontology Center for allowing the curation of ammonites from the Talkeetna Mountains. This research was also greatly aided by the discussions and suggestions of Jim Haggart, Guillaume Dera, Steve Calvert and Mike Foote who provided much insight toward various parts of the extinction and geochemical analyses. I would also like to thank many people who helped me along the way, either in the field by collecting ammonites (the fun part!) or by collecting samples  x  for geochemistry (the not-so-fun part!); or in lab by keeping my spirits up through their constant enthusiasm and humor. A very special thanks to Luke Beranek, Mike Valentine, Allison Perrigo, Erika Kercher, and Malcolm Brown for assistance while in the field; and to my lab-mates Melissa Grey, Louise Longridge, Emily Hopkin, Farshad Shirohammad, Martyn Golding, Pengfei Hou, and Sarah Porter. Lastly, this work could not have been accomplished without the support, encouragement and love of my girlfriend (Ellie Young), parents (Jennie and Marvin Caruthers), aunt (Sally Smoly ‘the great one’), baba (Mary Smoly), as well as the rest of my family (Jon Caruthers, Alexander Caruthers, Fiona CaruthersJucker, and Amy Nafziger). Thanks for being there for me through the tough times! And finally, how could I forget my furry companions, AnnaBelle, Monkey and cousin Cheryl, who always kept my spirits up by providing a constant source of entertainment and companionship.  xi  Dedication To my parents Jennie and Marvin Caruthers, who were always insistent that I “Do my homework” and constantly urged me to see the bigger picture of life and the potential opportunities ahead. You are my biggest inspiration!  xii  Chapter 1 Introduction 1.1 Mass extinctions Mass extinctions involve the global disruption of a broad suite of environments causing a significant proportion of the world’s biota to disappear over a narrow interval of geological time (Hallam and Wignall, 1997). They force the biosphere to restructure and, along with other background extinctions and recoveries, help form the basis of biostratigraphy (Raup, 1994). Recently, mass extinctions have become an important topic of interest because several studies have indicated that the world is currently experiencing a major decline in biodiversity that may be human-induced (Wilson, 1993; Lawton and May, 1995; Leakey and Lewin, 1995; Pimm et al., 1995; Thomas et al., 2004a,b). Understanding the dynamics, causes and extent of extinctions is a central question for science and society. Throughout the Phanerozoic, five first order and up to 27 second order mass extinction events have been identified (Sepkoski, 1986; Raup, 1994; Hallam and Wignall, 1997). First order events show rapid declines in biodiversity that are measurable at the family, genus and species levels (Raup and Sepkoski, 1982; Sepkoski, 1993; Benton and Storrs, 1994; Jablonski, 1995; Benton, 1995; Hallam and Wignall, 1997). During the five well-known first order events (known as the ‘Big Five’), an estimated 75–95% of existing species are thought to have become extinct (Raup, 1994). These events occur in the Ordovician, Devonian, Permian, Triassic, and Cretaceous Periods (Figure 1.1A) and show extinction  1  intensities between 70-95% among families and 76-95% among genera (Jablonski, 1995; Hallam and Wignall, 1997; Bambach, 2006). Second order events, on the other hand, are less well understood. They are often restricted to a specific environment or biogeographic realm and are characterized by much lower extinction rates (estimated to be 8%, 26%, and 33–53% at the family, genus, and species levels respectively using data for two events discussed by Harries and Little, 1999). Due to the lower extinction intensity at the family level, second order events are often best identifiable at the genus and species levels.  2  Figure 1.1 - A) Extinction rate throughout the Phanerozoic for Paleozoic evolutionary fauna (adapted from Peters, 2008). Paleozoic evolutionary fauna refers to the low-biomass, epibenthic suspension feeding organisms that were dominant during the Paleozoic time period (Peters, 2008). Error bars show 95% confidence. Major ‘Big Five’ mass extinction events are indicated, sequentially, in red. B) Potential cause and calculated extinction intensity for major ‘Big Five’ mass extinctions (adapted from Hallam and Wignall, 1997). O = Ordovician, S = Silurian, D = Devonian, C = Carboniferous, P = Permian, Tr = Triassic, J = Jurassic, K = Cretaceous, Pg = Paleogene, Ng = Neogene.  3  Currently there are a wide variety of causal factors and mechanisms invoked to explain first and second order mass extinction events (Figure 1.1B). Nearly all of these factors are Earth-bound and associated with periods of prolonged environmental change relating to events such as eustatic sea-level change, global warming and/or global cooling (Figure 1.1B). Recently, one mechanism-chain in particular has emerged that links global warming to both first and second order extinction events throughout the Late Paleozoic and Mesozoic (Figure 1.2). It is referred to as the ‘Volcanic Greenhouse Scenario’ by Wignall (2005) and it proposes that the eruptions forming large igneous provinces (LIPs) are responsible for initiating a cascade of events which include global warming, marine anoxia and mass extinction (Pálfy and Smith, 2000; Wignall, 2001; 2005). It is this model that is explored in the current study.  4  Figure 1.2 - The ‘Volcanic Greenhouse Scenario’ a model showing the potential chain of events leading to environmental change and mass extinction (modified from Wignall, 2001; 2005). Note, figure is composite from many individual events, not every isotopic system is necessarily affected. Green bubbles are specific events that are addressed in my study. The volcanic greenhouse scenario model invokes outgassing of volcanogenic CO2 and other greenhouse gasses during LIP eruption (Figure 1.2) caused an escalation of temperatures and prolonged global warming (Pálfy and Smith, 2000; Wignall, 2001). Warmer water temperature is thought to have initiated a series of events that resulted in large-scale environmental change (Pálfy and Smith, 2000; Hesselbo et al., 2000; Wignall, 2001). These events are summarized in Figure 1.2 and include: 1) the catastrophic release of methane  5  hydrate along the continental shelf sediment reservoir, 2) disruption of the thermohaline circulation of the ocean, 3) increased weathering and erosion on the continents and 4) ocean stagnation and marine anoxia. The resulting geological evidence of these large-scale events is a geochemical record that contains sizable perturbations in many isotopic systems (Figure 1.2); these include carbon- (Jenkyns, 1988; 2003; 2010; Holser et al., 1989; Magaritz et al., 1992; Gröcke et al., 1999; Hesselbo et al., 2000; 2002; 2004; Wignall, 2001; Pálfy et al., 2001; Ward et al., 2001; Payne et al., 2004; Williford et al., 2007; Whiteside et al., 2007; 2010; Suan et al., 2008; 2010;), nitrogen- (Jenkyns et al., 2001), molybdenum- (Pearce et al., 2008; McArthur et al., 2008), sulfur- (Gill et al., 2011), strontium- (McArthur et al., 2000) and osmium- (Cohen and Coe, 2002; 2007; Cohen et al., 2004; Kuroda et al., 2010) isotopes. It should be noted that Figure 1.2 is a composite diagram showing the total number of isotopic systems potentially affected by LIP eruption and mass extinction, although not every isotopic system listed in Figure 1.2 is necessarily affected simultaneously in each mass extinction event. The development of regionally (or globally?) extensive marine anoxia may result from four factors that are attributed to warmer water temperatures: 1) warm water holds less dissolved oxygen than cold water and therefore the greenhouse effect caused by increasing CO2, from LIP eruption alone, can contribute directly to anoxic marine conditions (Wignall, 2005), 2) warmer surface and bottom water on continental shelves causes a destabilization in the methane hydrate reservoir, this releases a large amount of methane into the water column and the  6  atmosphere which oxidizes quickly into CO2 and thereby reinforces greenhouse conditions and reduces the concentration of dissolved oxygen in the water column (Hesselbo et al., 2000; Wignall, 2005), 3) warmer water disrupts water current dynamics and oceanic circulation patterns by decreasing the amount of cold and dense surface water in polar regions, this decreases the amount of oxygen that can descend to deep water environments and contributes to marine anoxia (Wignall, 2005), 4) global warming causes an increase in continental weathering and runoff, through increased precipitation, which increases the supply of nutrients into the ocean, this increases productivity and lowers the concentration of dissolved oxygen in the water column through the process of biomass decay (Wignall, 2005). Under normal marine redox conditions degradation (oxidation) of organic matter by bacteria is accomplished through the consumption of oxygen, which produces an abundance of carbon dioxide, nitrate, and phosphate (Froelich et al., 1979). However, as oxygen becomes less abundant or absent from the environment, this process begins to consume nitrate. Nitrate is commonly known as the second-best (to oxygen) electron acceptor within this system (Froelich et al., 1979) and under extreme conditions, after the available nitrates are depleted, sulfate-reducing bacteria continue this process. These bacteria reduce a variety of other compounds and minor elements which include manganese, iron hydroxide, sulfate, and methane fermentation. This produces free hydrogen sulfide in the water column and results in a geochemical record with perturbations in the affected isotopic systems (Figure 1.2).  7  Recently, work has also linked anoxic marine conditions with perturbations (in the geological record) in molybdenum and molybdenum-isotopes (McArthur et al., 2008; Pearce et al., 2008). It is believed that in the presence of free hydrogen sulfide, molybdenum is removed from the anoxic (or euxinic) water column by reduction to thiomolybdate complexes (MoOxS2–4–x) and is sequestered in organic matter and/or pyrite, which is also formed in the water column (Helz et al., 1996; Erickson and Helz, 2000; Vorlicek and Helz, 2002; Tribovillard et al., 2004a,b, 2006, 2008; McArthur et al., 2008). This process of anaerobic decay of biomatter altogether is less efficient than that mediated by aerobic microbial processes, so that more of the settling particulate flux survives transport through the water column, leading to organic enrichment of the resulting sediment (Figure 1.2). These periods of increased organic carbon (wt% TOC) in the sedimentary record have been used to identify periods of suspected ocean anoxia in the geological past (Jenkyns, 1988; Hallam, 1995; Jenkyns and Clayton, 1997; Hallam and Wignall, 1997; Pálfy, 2003; Jenkyns, 2003; 2010). The study presented herein is a critical evaluation of this model. It utilizes aspects of marine invertebrate paleontology as well as carbon- and nitrogenisotope geochemistry to investigate the second order Pliensbachian–Toarcian mass extinction and its correlation with the eruption of the Karoo–Ferrar LIP. It specifically addresses the timing and geographic extent of this correlation with respect to: 1) mass extinction, 2) methane release, and 3) marine anoxia (green bubbles in Figure 1.2).  8  1.2 Mass extinction and the eruption of large igneous provinces Large igneous provinces are massive, large-scale, eruptions of predominantly mafic extrusive and associated intrusive igneous rocks (Figure 1.3). They include continental and oceanic flood basalts, volcanic passive margins, oceanic plateaus, submarine ridges and seamount groups (Coffin and Eldholm, 1994). They consist mainly of subaerial sheet flows that can range in volume from several hundred to several thousand cubic kilometers (Baksi, 1990; Arndt et al. 1993; Coffin and Eldholm, 1993; 1994; Courtillot et al., 1999; Wignall, 2001; Bryan et al., 2010) and occur globally on continents as well as oceanic basins (Figure 1.3A). With the exception of basalt eruptions associated with spreading centers, LIPs are the largest accumulations of mafic material on the Earth’s surface (Coffin and Eldholm, 1994). The formation of LIPs, which are largely thought to be the result of mantle plumes or hotspots (Wilson, 1963; Morgan, 1971; 1972; 1981; Coffin and Eldholm, 1994) produce considerable quantities of CO2 and other greenhouse gasses. As previously discussed, the long-term outgassing during LIP eruptions is thought to cause a significant global climate change (Pálfy and Smith, 2000; Wignall, 2001; 2005).  9  Figure 1.3 - A) Map of the world showing areal extent of various continental and oceanic Large Igneous Provinces (after Wignall, 2001). Karoo Traps include main lavas, dike swarms, and sills (areal extent after Jourdan et al., 2008). Position and areal extent of Ferrar Group after Morrison and Reay (1995). B) List of Large Igneous Provinces with their radiometric age, estimated volume, and corresponding extinction event (compiled from data in Wignall, 2001; Duncan, 2002; Courtillot and Renne, 2003; Jourdan et al., 2008; Wignall et al., 2009).  10  Extensive research over the last two decades has revealed that many mass extinction events and LIP eruptions are approximately coeval (Courtillot, 1999; Pálfy and Smith, 2000; Wignall, 2001; 2005 Courtillot and Renne, 2003; Suan et al., 2008; Caswell et al., 2009). Studies have shown that of the 12 major LIP episodes, 10 coincide with known extinction events (Figure 1.3B, 1.4). Four of these occur consecutively throughout the mid-Phanerozoic and include, in stratigraphic order: 1) the Emeishan flood basalts (~263 Ma; in Sun et al., 2010) and the Middle Permian extinction; 2) the Siberian Traps (~250 Ma) and the end Permian extinction (Wignall, 2001); 3) the Central Atlantic Magmatic Province (CAMP) flood basalt (~200 Ma) and the end Triassic extinction (Wignall, 2001); and 4) the Karoo–Ferrar flood basalt (~183 Ma) and the Pliensbachian–Toarcian extinction (orange circle in Figure 1.4; Pálfy and Smith, 2000). In the case of the Deccan Traps (~65 Ma) and the end Cretaceous extinction, this correlation has been validated (Wignall, 2001; Keller, 2008); however there is strong evidence that this mass extinction event was also heavily influenced by a major bolide impact (Wignall, 2001 and references therein). Other mass extinctions occur at times of low sea level (Figure 1.1B) when loss of shelf space would certainly affect biodiversity. Regression may reflect the eustatic influence of glaciations or tectonics (changes in mid-oceanic ridges) or the more regional, epeirogenic influence of continental doming (followed by fragmentation and the formation of LIPS). However, the interrelationships between regression and other agents of environmental change are beyond the scope of my study.  11  Figure 1.4 - A diagram comparing the timing of mass extinction versus the eruption ages of Large Igneous Provinces over the past 300 million years (after Courtillot, 1999). Note that ages of the Karoo–Ferrar Traps are approximated from data in Jourdan et al. (2008) and that the Pliensbachian–Toarcian extinction event, the subject of this study, is denoted in orange. One of the most intriguing aspects of this correlation is the affect that voluminous outgassing of CO2 during LIP eruption has on the global carbon reservoir. Several studies have shown that during times of LIP eruption, there are large perturbations in the carbon-isotope (δ13C) record, which form the basis for hypotheses concerning environmental change (Pálfy and Smith, 2000; Hesselbo et al., 2000; Wignall, 2001; Jenkyns, 2010). Recently, however, discrepancies have been noted for some of these events concerning the relative timing of extinction, flood basalt eruption, perturbation in the carbon-isotope record, and  12  marine anoxia (Wignall et al., 2005; Wignall and Bond, 2009; Wignall et al., 2009). In some instances it is argued that the main extinction interval precedes either flood basalt eruption or the initial perturbation in carbon isotopes, which causes concern for the validity of a LIP/mass extinction correlation. The Pliensbachian–Toarcian extinction is one such event in which this timing has been questioned (Wignall et al., 2005). In the study by Wignall et al. (2005), a composite dataset was generated from a well-known Early Toarcian succession in northwest Europe showing a clear discrepancy in timing between the extinction interval (with co-occurring anoxic marine facies) and the large negative perturbation in carbon-isotope values (Figure 1.5). Their conclusions have therefore cast some doubt on the timing within the Volcanic Greenhouse Scenario and further motivate this study.  13  Figure 1.5 - Summary of Early Toarcian sedimentation, ichnofacies, benthic macrofauna, and carbon-isotope gechemistry along the Yorkshire coast at Kettleness (UK), modified from Wignall et al. (2005). Figure shows: 1) a good correlation between the extinction horizon of benthic organisms at Kettleness and the appearance of organic-rich shale (indicating anoxic conditions), and 2) a discrepancy in timing between these events and the well-known negative CIE. Dashed red line denotes main extinction of benthic macrofauna (in Wignall et al., 2005) and approximate position of extinction horizon (ii) in Caswell et al. (2009). Diagonal lines indicate the approximate lower limit of anoxic facies, as indicated in Wignall et al. (2005). Serp. = Serpentinum, Exar. = Exaratum, Jet = Jet Rock, HCS = Hummocky cross-stratification.  1.3 The Pliensbachian–Toarcian extinction event The Pliensbachian–Toarcian extinction is best recognized in marine organisms at the family, genus and species levels of northwest Europe  14  (Sepkoski, 1982; 1992; 1996; Hallam, 1986; 1987; Little, 1996; Little and Benton, 1995; Harries and Little, 1999; Caswell et al., 2009). A calculation by Harries and Little (1999, p.45) suggests that this event significantly impacted taxa at the species level in northwest Europe, with a species extinction intensity of 90% among ammonites, belemnites, bivalves (deposit-feeding infaunal, suspensionfeeding infaunal, and semi-infaunal), gastropods, and brachiopods. However, a more recent calculation of ammonite species extinction in the northwest Tethyan and Arctic domains by Dera et al. (2010) suggests this group sustained losses that are measured between 40–65% at the subzone level and 70–90% at the zone level. In Europe, this extinction event also affected a variety of other benthic and pelagic taxonomic groups including: radiolarians, foraminifera, ostracods, dinocysts, crinoids, asteroids, crustaceans, coleoids, marine reptiles, and fish (Hallam, 1986; 1987; Little and Benton, 1995; Hallam and Wignall, 1997; Harries and Little, 1999; Vörös, 2002; Cecca and Macchioni, 2004; Zakharov et al., 2006; Caswell et al., 2009; García Joral et al., 2011).  15  Figure 1.6 - Number of extinct marine families throughout the Early Jurassic (from Little and Benton, 1995). Data show heightened global family-level extinction in five ammonite biozones spanning the Late Pliensbachian to Middle Toarcian. NW = northwest, Eur. = Europe, Teth. = Tethys, Aust. = Austria, Bor. = Boreal, G. = Grammoceras, D. = Dactylioceras, A. = Arnioceras. Family level data show a broad decline in marine biodiversity that spans five ammonite biozones from the Late Pliensbachian to Early Toarcian with extinction peaks in the Late Pliensbachian Margaritatus and Early Toarcian Serpentinum (Falciferum) Zones of northwest Europe (Figure 1.6). Throughout this interval, 33 of 49 global and 18 of 33 northwest European families went extinct (Little and Benton, 1995).  16  At the species and genus level, there is a major decline in biodiversity that occurred within the Early Toarcian at a level correlative with the Tenuicostatum / Serpentinum (Falciferum) Zone boundary (Benton, 1993; Little and Benton, 1995; Harries and Little, 1999; Zakharov et al., 2006; Caswell et al., 2009; Bilotta et al., 2010). This boundary is considered to be the main extinction interval of this event and is thought to have affected many benthic and pelagic marine groups (Figure 1.7). It has been observed in many Tethyan and Boreal regions and has been argued to be global in geographic extent (Zakharov et al., 2006; Caswell et al., 2009; Dera et al., 2010). It is this correlative boundary interval that some argue, coincides with the deposition of organic rich black shale in northwest Europe (Hallam, 1987; Jenkyns, 1988; Little, 1995; Little and Benton, 1995; Jenkyns and Clayton, 1997; Harries and Little, 1999). However the exact timing of extinction at the subzonal level with respect to black shale deposition within northwest Europe (Figure 1.5), and in other outcrops, has been questioned (Wignall et al., 2005; Gomez et al., 2008; McArthur et al., 2008; Garcia Joral et al., 2011).  17  Figure 1.7 - A summary of composite taxonomic ranges of Late Pliensbachian to Middle Toarcian marine species in four localities of northwest Europe (originally compiled by Harries and Little, 1999). Ranges show heightened extinction among marine groups at two main intervals, one at the Pliensbachian / Toarcian boundary and the other within the Early Toarcian at the Tenuicostatum / Serpentinum Zone boundary (denoted as dashed-grey boxes). Intervals of heightened extinction were originally marked by Harries and Little (1999) as four separate steps (dashed line at upper and lower boundary) and were subsequently shown to also occur in the Tethyan and Arctic Realms, steps (i) and (iii) in Caswell et al. (2009). Tenui. = Tenuicostatum, Sp. = Spinatum. Specific identification of marine taxa can be found in Harries and Little (1999). Ammonite zonal scheme after Page (2003). A study by Dera et al. (2010) of Pliensbachian–Toarcian ammonite biodiversity and morphology in the NW Tethyan and Arctic domains has greatly expanded the known temporal extent of this event. In this study a multi-phased event is shown where species and genus level biodiversity declined over six specific intervals of time throughout the Pliensbachian–Toarcian Stages (Figure 1.8). These drops in biodiversity include the: 1) Ibex–Davoei Zone, 2) the  18  Gibbosus subzone, 3) the Pliensbachian / Toarcian boundary, 4) Semicelatum subzone, 5) Bifrons–Variabilis Zones and 6) Dispansum Zone, with the major main-phase extinction event at a level that is correlative with the Tenuicostatum / Serpentinum Zone boundary. In Dera et al.’s work, events 2–6 are interpreted as contributing to the Pliensbachian–Toarcian extinction. Event number 1, occurring in the Early Pliensbachian at the Ibex–Davoei Zone boundary, was originally identified and interpreted by Dommergues et al. (2009) as a regional decline in ammonite species diversity across the Mediterranean and northwest European parts of the Tethys Ocean.  19  Figure 1.8 - Species and generic ammonite diversity in the Pliensbachian and Toarcian of Europe (modified from Dera et al., 2010). Six major declines in biodiversity are recognized with the largest being in the Early Toarcian at the Tenuicostatum / Serpentinum Zone boundary. Major declines in biodiversity are numbered and in red. ‘Weighted richness’ refers to the total number of taxa per measured time interval; ‘geographic singletons’ refer to taxa that are confined to a specific geographic area; ‘single-interval taxa’ are confined to a specific subzone; and ‘boundary crossers’ refer to taxa that occur in two separate time intervals. Marg. = Margaritatus, Sp. = Spinatum, Ten. = Tenuicostatum, Ser. = Serpentinum, Bif. = Bifrons, Variab. = Variabilis, Thouar. = Thouarsense, Dis. = Dispansum, Ps. = Pseudoradiosa, Aal. = Aalensis. To date, this multi-phased extinction has not been established in other taxonomic groups, except for the two phases at the Pliensbachian / Toarcian boundary and the Tenuicostatum / Serpentinum boundary within the Early Toarcian where declines in diversity have been observed in ammonites, bivalves, brachiopods and other previously mentioned taxonomic groups (Harries and Little, 1999; Vörös, 2002; Caswell et al., 2009). Neither has the multi-phased extinction been shown to be global in extent. Except for the decline across the Ibex / Davoei Zone boundary in the Early Pliensbachian, Dera et al. (2010) speculate that the multi-pulsed volcanic activity of the Karoo–Ferrar LIP could have been a trigger for the extinction phases in the Pliensbachian–Toarcian time. However, in order to better assess this control mechanism, it is necessary to first discuss the many changes in isotope geochemistry that also occur.  1.3.1 Geochemical data from northwest Europe and the Mediterranean Tethys Within the epicontinental seaway of NW Europe and the western Tethys Ocean there is a well-documented Early Toarcian disruption in seawater geochemistry that is thought to co-occur with widespread black shale deposition and extinction of marine organisms (Jenkyns, 1988; Jones et al., 1994; Hesselbo 20  et al., 2000; McArthur et al., 2000; Jenkyns et al., 2001; Jones and Jenkyns, 2001; Jenkyns et al., 2002; Cohen et al., 2004; McElwain et al., 2005; Kemp et al., 2005; Caswell et al., 2009). Carbon-isotope data (Figure 1.9) reveal a small negative shift at the Pliensbachian / Toarcian boundary (Hesselbo et al., 2007; Suan et al., 2008; Littler et al., 2010; Bodin et al., 2010) that is subsequently followed by a broad positive shift that extends into the Serpentinum Zone. It is then interrupted at the Tenuicostatum/Serpentinum Zone boundary by a pronounced negative δ13C excursion of ~5–7‰ (Hesselbo et al., 2000, 2007; Kemp et al., 2005; Suan et al., 2008; Sabatino et al., 2009; Hermoso et al., 2009a).  21  Figure 1.9 - Late Pliensbachian–Early Toarcian carbon-isotope records from northwest Europe and the Mediterranean illustrating correlative negative excursions at the Pliensbachian / Toarcian boundary and within the Early Toarcian at the correlative Tenuicostatum / Serpentinum Zone boundary. Figure is adapted from a compilation in Caruthers et al. (2011). Locality numbers at top refer to paleogeographic location in Figure 1.11. This negative CIE has been recorded in many geologic materials including: bulk organic carbon (Hesselbo et al., 2000; Kemp et al., 2005; Sabatino et al., 2009), carbonate (Hesselbo et al., 2007, Hermoso et al., 2009a; Sabatino et al., 2009; Bodin et al., 2010), phytoplanktonic organic matter and organic biomarkers (Schouten et al., 2000; van Breugel et al., 2006), benthic  22  shells (Suan et al., 2008, 2010), fossil wood and phytoclasts (Hesselbo et al., 2000; McElwain et al., 2005; Hesselbo et al., 2007; Hesselbo and Pienkowski, 2011) suggesting a close relationship between the carbon-isotope composition of the carbon reservoirs in the ocean and atmosphere. At the onset of this present study, the geographic extent of this large negative excursion in carbon-isotopes was challenged. In a study of δ13C from belemnite calcite in Early Toarcian sedimentary successions in England (Yorkshire) and Germany, van de Schootbrugge et al., (2005) were unable to identify this negative CIE and thereby challenged the geographic extent of this excursion by suggesting it is a regional phenomena that was principally derived from the upwelling of 12C enriched bottom water into a saline-stratified body of water within the epicontinental seaway of northwest Europe. However, this current research and recent contributions by other groups has shown that the negative CIE is recorded in many coeval Early Toarcian successions that are geographically outside the Tethys Ocean area (Al-Suwaidi et al., 2010, Caruthers et al., 2011; Suan et al., 2011; Gröcke et al., 2011; Izumi et al., 2012). This new evidence of an Early Toarcian negative CIE that is outside the Tethys Ocean area strongly indicates a globally extensive perturbation to the carbon reservoir requiring a global control mechanism. In the epicontinental seaway of NW Europe and parts of the Mediterranean Tethys, the negative CIE coincides with dramatic increases in organic carbon burial (Jenkyns, 1988; Jenkyns and Clayton, 1997; Hesselbo et al., 2000) that are also correlative with perturbations in many geochemical  23  systems (Figure 1.10). These include: positive shifts in nitrogen-isotopes (Jenkyns et al., 2001), osmium-isotopes (Cohen et al., 2004), manganese concentrations (Sabatino et al., 2011), sulfur-isotopes (Gill et al., 2011) and sulfur concentrations (McArthur et al., 2008), turning points in strontium-isotope ratios (McArthur et al., 2000), and a minimum in molybdenum (Mo ppm, Mo/total organic carbon or TOC, and δ98Mo) values. In northwest Europe, the minimum in molybdenum values is subsequently followed by a positive shift in Mo ppm, Mo/TOC, and δ98Mo that coincides with a positive shift in δ13C in the later part of the Early Toarcian (Pearce et al., 2008; McArthur et al., 2008). Furthermore, a recent comparison of carbon- and oxygen-isotope ratios in the Mediterranean Tethys region shows that there was a sharp negative shift in δ18Obrachiopod values at the Pliensbachian / Toarcian boundary and a subsequent broad decrease during the Early Toarcian negative CIE interval (Suan et al., 2008, 2010). As previously explained by the Volcanic Greenhouse Scenario (Figure 1.2), this nearly simultaneous disruption in seawater geochemistry is compelling evidence for climate change and ocean anoxia in this region (Jenkyns et al., 2001; Jenkyns, 2003; 2010; McArthur et al., 2008).  24  25  Figure 1.10 - Correlative chemostratigraphy throughout the Late Pliensbachian Middle Toarcian successions of the Cleveland and Cardigan Bay Basins (northwest Europe) showing geochemical profiles and concentrations of (from left to right): organic carbon, molybdenum, sulfur, nitrogen-isotopes, osmiumisotopes, carbon-isotopes, and strontium-isotope values (modified from McArthur et al., 2008). Spin. = Spinatum, Tenui. = Tenuicostatum, Cleveland. = Clevelandicum. Locality numbers at top refer to paleogeographic location in Figure 1.11. Note that temporal correlation of manganese concentrations from Monte Mangart (Italy / Slovenia border) was not given at the zone-level in Sabatino et al. (2011). Black arrow denotes extent of Jet Rock.  1.3.2 Controlling mechanisms Currently there are two leading hypotheses as to the cause of the CIE in the Early Toarcian, invoking global controls in one case and regional controls in the other. The global hypothesis is synonymous with the Volcanic Greenhouse Scenario (Figure 1.2). It is argued that in the Early Toarcian, rift-related tectonic activity during the break-up of Gondwana initiated the eruption of the Karoo– Ferrar LIP (Figure 1.11). Outgassing of volcanogenic CO2 from this eruption started a prolonged period of global warming (Pálfy and Smith, 2000) that subsequently destabilized and released ~5000 Gt of methane hydrate from continental shelf sediments (Hesselbo et al., 2000; Beerling et al., 2002). The rapid injection and oxidation of methane reduced oceanic O2 levels and increased organic carbon burial, creating euxinic conditions (known as the Toarcian Ocean Anoxic Event or T–OAE). The T–OAE is largely evidenced by correlative perturbations in seawater geochemistry including: organic carbon, nitrogen, sulfur, molybdenum, and manganese (Figure 1.2; Jenkyns, 2010 and references therein; Gill et al., 2011). This severe environmental perturbation is further thought to have: 1) increased seawater temperatures by ~7–10˚C, as  26  evidenced by a sudden decrease in brachiopod oxygen-isotope ratios (Suan et al., 2010); 2) caused a significant increase in continental weathering rates, evident from increases in strontium- and osmium-isotope ratios (Jones and Jenkyns, 2001; McArthur et al., 2000; Cohen et al., 2004); 3) caused a major reduction in biocalcification (Mattioli et al., 2004; Tremolada et al., 2005; Suan et al., 2010); and 4) escalated the extinction intensity of marine organisms at the Tenuicostatum / Serpentinum ammonite Zone boundary (see Caswell et al., 2009 and references therein; Bilotta et al., 2010; Dera et al., 2010). Astronomical precession may have also helped to reinforce the effects of long-term global warming, resulting in the cyclic (at least three separate pulses) release of methane hydrate lasting an estimated ~60–350 kyr per cycle (Kemp et al., 2005; Cohen et al., 2007; Hesselbo and Pienkowski, 2011; Kemp et al., 2011).  27  Figure 1.11 - A and B) Global paleogeographic maps of the Early Toarcian depicting the location of the Tethys and Paleo–Pacific Oceans, the Karoo–Ferrar Large Igneous Province (LIP), and location of various localities in northwest Europe (after Hesselbo et al., 2007; Caswell et al., 2009), the Mediterranean (after Sabatino et al., 2010), South America (after Al-Suwaidi et al., 2010) and this study. 1 = Yorkshire (UK), 2 = Peniche (Portugal), 3 = Mochras (Wales), 4 = Sancerre (France), 5 = Monte Mangart (Italian–Slovenian border), 6 = Neuquén Basin (Argentina), 7 = Wrangellia (Smith, 2006), 8 = northern Alaska (approximate position based on Smith, 2006; Imlay, 1981). As discussed earlier, the timing and geographic extent of the T–OAE has been disputed. Questions have been raised concerning: 1) differences in the timing of black shale deposition, mass extinction, carbon-isotope perturbation, and sea level change between northern (Boreal) and southern (Tethyan) Europe (Figure 1.5; Wignall et al., 2005; Bilotta et al., 2010); 2) the geographic extent and concentration of organic carbon in what is regarded as ‘globally distributed black shale units’, inferred to have been deposited in euxinic conditions (McArthur et al., 2008); and 3) the reported geographic extent for a disruption in molybdenum (Mo ppm) and molybdenum-isotope ratios that are thought to be an indication of euxinic conditions in Early Toarcian marine sediments of NW Europe (Pearce et al., 2008; McArthur et al., 2008). These questions have prompted an alternative hypothesis suggesting the T–OAE may have been caused by local controls and was therefore restricted to the epicontinental seaway of NW Europe (McArthur et al., 2008). This hypothesis proposes that the European epicontinental seaway was the site of several silled basins (Figure 1.12) where black shales, 12C-enriched bottom waters and the previously discussed perturbations in isotope geochemistry formed in response to euxinic conditions brought on by a salinity-driven pycnocline that developed as  28  a result of increased fresh water input (Küspert, 1982; Jenkyns, 1988; Sælen et al., 1996; Schouten et al., 2000; McArthur et al., 2008). In this model, the negative CIE occurred by diachronous upwelling of 12C-rich bottom water in separate marine basins in the Tethys Ocean area and was therefore restricted geographically.  Figure 1.12 - Schematic diagram showing environmental dis-tress, saline stratification, the development of anoxic (or euxinic) redox conditions, and subsequent deposition of organic rich black shale in restricted basins of northwest Europe during the Early Toarcian (modified from McArthur et al., 2008). It is possible that both hypotheses are applicable and that regional and global controls could have occurred simultaneously thereby reinforcing each other. As mentioned, at the onset of this research the negative CIE was unknown outside the area of the Tethys Ocean. The first discovery of this excursion in carbon-isotopes (outside the Tethys Ocean) came from the Neuquén Basin in  29  Argentina, where it was described from two sections (Al-Suwaidi et al., 2010; Mazzini et al., 2010). Its presence was interpreted as evidence that regional controls were not operating alone. However, these data are somewhat problematic. Al-Suwaidi et al. (2010) demonstrate a 6‰ negative shift in δ13Corg and δ13Cwood that is truncated by slumping and erosion at the top of the section so that only the lower part of the Early Toarcian CIE is evident. The data presented by Mazzini et al. (2010) are poorly constrained biostratigraphically but show an 8‰ negative excursion in δ13Corg that, in contrast to conclusions drawn in Al-Suwaidi et al. (2010), seems to occur below the correlative of the Tenuicostatum / Serpentinum Zone boundary. If correct, this would suggest a diachronous, rather than isochronous, negative CIE when compared to European data. Although geographically far-removed from the Tethys Ocean, the Neuquén Basin was similarly somewhat restricted and therefore could have been subjected to coeval regional and global controls as postulated for Europe. It is therefore necessary to seek correlative and preferably more complete data from sediments that were deposited in a marine environment outside the Tethys Ocean area that was not part of a restricted basin.  1.4 Study objectives The study presented herein forms a critical evaluation of the Pliensbachian–Toarcian extinction and its hypothesized control mechanisms. By examining the calibrated biostratigraphy and chemostratigraphy in western North America and comparing it with correlative data in Europe, it will assess: 1) the geographic extent of the multi-phased ammonite extinction recorded by Dera et 30  al. (2010); 2) if other taxonomic groups, such as foraminifera, showed similar declines in biodiversity that are correlative with those observed in ammonites; 3) the geographic extent of the well-documented European perturbation in carbonisotopes (negative CIE); and 4) the long-term (Pliensbachian–Toarcian) carbonand nitrogen-isotope records that are related to a potentially global T–OAE. Carbon-isotope chemostratigraphy in this research is generated from two temporally constrained (to the zonal level) and well-studied areas of western North America, representing deposition along terrane and craton margins. A primary goal is to calibrate the stratigraphic record of biodiversity changes (from the Late Pliensbachian – Late Toarcian strata) with a newly generated North American carbon isotope curve for the Pliensbachian–Toarcian interval. Results will be compared with established records in the Tethys Ocean in order to assess this extinction and its controlling mechanisms on a global scale.  31  Chapter 2 The Pliensbachian–Toarcian time scale In order to compare paleontological and geochemical signals of western North America with the well-established records from Europe and parts of the Mediterranean, it is imperative that coeval strata be identified. Therefore, it is important to establish a geologic time scale for the Pliensbachian and Toarcian that can be applied globally. Jurassic ammonites have been utilized for this purpose and are commonly held as one of the most useful index fossils for zone and subzone level temporal constraint. Throughout the Early Jurassic, ammonites are known to have evolved rapidly producing a broad variety of shell morphologies. Their pelagic mode of life also facilitates a wide geographic dispersal which included both deep- and shallow-water marine habitats. Ammonites of the Jurassic have had a long history of research in Europe dating back to the mid 1800’s with the work of Quenstedt (1845–1849; 1856– 1858; 1882–1885) and Oppel (1853; 1862). These studies laid the framework for a comprehensive Early Jurassic ammonite zonal scheme that was applicable to many successions throughout northwest Europe (Dean et al., 1961). Since this scheme was established, it has undergone refinement (Mouterde et al., 1971; Schlatter, 1980; Howarth, 1992; Page, 2003) and is currently used as a global reference standard with which to compare and calibrate other schemes (Braga et al., 1982; Hillebrandt, 1987; Smith et al., 1988; Jakobs et al., 1994; Zakharov et al., 1997; Page, 2003). 32  In western North America, Early Jurassic interbedded fossiliferous marine sedimentary and volcaniclastic successions are extremely useful for calibrating biochronologic and geochronologic time scales. Previous studies have established an ammonite-based zonal scheme for Pliensbachian and Toarcian successions (Smith et al., 1988; Jakobs et al., 1994; Smith and Tipper, 1996; Jakobs, 1997) that is calibrated with absolute age determinations from interbedded zircon-bearing ash beds (Pálfy et al., 1997; 2000). This calibrated time scale has been further correlated with the ammonite zonal schemes of NW Europe, parts of the Mediterranean, and South America (Figure 2.1). This zonal scheme integrates associations of ammonite species that have similar, overlapping, stratigraphic ranges and provides the basis for zonal boundaries (Smith et al., 1988; Jakobs et al., 1994; Smith and Tipper, 1996). Within this scheme, the stratigraphic range of a particular index (or characterizing, or name bearing) species is not necessarily confined to a zone and, furthermore, its presence is not always necessary in order to recognize the zone (Smith et al., 1988). The study presented herein uses the ammonite based zone schemes of Smith et al. (1988) and Jakobs et al. (1994) to correlate paleontological and geochemical results from western North America with well established coeval records from northwest Europe and parts of the Mediterranean.  33  Figure 2.1 - Correlative Pliensbachian–Toarcian time scales for Northwest Europe and the Mediterranean (Dean et al., 1961; Schlatter, 1980; Braga et al., 1982; Howarth, 1992; Page, 2003); High-Arctic (Zakharov et al., 1997; Nikitenko et al., 2008); western North America (Smith et al., 1988; Jakobs et al., 1994) and South America (von Hillebrandt, 2006). Absolute U–Pb and Ar–Ar age data from Pálfy et al. (1997; 2000). Absolute age dates in bold-face font have an error range that is less than 5 Ma and are interpreted as good quality. Note that the Pliensbachian zone level biostratigraphy in the Mediterranean Province refers to areas of Spain (Page, 2003).  34  2.1 Pliensbachian zonal scheme The Pliensbachian of western North America is divided into five separate zones that include, from the earliest to latest, the Imlayi, Whiteavesi, Freboldi, Kunae, and Carlottense zones (Figure 2.1). The Early Pliensbachian includes the Imlayi, Whiteavesi, and Freboldi Zones and is derived primarily on the stratigraphic ranges of the families Polymorphitidae and Eoderoceratidae (Smith et al., 1988). The combined duration of the chronozones is just short of being equivalent to the combined time frame of the Jamesoni, Ibex and Davoei Zones of northwest Europe. The later part of the Pliensbachian includes the Kunae and Carlottense Zones and, at present, its definition is based primarily on the gradual evolution (and stratigraphic ranges) of two ammonite superfamilies: the Hildoceratoidea and Eoderoceratoidea (Meister, 2010). However, the occurrence of these superfamilies are restricted to specific paleogeographic domains and therefore show limited overlap. The transition between the Sinemurian and Pliensbachian remains poorly defined. Smith et al. (1988) recognize that throughout much of western North America, the definition of the boundary is hampered by unfossiliferous intervals that separate ammonite species of Sinemurian affinity from those of the Pliensbachian. A more recent study, however, has identified a somewhat transitional ammonite fauna from areas of Haida Gwaii (meaning “islands of the Haida people” and formerly the Queen Charlotte Islands), identified as the Tetraspidoceras Assemblage (Pálfy et al., 1994). This so-called Tetraspidoceras Assemblage contains a mix of ammonites that are known from the uppermost Sinemurian and lowermost Pliensbachian. The presence of this assemblage may  35  be an indication that the base of the Imlayi Zone is not the base of the Pliensbachian. However, in western North America the Tetraspidoceras Assemblage has also not yet been demonstrated to occur outside Haida Gwaii and therefore it was not included in Figure 2.1. In the present study, successions yielding ammonite species from the Tetraspidoceras Assemblage were not observed. The Imlayi Zone is characterized by the zonal index species Pseudoskirroceras imlayi and a variety of other taxa (Figure 2.2). The Whiteavesi Zone is marked by the disappearance of Pseudoskirroceras imlayi and the appearance of Tropidoceras masseanum (Smith et al., 1988). Also within this zone, Acanthopleuroceratids such as the zonal index A. whiteavesi occur in abundance. The overlying Freboldi Zone is marked by the appearance of Dubariceras freboldi above the last acanthopleuroceratid (Smith et al., 1988). D. freboldi extends stratigraphically throughout the zone and is described as being widely distributed and endemic to the East Pacific (Frebold, 1970; Imlay, 1968; 1981; Smith, 1983; Dommergues et al., 1984; Smith et al., 1988).  36  Figure 2.2–2.4 - Pliensbachian and Toarcian ammonite biostratigraphy of western North America compiled from previously published data of Imlay (1955; 1981), Smith et al. (1988), Thomson and Smith (1992), Jakobs (1992; 1995; 1997), Jakobs et al. (1994), Smith and Tipper (1996), Jakobs and Smith (1996), Smith et al. (2001), and Caruthers and Smith (2012).  37  In western North America the first appearance of the genus Fanninoceras, primarily above the species Dubariceras freboldi is used to define the base of the Kunae Zone (Smith et al., 1988). In the uppermost Freboldi Zone there is a brief overlap in the stratigraphic ranges of D. freboldi and F. bodegae (Figure 2.2, 2.3) and therefore occurrences of Fanninoceras above this interval of overlap are used to define the Kunae Zone. Although the Kunae Zone lies predominantly within the Late Pliensbachian, its base is situated within the uppermost part of the Early Pliensbachian (Figure 2.1). This zone is characterized further by a variety of species from numerous genera (Figures 2.2, 2.3). The Carlottense Zone is the uppermost zone of the Pliensbachian. Its base is marked by the first appearance of the index species Fanninoceras (Fanninoceras) carlottense (Smith et al., 1988) and is further characterized by various other species of Lioceratoides (Lioceratoides), Protogrammoceras, Amaltheus, Fieldingiceras, Lioceratoides (Pacificeras), and Tiltoniceras (Figure 2.3). In western North America the species Lioceratoides (Pacificeras) propinquum, Protogrammoceras paltum, and Tiltoniceras antiquum occur well below the first appearance of Dactylioceras, which is known globally as the earliest Toarcian ammonite genus (Smith et al., 1988; Smith and Tipper, 1996). However, in Europe and other parts of the world the first appearance of L. propinquum, P. paltum, and T. antiquum parallel the first appearance of Dactylioceras and therefore suggest a younger age. This could suggest that these species evolved in the paleo Pacific Ocean in the latest Pliensbachian and then radiated to other parts of the world by the Early Toarcian.  38  39  2.2 Toarcian zonal scheme In western North America the Toarcian is divided into five ammonite zones (Jakobs et al., 1994; Jakobs, 1997). In ascending stratigraphic order these include the Kanense, Planulata, Crassicosta, Hillebrandti, and Yakounensis Zones (Figure 2.1). Within this scheme, the Kanense Zone comprises the entire Early Toarcian time period and is equivalent to the combined time frame of both the Tenuicostatum and Serpentinum Zones of northwest Europe. The Middle Toarcian contains the Planulata and Crassicosta Zones and correlates with the Bifrons and Variabilis Zones of northwest Europe, and the Late Toarcian contains the Hillebrandti and Yakounensis Zones and correlates with the combined time frame of the Thouarsense, Dispansum, Pseudoradiosa, and Aalensis Zones of northwest Europe (Figure 2.1). As previously discussed, in northwest Europe and in parts of the Mediterranean, the main-phase of the Pliensbachian–Toarcian mass extinction and the accompanying large negative CIE occurs at the Tenuicostatum / Serpentinum Zone boundary within the Early Toarcian (Figures 1.7; 1.9). Also mentioned, in western North America, the Kanense Zone is approximately equivalent to the combined Tenuicostatum and Serpentinum Zones of northwest Europe (Figure 2.1; Jakobs et al., 1994; Jakobs, 1997). Therefore, if the Early Toarcian mass extinction and the negative CIE are isochronous global events, they should occur within the Kanense Zone and would therefore indicate a stratigraphic position that is approximately coeval with the Tenuicostatum / Serpentinum Zone boundary.  40  The base of the Kanense Zone also marks the base of the Toarcian. It is characterized by the first appearance of Dactylioceras above the last occurrences of Amaltheus and Fanninoceras species (Smith et al., 1988; Jakobs et al., 1994; Jakobs, 1997). The Kanense Zone also has several other characterizing species from the genera Dactylioceras, Taffertia, Hildaites, Harpoceras, and Cleviceras (Figure 2.3). Of these, D. kanense, D. aff. comptum, D. cf. compactum, D. alpestre, and T. taffertensis are restricted to the lower part of the zone, whereas Hildaites spp. Hildaites cf. serpentinum, Hildaites murleyi, and Cleviceras cf. chrysanthemum are known from its upper portion (Figures 2.3, 2.4).  41  42  The Middle Toarcian Planulata Zone is marked by the first appearance of Rarenodia planulata above the last occurrence of Hildaites murleyi (Jakobs et al., 1994; Jakobs, 1997). Rarenodia planulata is the index species of the Planulata Zone and unlike D. kanense whose taxonomic range is restricted to the lower part of the underlying Kanense Zone, the taxonomic range of R. planulata is known to extend throughout the zone (Figure 2.4). The base of the Crassicosta Zone is defined by the first appearance of Phymatoceras crassicosta and Peronoceras cf. moerickei above the last Rarenodia planulata (Jakobs et al., 1994). Phymatoceras crassicosta, the zonal index, extends throughout much of the zone and disappears near its top (Figure 2.4). The Late Toarcian Hillebrandti Zone is marked by the appearance of Grammoceras thouarsense, Podagrosites latescens, and the index species Phymatoceras hillebrandti above the last occurrence of P. crassicosta (Jakobs et al., 1994; Jakobs and Smith, 1996). This zone is regarded as having one of the lowest species-level diversities in the Toarcian (Jakobs and Smith, 1996), with the majority of its fauna confined to the zone. The Yakounensis Zone is the uppermost Toarcian ammonite zone in western North America. Its base is defined by the first appearance of Yakounia silvae, Pleydellia maudensis, Hammatoceras speciosum, and Pleydellia spp. above the last occurrence of Phymatoceras hillebrandti (Jakobs et al., 1994). In general, ammonite species of the Yakounensis Zone are largely confined to the zone (Figure 2.4). Species include the index Yakounia yakounensis and a host of other characterizing species (Jakobs et al., 1994). To date, only certain species belonging to the  43  family hammatoceratinae (in Jakobs et al., 1994; Jakobs, 1997) and possibly Holcophylloceras calypso (in Jakobs and Smith, 1996) appear to extend into the Aalenian Stage of the Middle Jurassic.  2.3 Calibration with U–Pb ages The Jurassic time scale has been calibrated using a database of U–Pb and 40Ar/39Ar ages to integrate bio- and geochronologic time scales in many successions throughout western North America, providing age estimates for stage and zonal boundaries throughout this period of time (Pálfy et al., 2000). Two methods were largely employed: 1) A direct dating method, where the age of a biochronologically defined boundary is determined through isotopic dating of a particular volcanogenic layer that is in close proximity to the boundary itself, and 2) the chronogram method, where the age of the boundary in question is estimated using age determinations from adjacent units (discussed in detail in Pálfy et al., 2000). Age determinations are considered ‘good quality’ if they contained a 2σ error range that is less than 5 Ma (age determinations in bold font in Figure 2.1). Although the direct dating method is preferred, within the Early Jurassic time scale it could only be used to date the Triassic–Jurassic boundary and the initial boundary of the Crassicosta Zone of the Middle Toarcian (Pálfy et al., 2000). U–Pb and 40Ar/39Ar age estimates for the Pliensbachian and Toarcian are as follows. The base of the Pliensbachian is poorly constrained at 190.7 +2.7–3.9 Ma; the base of the Freboldi Zone is constrained at 186.7 +1.8–1.6 Ma; the base of the Kunae Zone is constrained at 185.7 +0.5–0.6 Ma; the base of the Carlottense  44  Zone is constrained at 184.1 +1.2–1.6 Ma; the base of the Toarcian (base of the Kanense Zone) is constrained at 183.6 +1.7–1.1 Ma; the base of the Planulata Zone is poorly constrained at 182 +3.3–4.9 Ma; the base of the Crassicosta is constrained at 181.4 +1.2–1.2 Ma (using data from Pálfy et al., 1997); the base of the Yakounensis Zone is constrained at 180.1 +0.7–3.0 Ma; and the base of the Aalenian is constrained at 177.6 +1.4–1.1 Ma (Figure 2.1). This data indicates that the Early Pliensbachian lasted for ~5.8 Ma, the later part of the Pliensbachian lasted for ~2.1 Ma, the Early Toarcian lasted for ~1.6 Ma, and the Middle and Late Toarcian combined lasted for ~4.4 Ma.  2.4 Zonal standards for the high-Arctic Early Jurassic fossiliferous marine successions are widespread and well studied throughout much of northeast Russia, Siberia, Barents Sea, and northern Alaska (Nikitenko 1994; 2008; Zakharov et al., 1997; Shurygin et al., 2000; Nikitenko and Mickey, 2004; Nikitenko et al., 2008; Shurygin et al., 2011). To date, ammonite biostratigraphy within these fossiliferous successions provides the basis for a Pliensbachian and Toarcian Zonal scheme for the Arctic that is applicable to many circum Polar Regions (Zakharov et al., 1997). The Pliensbachian portion of this zone scheme is composed of four ammonite-based zones that include, in ascending stratigraphic order, the Polymorphites, Stokesi, Margaritatus, and Viligaensis Zones (Figure 2.1). In this scheme the Early Pliensbachian consists only of the Polymorphites Zone and a gap interval for which zone-level constraint is currently unavailable. The basal Pliensbachian Polymorphites Zone overlies the Upper Sinemurian Colymicum  45  Zone; its base is consistent with the base of the Pliensbachian (Zakharov et al., 1997). Although the top of the Polymorphites Zone is currently not known due to the overlying gap interval, it has been illustrated as correlative with the Jamesoni Zone of northwest Europe (Zakharov et al., 1997). The Late Pliensbachian part of the zone scheme consists of the Stokesi, Margaritatus, and Viligaensis Zones of which the Stokesi Zone is temporally equivalent with the Stokesi Subzone of northwest Europe (Figure 2.1) and the top of the overlying Margaritatus Zone is consistent with that of northwest Europe (Zakharov et al., 1997). The Viligaensis Zone is the uppermost zone of the Late Pliensbachian and it is considered to be correlative with the Spinatum Zone of northwest Europe (Figure 2.1). The Toarcian part of this circum Polar zonal scheme is somewhat different than in European schemes, in that the Middle Toarcian is not formally recognized (as illustrated in Zakharov et al., 1997; Nikitenko and Mickey, 2004; Zakharov et al., 2006; Nikitenko et al., 2008). This scheme consists of eight zones that are subdivided into the Early and Late Toarcian (Zakharov et al., 1997; Nikitenko et al., 2008). These include, in ascending stratigraphic order, the Antiquum, Falciferum, Commune, Monestieri, and Spinatum Zones of the Early Toarcian and the Compactile, Wuerttenbergeri, and Falcodiscus Zones of the Late Toarcian (Figure 2.1). The basal zone of the Toarcian was originally identified as the Propinquum Zone in Zakharov et al. (1997), however a more recent work by Nikitenko et al. (2008) re-names this zone as the Antiquum Zone; its base is consistent with the base of the Toarcian (Figure 2.1). Within this Early Toarcian  46  zone scheme, the Antiquum and overlying Falciferum Zones are correlative with the Early Toarcian Tenuicostatum and Falciferum Zones of northwest Europe. The overlying Commune, Monestieri, and Spinatum Zones together are temporally equivalent to the Middle Toarcian Bifrons Zone of northwest Europe (Zakharov et al., 1997; Nikitenko et al., 2008). In the Late Toarcian part of the zone scheme, the Compactile Zone correlates with the Middle Toarcian Variabilis Zone of northwest Europe and the Wuerttenbergeri and Falcodiscus Zones are temporally equivalent to the Late Toarcian Thouarsense, Dispansum, Pseudoradiosa, and Aalensis Zones of northwest Europe (Figure 2.1).  2.4.1 Foraminiferal zone scheme for the high Arctic More recent studies have constructed separate zonal schemes that are specific to high Arctic regions and are based on stratigraphic ranges of a variety of macro and microfossils, such as bivalves, foraminifera, and ostracodes (Nikitenko and Mickey, 2004; Nikitenko et al., 2008; Shurygin et al., 2011). The research presented here utilizes the foraminifera-based zone scheme (illustrated in Figure 2.1) in conjunction with ammonite biostratigraphy, to temporally constrain well core data from northern Alaska. This foraminifera-based zone scheme is based, in part, on previously identified specimens that were recovered from various well cores in the National Petroleum Reserve of northern Alaska (Tappan, 1955; Bergquist, 1966). The foraminiferal zone scheme is applicable to the Early–Middle (Aalenian only) Jurassic of the circum Polar Region and is composed of 17 zones that are based on assemblages of foraminifera that have similar overlapping stratigraphic ranges (Nikitenko and Mickey, 2004; Nikitenko et al., 2008). Each zonal assemblage is identified as a JF# (e.g. JF9) with zonal 47  boundaries, discussed herein, referring to the northern Alaska portion of the scheme only. They are calibrated with the ammonite zone scheme (Figure 2.1) and are presented in detail by Nikitenko and Mickey (2004). It should be noted, however, that this foraminiferal zone scheme contains a regionally extensive component; in that each geographic region (in the circumpolar area) contains a slightly different representation of individual zones and their boundaries. This is represented when comparing the Upper Pliensbachian part of the zonal scheme from northeast Siberia and Russia versus northern Alaska (Figures 2 and 4 in Nikitenko and Mickey, 2008). Within these two geographic areas, northeastern Siberia and Russia contain the JF5 and JF6 zones, whereas in northern Alaska these zones do not seem to appear. Furthermore, this foraminiferal zone scheme often contains zones that are temporally divided into both large- and fine-scales, which is represented by a particular zone that has a long stratigraphic range and contains shorter ranging zones within its boundaries (e.g. the long ranging JF4 Zone with shorter ranging JF7–JF9 zones that are contained within its boundary in Figure 2.1). This could potentially indicate the presence of a regionally extensive subzone-level temporal constraint. The Pliensbachian and Toarcian portion of this zone scheme for northern Alaska contains nine foraminiferal zones (Figure 2.1). Within this scheme, the base of the Pliensbachian does not occur at a zone boundary but rather within the Trochammina inusitata–Turritellella volubilis JF2 Zone. The JF2 Zone is long ranging and extends from the Upper Sinemurian to the lower part of the Upper  48  Pliensbachian (Nikitenko and Mickey, 2004). The upper boundary of the JF2 Zone occurs within the Late Pliensbachian Stokesi Zone (Figure 2.1). Overlying the JF2 Zone is the Trochammina lapidosa JF4 Zone (Figure 4 in Nikitenko et al., 2008). The JF4 Zone occurs in northwest Siberia where it contains the species Ammodiscus siliceus, Hyperammina ex gr. odiosa, Saccammina sp. and Jaculella jacutica (Nikitenko and Mickey, 2004). This zone is long ranging, extending from near the base of the Upper Pliensbachian to the top of the Antiquum Zone of the Early Toarcian (Figure 2.1). The Anmarginulina arctica–A. gerkei JF7–JF8 Zone occurs within the boundaries of the JF4 Zone and is known to be strictly Late Pliensbachian in age (Figure 4 in Nikitenko et al., 2008). The JF7–JF8 Zone extends from the upper part of the Stokesi Zone to the middle of the Viligaensis Zone (Figure 2.1) and includes a wide variety of taxa (Nikitenko and Mickey, 2004). Overlying the JF7– JF8, and still within the stratigraphically broader JF4 Zone, is the Recurvoides taimyrensis JF9 Zone. This zone spans the Pliensbachian / Toarcian boundary and extends to the top of the Antiquum ammonite Zone, which is also coincident with the top of the JF4 zone (Figure 2.1). In particular, they note the disappearance of JF7–JF8 Zone taxa in the lower part of the zone and the appearance of Early Toarcian taxa in its upper portion, which is taken to denote the Pliensbachian / Toarcian boundary (Nikitenko and Mickey, 2004). The stratigraphically lowest Toarcian taxa, as noted by Nikitenko and Mickey (2004), include Trochammina kisselmani, Triplasia kingakensis and Ammodiscus glumaceus. At the upper boundary of the JF9 Zone there is a distinct change in  49  the taxonomic composition of foraminifera which marks the beginning of the overlying Trochammina kisselmani JF10 and the Ammobaculites lobus– Trochammina kisselmani JF11 zones (Nikitenko and Mickey, 2004; Nikitenko et al., 2008). The JF10 Zone is another long ranging zone and extends to the top of the Lower Toarcian, correlating with the Middle Toarcian of northwest Europe (Figure 4 in Nikitenko et al., 2008). It contains the entire JF11 Zone and the lower portion of the Astacolus praefoliaceus–Lenticulina multa JF12 Zone (Figure 2.1). The JF11 Zone extends to a stratigraphic level that is equivalent with the lower part of the Commune Zone. Lastly, the JF12 Zone extends to the Lower Aalenian and includes the species Lenticulina toarcense, and Nodosaria pulchra (Nikitenko and Mickey, 2004).  50  Chapter 3 Methods: extinction & geochemical analyses 3.1 Extinction analysis Pliensbachian and Toarcian ammonoid and foraminiferal species-level diversities in western North America were measured from compiled stratigraphic range charts derived from many previous accounts (Figures 2.2–2.4; 3.1–3.4). New biostratigraphy from the Talkeetna Mountains (Alaska) and Haida Gwaii (British Columbia), presented in this study, helped to supplement and modify previously known stratigraphic ranges of several Late Pliensbachian and Early Toarcian ammonoids including the species Fanninoceras (Charlotticeras) cf. maudense, Leptaleoceras? sp., Lioceratoides (Lioceratoides) cf. L. involutum, Dactylioceras cf. compactum, Cleviceras spp., Hildaites spp. and Hildaites murleyi (Caruthers et al., 2011; Caruthers and Smith, 2012). Stratigraphic range charts, used herein, incorporate occurrences of fossils from many areas of western North America. Species ranges were then analyzed using various metrics established by Foote (2000) as well as Hammer and Harper (2006) and include: Total diversity, total diversity minus singletons, estimated mean standing diversity, per-taxon rate, Van Valen rate, and estimated per-capita rate (defined on Figure 3.7). The resulting values are a measure of diversity and extinction/origination rates for these two taxonomic groups throughout the Pliensbachian–Toarcian interval.  51  52  Figures 3.1–3.4 - Combined stratigraphic ranges of foraminifera species for two areas of western North America throughout the Pliensbachian and Toarcian. Stratigraphic ranges of foraminifera species are from various sections in Kottachchi et al. (2002, 2003) and well core samples in Tappan (1955). Localities used here are presented in Figures 3.5 and 3.6 of this study. Ranges are plotted with respect to the ammonite zone scheme of Smith et al (1988) and Jakobs et al. (1994). Note that species in bold occur in both areas.  53  54  55  Throughout western North America, Pliensbachian and Toarcian ammonoid faunas are common and are widely distributed throughout many stratigraphic successions (red circles in Figures 3.5, 3.6; compiled from data in Smith et al., 1988; Jakobs, 1992; Jakobs et al., 1994; Jakobs, 1995; Smith and Tipper, 1996; Jakobs and Smith, 1996; and Jakobs, 1997). In comparison, Pliensbachian and Toarcian foraminiferal faunas are much less well understood. They have only been documented from two areas of western North America, 56  Haida Gwaii and northern Alaska (green circles in Figures 3.5; Figure 3.6; compiled from data in Tappan, 1955; Kottachchi et al., 2002, 2003). Ammonoids were pelagic organisms whereas Early Jurassic foraminifera had a benthic mode of life. Planktonic foraminifera did not evolve until the Middle Jurassic (Hart, 1999; Hart et al., 2002). Therefore, together, these two groups occupied many parts of the marine ecosystem and were subjected to changes in environment (including bottom-water anoxia) that may have occurred in both the water column and on the sea floor. They therefore play a key role in understanding the extinction event, which is a primary focus of the study presented herein.  57  Figure 3.5A–C - Maps showing the location of previously published Pliensbachian and Toarcian stratigraphic sections in western North America. Sections contain occurrences of ammonites (red circles) and foraminifera (green circles) that are used in this study. Locality numbers 1–22 refer to the section correlation chart (Figure 3.6). YK = Yukon, NWT = Northwest Territories, BC = British Columbia, AK = Alaska, OR = Oregon, NV = Nevada, WA = Washington, MT = Montana.  Figure 3.6 - Pliensbachian and Toarcian stratigraphic sections in western North America. Sections contain occurrences of ammonites and foraminifera that are used in this study to compile stratigraphic range charts. SB#3 = South Barrow #3 well core, HC = Hicks Creek, CC = Camp Creek, YR Sec. = Yakoun River Section. Occurrences of foraminifera were placed in context of the Pliensbachian and Toarcian ammonite zone schemes in order to develop comparable 58  stratigraphic range charts, which seems to be the first reported attempt of this type of integration in Early Jurassic successions of western North America. Foraminifera were only used from sections that contain coeval ammonite faunas. In the case of Early Jurassic foraminiferal occurrences from Haida Gwaii, Pliensbachian and Toarcian successions could only be used if previous accounts had established a zone-level ammonite biostratigraphy for the section (green circles in Figures 3.5, 3.6). In northern Alaska, foraminifera are known primarily from a variety of well cores originally extracted from the National Petroleum Reserve on the North Slope (Tappan, 1955; Bergquist, 1966). Within these well cores, co-occurring Pliensbachian–Toarcian ammonite and foraminifera faunas are exceedingly rare with the exception of a single well core known as the United States Navy South Barrow #3 Core (identified herein as SB #3 core). The SB #3 core contains a diverse foraminifera fauna that is temporally constrained to the Late Pliensbachian–Early Toarcian part of the ammonoid zone scheme for the high Arctic of Zakharov et al. (2006) Nikitenko and Mickey (2004) and Nikitenko et al. (2008). Stratigraphic ranges of foraminifera from the SB #3 core (orange bars in Figures 3.1–3.4) are calibrated with the ammonite zone scheme of western North America (Figure 2.1). Open nomenclature taxa (Bengston, 1988) are included, herein, in the stratigraphic range charts of this study. The term “aff.” denotes a new species that could not be named due to a small sample size, stratigraphic uncertainties, or other ambiguities. The term “cf.”, short for “Confer”, is used to denote a  59  tentative or provisional level of identification, mostly used when preservation is poor. The terms “sp. A”, “sp. B”, “sp. C” (etc.) are used to indicate separate but as yet unnamed species within a particular genus, while “species indet.” refers to an indeterminate species of a particular genus. The term “ex gr.” or “gr.” refers to “of the group” and, lastly, a question mark is used to denote uncertainty, mostly at the genus level. Despite the uncertainty, these somewhat tentatively identified species are nonetheless included within the stratigraphic range charts. This is largely because they represent taxa that have specific stratigraphic ranges and may infact be an individual (new?) species that could be separated from the well-known species with which it is being compared. They should be included in diversity and rate estimates. When plotting stratigraphic ranges of foraminiferal species (Figures 3.1– 3.4), updated taxonomic identifications of foraminifera were incorporated into the stratigraphic range charts. Systematic descriptions of foraminiferal species from Haida Gwaii were not provided in Kottachchi (2001) or Kottachchi et al. (2002; 2003). Consequently, the taxonomic relationships of this fauna to the coeval fauna from northern Alaska (e.g. Tappan, 1955) are not certain. However, updated taxonomic identification by Nagy and Johansen (1991) and Nikitenko and Shurygin (1992) re-assign certain species from northern Alaska (in Tappan, 1955) to different genera. This re-assignment is maintained here, not only within the foraminiferal fauna from northern Alaska, but also with similarly identified species from Haida Gwaii. Updated taxonomic identifications (re-assigned taxa)  60  include: Kutsevella barrowensis, Laevidentalina pseudocommunis, Grigelis apheilolocula, Reussoolina aphela, Ammovertellina irregularis, and Ammoglobigerina canningensis, Ammodiscus asper, Reophax metensis, Ammobaculites lobus and Ammodiscus siliceus.  3.1.1 Diversity measurements Ammonoid and foraminiferal species diversity is measured using methodology presented in Foote (2000) as well as Hammer and Harper (2006). Diversity is measured from the recently compiled species level stratigraphic range charts for the Pliensbachian and Toarcian stages (Figures 2.2–2.4; 3.1– 3.4), first at the zone level and then re-measured with an informally divided zone scheme where zones were grouped into three intervals (lower, middle, and upper; Figure 3.7A). Using these intervals, the stratigraphic range of each species was divided into one of four mutually exclusive categories (Figure 3.7B) following Barry et al. (1995) and Foote (2000). These include: 1) taxa whose first and last appearance are confined to the interval (Nfl), 2) taxa that cross the bottom boundary and disappear within the interval (Nbl), 3) taxa that appear within the interval and cross the upper boundary (Nft), and 4) taxa that range through the entire interval and cross both stratigraphic boundaries (Nbt). Taxa were then assigned a unitary weight depending on their stratigraphic range within the interval (Sepkoski, 1975; Foote, 2000; Hammer, 2003; Hammer and Harper, 2006). Taxa that range through the interval (Nbt) were counted as one unit each, taxa that disappeared (Nbl) or appeared (Nft) within the interval were only counted as a half of a unit, and taxa that were confined to the interval (Nfl) were counted as a third of a unit (Hammer, 2003). Hammer and Harper (2006) justify these 61  unitary weights by the rational that the first appearance (Nft) and last appearance (Nbl) datum are uniformly distributed throughout the measured interval and that the average portion of the interval length that is occupied by both of these types of single-ended taxa is approximately 0.5. Similarly, taxa that are confined to the interval (Nfl) are thought to occupy approximately one-third of the interval length and therefore should be counted as such (Hammer, 2003; Hammer and Harper, 2006). Categories were then totaled with respect to each stratigraphic interval (zone or informally divided zonal unit) and then analyzed using various metrics to calculate taxonomic diversity and extinction/origination rates (Figure 3.7C). Lastly, the term singleton has been used to denote either: 1) species that are represented by a single specimen (Buzas and Culver, 1994; 1998) or 2) taxa that are confined to a single stratigraphic interval, essentially any species that is in the Nfl category (Foote, 2000). In this study, the definition of ‘singleton taxa’ by Foote (2000) is used, suggesting taxa that that are confined to a particular interval.  62  Figure 3.7 - A) Example of stratigraphic ranges of hypothetical species, divided at the zone- and informal subzone-levels, used in this study for measuring taxonomic diversity and extinction/origination rates. Taxa from within each informal subzone are identified as being from the ‘lower’, ‘middle’, or ‘upper’ part of the zone. B) Four fundamental classes of taxa (after Foote, 2000) used in this study to quantify stratigraphic ranges of species in western North America. C) Definitions of diversity and rate measures used in this study (after Foote, 2000). O = Origination rate, E = Extinction rate.  3.2 Geochemical analysis Over the 2008–2010 field seasons, samples were collected for stable isotope geochemistry (δ13Corg, δ13Cwood, δ15Norg) total organic carbon, and total nitrogen analysis from two well studied and temporally constrained stratigraphic sections on Haida Gwaii (British Columbia, Canada) and the SB #3 core (Figure 63  3.5). On Haida Gwaii, samples were collected from Whiteaves Bay on northern Moresby Island and along the Yakoun River on central Graham Island (Figure 3.5C1,C2). In northern Alaska (Figure 3.5A), the SB #3 core has been split (lengthwise) with half of the split being stored at the United States Geological Survey (USGS) in Denver, Colorado and the other half stored at the Geological Materials Center (GMC) in Eagle River, Alaska. The split that is stored at the GMC in Eagle River was sampled in the current study. In the field on Haida Gwaii, each section was measured and samples were collected at regular stratigraphic intervals ranging from 5cm to 40cm spacing. Co-occurring ammonites were also collected. Samples were prepared at UBC for isotopic analysis in the paleontology laboratory using methods that are consistent with the Stable Isotope Biogeochemistry Laboratory (SIBL) protocol at Durham University (UK). All samples were air-dried in their sample bag with the top opened. Shale samples were prepared separately from wood samples using slightly different preparation methods. Shale samples were ground to a powder and homogenized using an agate mortar and pestle and then placed into glass 25ml screw-top vials according to sample number. In between samples, the agate mortar and pestle were ground using quartz sand and wiped clean using acetone to avoid cross contamination of samples. The powdered sample was then placed into a 50ml centrifuge tube with enough powder to cover the conical portion of the tube. Three Molar HCl was added up to the 45ml mark on each tube and left to decalcify for ~8 hours (or overnight). To avoid overflow, HCl was sometimes  64  added slowly if a high concentration of calcite was present. After the HCL was fully added, the lid of the centrifuge tube was tightened and the vial shaken by hand until all of the sediment was suspended in solution. The lid was slowly uncapped to allow the CO2 gas to escape, and then this step was repeated. The lid was then loosened and left slightly ajar overnight. After ~8 hours or overnight, the 3M HCl supernatant was decanted, the centrifuge tube refilled with deionized water (thus rinsing the sample), and then the sample was centrifuged at ~3000 rpm for 6 minutes. This rinsing step was performed four times with the deionized water being decanted off after each centrifuge step. After the last rinse, the sample was placed in a drying oven set at 60ºC for 24 hours or until dry. Once dry, the sample was then re-ground to a powder using an agate mortar and pestle and cleaned in-between samples with quartz sand and acetone (same as above). The decalcified, powdered sample was then transferred to a small 1dram glass vial and transported to the SIBL laboratory in Durham (UK) for stable isotope analysis. Preparation of wood samples involved a slightly different methodology in comparison to the shale. After the previously described drying step, fossil wood was separated from the surrounding sediment using tweezers under a dissecting microscope. The wood sample was ground using an agate mortar and pestle, which was cleaned in between each sample using the same methodology as above. Once ground and homogenized, the wood sample was placed into a 15ml centrifuge tube with as much sample wood as possible, or until the sample reached the top of the conical portion of the tube. Three Molar HCl was then  65  added up to the 12ml in the tube, shaken (same as above), and then left to decalcify overnight. The supernatant was then decanted, the sample rinsed with de-ionized water, and centrifuged using the same methodology as above. Rinsing was completed four times and then the sample was placed into a drying oven, set at 60ºC for 24 hours. Once dry, the sample was re-ground (same as above) and then transferred to a small 1-dram glass vial and shipped to Durham (UK) for stable isotope analysis. At the SIBL laboratory each sample was weighed on a mass balance according to the approximated concentration of organic carbon in each sample. For this study, between 2–4 mg of powder was weighed per decalcified shale sample and < 1 mg of powder was weighed per decalcified wood sample. Weighed samples were then placed into tin sample cups and analyzed by mass spectrometry to obtain δ13C, total organic carbon, δ15N and total nitrogen values. Anomalous isotope values were re-analyzed if they were either very different (more negative or positive) in comparison to stratigraphically adjacent values or were considered suspect from any technical/mechanical problems that may have occurred in the analysis.  66  Chapter 4 Regional geology The Cordilleran region of western North America (Figure 4.1) consists of allochthonous, fault-bounded, tectonostratigraphic terranes that are composed of sedimentary, igneous and metamorphic rock units (Coney et al., 1980; Plafker and Berg, 1994; Miller et al., 2002; Haggart et al., 2006; Ridgway et al., 2007; Blodgett and Stanley, 2008; Colpron and Nelson, 2009). Currently it is believed that many of the more inboard terranes (e.g. Yukon–Tanana, Quesnellia, and Stikinia) originated in the mid Paleozoic along the western margin of western North America as a series of rifted continental fragments (Colpron and Nelson, 2009 and references therein). This differs from the more outboard terranes (e.g. Wrangellia, Alexander, Peninsular, and Arctic Alaska–Chukotka) containing Early Paleozoic and older basement sequences that are thought to have originated far from western North America in a variety of geographically distant areas that include Baltica, Siberia, and the northern Caledonides (Colpron and Nelson, 2009 and references therein). Although the timing of terrane accretion to the western margin of North America has been debated, it is currently believed that many of the displaced terranes of the North American Cordillera began to collide with the western margin of North America in the Late Paleozoic as a series of belts that continued accreting throughout the Late Cretaceous (Coney et al., 1980; Plafker and Berg, 1994; Trop et al., 2002; Colpron and Nelson, 2009).  67  Figure 4.1 - Tectonic map of western North America showing major fault systems and terrane boundaries (modified from Colpron and Nelson, 2009; Beranek, 2009). Terranes that are discussed in this study are coloured. Boundaries of the Peninsular terrane are from Trop et al. (2002; 2005). Red boxes refer to areas discussed herein.  68  Of particular importance is the western-most outboard crustal fragment known as the Wrangellia composite terrane (Plafker and Berg, 1994; Nokleberg et al., 1994) and the northern-most Arctic Alaska–Chukotka terrane (Till and Dumoulin, 1994; Amato et al., 2004; 2009). The Wrangellia composite terrane (Figure 4.1) is a subcontinent sized crustal fragment that consists of three large tectonostratigraphic terranes, which include the Wrangellia, Alexander, and Peninsular terranes (Plafker and Berg, 1994; Nokleberg et al., 1994). The Arctic Alaska–Chukotka terrane, also illustrated on Figure 4.1, encompasses much of present-day northern Alaska (Brooks Range and Seward Peninsula), Russia (Chukotka Peninsula), East Siberia and Bering continental shelves (Amato et al., 2004; 2009; Colpron et al., 2007). During the Pliensbachian and Toarcian, it is thought that the Wrangellia composite terrane was located in the northeast paleo Pacific Ocean at a paleolatitude that is comparable with the modern US / Canada border (Smith, 2006) and the Arctic Alaska–Chukotka terrane was located adjacent to Laurentia close to its present-day position in the high Arctic (Miller et al., 2006; Amato et al., 2009).  4.1 Haida Gwaii The Haida Gwaii, are located off mainland British Columbia, Canada in the eastern Pacific Ocean (Figure 3.5C). This group of islands is part of Wrangellia, which is currently distributed along the western margin of the North American Cordillera from Vancouver Island to southern Alaska (Jones et al., 1977; Coney et al., 1980). It has been informally divided into two portions based on modern geographical placement (Green et al., 2010): 1) a northern portion, which  69  includes The Wrangell Mountains (Alaska) and parts of the Yukon (Canada) and 2) a southern portion, which includes Vancouver Island and Haida Gwaii (British Columbia). Wrangellia was originally defined as a fault-bounded section of upper crust containing correlative successions of Middle–Late Triassic flood basalt with overlying Late Triassic carbonate that, altogether, overlie Paleozoic units (Jones et al., 1977; Greene et al., 2010). However, it has been determined that deposition began in the Devonian and continued intermittently throughout the Cretaceous until its collision with Western North America by Latest Cretaceous time (Csejtey et al., 1982; McClelland et al., 1992; Nokleberg et al., 1994; Plafker and Berg, 1994; Hillhouse and Coe, 1994; Trop et al., 2002; Umhoefer and Blakey, 2006; Trop and Ridgway, 2007; Greene et al., 2010). On Haida Gwaii, Upper Triassic to Lower Jurassic strata (Figure 4.2) were deposited continuously in a backarc, deep water, environment that produced the Karmutsen Formation (not included in Figure 4.2) overlain by the Kunga and Maude Groups (Andrew and Goodwin, 1989; Thompson et al., 1991; Tipper et al., 1991). The Maude Group encompasses the Pliensbachian–Aalenian part of the stratigraphy and is composed predominantly of volcaniclastic and pyroclastic sediments that were deposited at varying water depths in a shelf environment (Cameron and Tipper, 1985; Jakobs, 1997). In ascending stratigraphic order the Maude Group includes the Ghost Creek, Fannin, Whiteaves, and Phantom Creek formations (Figure 4.2). Outcrops of Maude Group sedimentary rocks are best exposed along the intertidal zone of the Skidegate Inlet area (Cameron and  70  Tipper, 1985), as well as along many road- and stream-cuts of the Yakoun River drainage in central Graham Island (Figure 3.5C1, C2).  Figure 4.2 - Lower–Middle Jurassic stratigraphy and correlative ammonite zonal scheme for Pliensbachian–Toarcian units on Haida Gwaii, British Columbia (Modified from Jakobs and Smith, 1996; Jakobs, 1997). Pliens. = Pliensbachian, Yakoun = Yakoun Group, R.J.M. = Rennell Junction Member.  4.1.1 Lithostratigraphy The Ghost Creek Formation is an approximately 60 m thick, recessive dark grey/black, organic-rich, shale and silty shale unit that contains abundant  71  pyrite and occasional sandstones, nodular limestones and rare ash or tuffaceous interbeds (Cameron and Tipper, 1985; Smith and Tipper, 1996). It rests conformably on the Sandilands Formation of the Kunga Group (Figure 4.2; Cameron and Tipper, 1985) and has been temporally constrained to the Imlayi and Whiteavesi Zones by Smith and Tipper (1996). The Ghost Creek Formation contrasts with the underlying Sandilands Formation in being softer with more shale, contains less volcanic material, and lacks lamination (Cameron and Tipper, 1985; Smith and Tipper, 1996). In the Skidegate Inlet area, the basal part of the Ghost Creek Formation shows highly contorted beds, calcite veins, and shale rip-up clasts which Cameron and Tipper (1985) suggest are the result of higher energy deposition but do not specify a specific environment of deposition. Near the top of the formation, in areas of Central Graham Island, the Ghost Creek Formation is more regularly bedded and contains glauconite which is interpreted as evidence for a slower rate of sedimentation (Cameron and Tipper, 1985). Benthic fossils and bioturbation are uncommon, but are noted to increase in abundance near the upper part of the unit (Cameron and Tipper, 1985; Smith and Tipper, 1996). Lying conformably above the Ghost Creek Formation is the Early Pliensbachian (Whiteavesi Zone) to Early Toarcian (Kanense Zone) Fannin Formation (Figure 4.2; Tipper et al., 1991; Smith and Tipper, 1996; Jakobs, 1997). Original description of this sequence by Cameron and Tipper (1985) identifies a lower unit called the Rennell Junction Formation. However, this was subsequently redefined as the Rennell Junction Member of the Fannin Formation  72  because the unit was unmappable (Tipper et al., 1991; Smith and Tipper, 1996). The Rennell Junction Member is described as a fine grained sandstone, siltstone, and shale unit that contains irregular limestone beds and nodules and ranges in thickness from 15 to 40 m (Smith and Tipper, 1996). In the Skidegate Inlet area, the base of the Rennell Junction Member occurs at the base of the Fannin Formation. The contact of the Rennell Junction Member and the overlying upper Fannin Formation occurs near the Freboldi / Kunae Zone boundary (Smith and Tipper, 1996). The upper part of the Fannin Formation is generally coarser-grained, more resistant, and contains thicker beds of crossbedded sandstone that are occasionally convoluted. The upper Fannin Formation also contains limestone interbeds, chamosite ooliths, locally developed concretions, tuffs, breccias, and conglomerate lenses (Cameron and Tipper, 1985; Smith and Tipper, 1996). The Fannin Formation also contains a fairly diverse benthic and pelagic fauna including: ammonites, nautiloids, bivalves, brachiopods, radiolarians, foraminifera, and bioturbation (Cameron and Tipper, 1985; Smith and Tipper, 1996; Kottachchi et al., 2002). The Early to Late Toarcian (Kanense–Hillebrandti Zone) Whiteaves Formation is a recessive unit that conformably overlies the Fannin Formation (Figure 4.2; Jakobs, 1997). It consists predominantly of grey/green medium to fine grained siltstone and mudstone with medium to fine grained sandstone interbeds of variable thickness, zircon-bearing ash beds, pale grey/light brown calcareous nodules, and concretionary limestones (Cameron and Tipper, 1985;  73  Jakobs, 1997). The abundance of siltstone and mudstone throughout the succession and no evidence of transport indicators within sandstone interbeds, suggest the Whiteaves Formation was deposited slowly in a moderately deepwater environment (Cameron and Tipper, 1985). The presence of glauconite in some sandstones suggest a depositional paleo-depth between 200–300 m (Cameron and Tipper, 1985 and references therein). However, the occasional local presence of fragmented bivalve shells in small sandy pockets suggests a minor degree of intermittent higher energy re-sedimentation (Cameron and Tipper, 1985) which was not observed in stratigraphic sections of the Whiteaves Formation within the study presented herein. The Late Toarcian (Yakounensis Zone) to Aalenian Phantom Creek Formation rests conformably on the Whiteaves Formation and is separated from the overlying Yakoun Group by an angular unconformity (Figure 4.2; Cameron and Tipper, 1985; Jakobs, 1997). It is an approximately 30 m thick, resistant, brown to buff weathering, fine–coarse grained, sandstone unit that contains partly calcareous and common argillaceous grains, thin shale interbeds in its lower part, massively bedded sandstone units in its upper part, and a rich fossil fauna throughout that includes ammonites, bivalves and belemnites (Cameron and Tipper, 1985). It has been informally subdivided into two members that are separated by an erosional hiatus (Cameron and Tipper, 1985; Jakobs, 1990; Jakobs, 1997). The lower member is known as the coquinoid sandstone member (Cameron and Tipper, 1985; Jakobs, 1997). It is a 3-4 m thick, well-bedded, grey  74  to greenish-brown sandstone unit that contains an abundant bivalve and ammonite fauna (Cameron and Tipper, 1985). Fossils within this unit are tightly packed, randomly oriented, and commonly broken. Overlying this unit is the belemnite sandstone member (Cameron and Tipper, 1985; Jakobs, 1997). It is an approximately 20 m thick unit composed of massively bedded, carbonaceous, brown weathering sandstone. This unit is characterized by an abundance of belemnites, pectinoid bivalves, gastropods and nautiloids, with rare occurrences of ammonites (Cameron and Tipper, 1985; Jakobs, 1997). The Phantom Creek Formation is interpreted as being deposited in a high-energy shallow water environment based on: 1) the disappearance of shale in its lower part, 2) the appearance of massively bedded coarse-grained sandstone in its upper part, and 3) appearance of coquinas that are densely packed, randomly oriented and contain highly fragmented fossils (Cameron and Tipper, 1985; Jakobs, 1997).  4.2 Northern Alaska The Arctic Alaska–Chukotka terrane (Figure 4.1) includes many different fragments of the present-day circum-Arctic region and is generally composed of continental rocks in northern Alaska and northeastern Russia (Moore et al., 1994; Amato et al., 2004; 2009; Miller et al., 2006; Colpron et al., 2007). It is an extremely large terrane that covers an area of approximately 3 million km2 which is more than 85% the area of Greenland (Amato et al., 2009). It is bounded to the North by the edge of the Arctic Alaskan outer shelf, and to the south by arc and ophiolite rocks of the southern Brooks Range (Moore et al., 1994; Miller et al., 2006).  75  The northern Alaska portion of the Alaska–Chukotka terrane lies beneath much of the present-day continental shelves of the U.S. Chukchi and Beaufort Seas and the North Slope of Alaska (Hubbard et al., 1987a,b; Sherwood et al., 2002; Houseknecht and Bird, 2004). It is comprised of sedimentary sequences from two connected basins, the Hanna Trough and Arctic Alaska Basin, that rest on a basement sequence of highly deformed Devonian and older crystalline rocks known as the Franklinian sequence (Moore et al., 1994; Sherwood et al., 2002; Miller et al., 2006; Amato et al., 2009). Today, these basins are bordered by the Barrow arch to the north, Brooks Range to the south, Herald Thrust to the West, and continental Canada (Laurentia) to the East (Figures 3.5A and 4.1). In northern Alaska, deposition of the Alaska–Chukotka terrane began in the Late Devonian (?) and continued intermittently until the Late Jurassic / Cretaceous, with sequences of marine carbonate and nonmarine–shallow marine siliciclastic strata that are approximately 6 km thick (Hubbard et al., 1987a,b; Bird, 1988). This sequence is widely known as the Ellesmerian sequence (Moore et al., 1994; Moore et al., 1997; Natal’in et al., 1999; Dumoulin et al., 2002; Sherwood et al., 2002; Amato et al., 2003; Houseknecht and Bird, 2004; Miller et al., 2006; Amato et al., 2009) and is thought to represent rift- (lower Ellesmerian) and subsidence- (upper Ellesmerian) phases of basin development (Sherwood et al., 2002). During the Jurassic (Late Jurassic in Sherwood et al. (2002) and Early–Middle Jurassic in Hubbard et al. (1987a,b)) the northeast part of the basin experienced a rifting event that was associated with the opening of the Arctic Ocean. This rifting is known as the Beaufortian sequence and it is thought to  76  have covered the Ellesmerian rocks with thick deposits of rift-related clastics, sandstones, siltstones and shales (Hubbard et al. 1987a,b; Sherwood et al., 2002; Houseknecht and Bird, 2004). Subsurface sediments in the South Barrow area of northern Alaska contain strata of the Beaufortian sequence (Figure 4.3) and were sampled in this study. Several other studies have addressed the paleogeographic history of the Arctic region (Lawver and Scotese, 1990; Blodgett et al., 2002; Lawver et al., 2002; Miller et al., 2006; Amato et al., 2009), some suggest that during the Neoproterozoic, the Arctic Alaska–Chukotka terrane was located in a geographic position northeast of Laurentia and had faunal ties with Siberia and Laurentia by the Early Ordovician (Blodgett et al., 2002; Amato et al., 2009). Other studies have suggested that sediments of the Ellesmerian and Beaufortian sequences were derived from an exposed landmass to the North, and were then subsequently deposited on a south-facing continental shelf (Moore et al., 1994; Sherwood et al., 2002; Houseknecht and Bird, 2004). However, it is thought that by the Late Paleozoic–Jurassic, the Arctic Alaska–Chukotka terrane as a whole was located in the high Arctic (above 50°N) near its present location (Amato et al., 2009).  77  Figure 4.3 - Schematic diagrams showing a summary of the inferred depositional sequence of the Kingak Shale through the National Petroleum Reserve in Alaska (modified from Houseknecht and Bird, 2004). A) shows the litho- and chronostratigraphy of the various formations that comprise the Kingak Shale and B) depicts a generalized South–North cross section of the strata through the Barrow Arch.  4.2.1 Lithostratigraphy The Barrow area of northern Alaska contains a thick sequence of Beaufortian rock generally identified as the Kingak Shale (Houseknecht and Bird, 2004 and references therein). It is a geographically extensive unit that is described from many surface outcrops along the North Slope and in subsurface  78  wells of the National Petroleum Reserve (NPRA) of northern Alaska. As illustrated by Houseknecht and Bird (2004), this all-inclusive unit contains three formations of Jurassic–Cretaceous age (Figure 4.3) and is stratigraphically confined by the Shublik and Sag River Formations (below) and a geographically extensive ‘pebble shale’ unit (above). The Kingak Shale is truncated to the North by a Lower Cretaceous unconformity which is thought to represent, among others, erosion of the unit by the uplift of the Barrow arch (Houseknecht and Bird, 2004). The formally defined Kingak Formation is the basal unit of the Kingak Shale and contains upper and lower units that are separated by a disconformity (Figure 4.3). The Lower Kingak Formation includes sedimentary sequences that are of Early Jurassic (Hettangian–Toarcian) age, while the Upper Kingak Formation contains mainly Late Jurassic (Oxfordian–Kimmeridgian) deposits (Imlay, 1981; Sherwood et al., 2002; Houseknecht and Bird, 2004; Mickey et al., 2006). The Lower Kingak Formation is a 900–1200 m thick fissile, light to dark grey, argillaceous siltstone and claystone with minor sandstone interbeds and glauconite (Collins, 1961; Imlay, 1981; Sherwood et al., 2002; Houseknecht and Bird, 2004). Temporal constraint of the Lower Kingak Formation is provided mainly through ammonite and foraminifera that were collected in subsurface well cores within the National Petroleum Reserve (Imlay, 1955; 1981; Tappan, 1955; Bergquist, 1966). However, collections of ammonites from surface outcrops along the North Slope have also provided temporal constraint for the formation (Imlay, 1955; 1981; Smith et al., 2001).  79  4.3 Southern Alaska The Peninsular terrane of southern Alaska (Figure 4.1) is one of the largest outboard tectonostratigraphic terranes in western North America (Coney et al., 1980; Wilson et al., 1985; Wang et al., 1988). It comprises much of the modern-day Alaskan Peninsula extending ~1200 km throughout the Talkeetna Mountains and is bordered by the Chugach terrane to the southeast and the Wrangell terrane to the east (Csejtey et al., 1978; Jones and Silberling, 1979; Csejtey and St. Aubin, 1981; Wilson et al., 1985). Sequences deposited on the Peninsular terrane are thought to have initially accumulated in the paleo Pacific Ocean far from the North American margin (Wilson et al., 1985; Wang et al., 1988; Plafker et al., 1989, 1994; Clift et al., 2005a,b; Blodgett and Stralla, 2006; Rioux et al., 2007). Deposition began in a shallow-water, tropical, backarc environment during Permian–Triassic time (Wang et al., 1988; Blodgett and Stralla, 2006) and shifted toward an intraoceanic, volcanic island arc-type environment throughout the Early Jurassic and into the early Middle Jurassic (Barker and Grantz, 1982). Island arc related magmas have recently been dated using U–Pb zircon and whole-rock isotope analysis indicating that magmatism occurred primarily between 202.1 and 181.4 Ma (Clift et al., 2005a,b; Rioux et al., 2007). By Middle–Late Jurassic time, it is thought that the depositional environment had shifted toward a forearc including deep-water fan-delta deposition, shelf sedimentation, post-depositional uplift, and strike-slip displacement (Trop et al., 2005). Current tectonic models postulate that the Peninsular terrane was amalgamated with Wrangellia and the Alexander terrane, forming the Wrangellia  80  composite terrane, which then accreted to the western margin of North America (Wang et al., 1988; Plafker et al., 1989; Ridgway et al., 2002; Trop et al., 2005). However, the timing of amalgamation versus the timing of accretion is a complex issue that is still unclear. Originally it was thought that the composite terrane was assembled by Triassic time (Wang et al., 1988; Plafker et al., 1989) and accreted to the continental margin of western North America between the Late Jurassic and Late Cretaceous (Csejtey et al., 1982; Jones et al., 1982; Pavlis, 1982; McClelland et al., 1992; Cole et al., 1999; Ridgway et al., 2002; Trop et al., 2002). Other studies have indicated that collision of the composite terrane with western North America may have occurred as early as Middle Jurassic–Early Cretaceous time (McClelland and Gehrels, 1990; McClelland et al., 1992; van der Heyden, 1992; Kapp and Gehrels, 1998; Gehrels, 2001). More recent work concludes that Late Jurassic deformation and synorogenic sedimentation could reflect either the initial collision of Wrangellia and the Peninsular terranes or their collision with western North America (Trop et al., 2005).  4.3.1 Lithostratigraphy There are no known crystalline basement rocks underlying the Peninsular terrane (Wang et al., 1988; Blodgett and Stralla, 2006). Previous work has shown that the oldest strata crop out at Puale Bay along the Alaska Peninsula. They consist of a small (~11 m thick) unnamed unit of volcanic agglomerate, volcaniclastic sandstone, and fossiliferous limestone that is thought to be of Middle Permian age (Hansen, 1957; Jeffords, 1957; Wang et al., 1988; Blodgett and Stralla, 2006). Lying unconformably above the Permian unit is the Upper Triassic (Norian) Kamishak Formation (Capps, 1923; Smith, 1926; Silberling in 81  Detterman et al., 1985; Wang et al., 1988; Newton, 1989; Silberling et al., 1997; Pálfy et al., 1999; Blodgett and Stralla, 2006; Blodgett, 2008). This unit crops out throughout much of the Alaska Peninsula and comprises ~700 m of shallow fossiliferous marine and biogenic carbonate interbedded with volcanic rocks in its upper part (Wang et al., 1988; Blodgett, 2008). The Early Jurassic Talkeetna Formation rests conformably on the Late Triassic Kamishak Formation (Newton, 1989; Pálfy et al., 1999). The Talkeetna Formation is a widespread unit that is exposed in many outcrops and cores throughout the Alaska Peninsula extending northeast to the Talkeetna Mountains (Barker and Grantz, 1982; Nokleberg et al., 1994; Sandy and Blodgett, 2000). It is primarily composed of volcanic (flows and pyroclastics) rocks with minor interbedded marine sedimentary sequences (volcaniclastics, sandstones, siltstones, and mudstones) that reach a thickness of ~3 km (Barker and Grantz, 1982; Nokleberg et al., 1994; Trop et al., 2005). The Talkeetna Formation is thought to be Hettangian–Late Toarcian in age (Imlay and Detterman, 1973; Imlay, 1981; Newton, 1989) with zone-level age constraints published for a Hettangian–Sinemurian sequence exposed in Puale Bay (Pálfy et al., 1999). For the Talkeetna Mountain sequences there are currently no age constraints at the zone-level. Imlay (1981) indicates the presence of Late Sinemurian–Late Toarcian ammonites from several isolated localities but does not provide detailed stratigraphic distributions. Stratigraphically above the Talkeetna Formation are the Middle Jurassic– Late Cretaceous Tuxedni (Bajocian-Bathonian), Chinitna (Bathonian-Callovian),  82  Naknek (Oxfordian-Kimmeridgian), Nelchina limestone (Valanginian-Hauterivian), and Matanuska formations (Albian-Maastrichtian) representing deposition in a shallow water marine shelf, forearc environment (Grantz, 1960a,b; Winkler, 1992; Trop et al., 2005). The lithology of these units is quite variable in that the Tuxedni and Chinitna formations are composed primarily of mudstone and fine-grained sandstone; the Naknek Formation of coarse-grained cobble conglomerate, sandstone, and mudstone; the Nelchina limestone of massively bedded limestone, and the Matanuska Formation of sandstone, conglomerate and mudstone (Trop et al., 2005 and references therein). The coarse-grained cobble conglomerate lithofacies of the Upper Jurassic Naknek Formation is thought to record collision of the Peninsular terrane with either adjacent Wrangellia or the North American margin (Trop et al., 2005 and references therein). 4.3.1.1 Biostratigraphy of the Hicks and Camp Creek sections Herein, 64 Late Pliensbachian ammonoid specimens are described from two sections of the Talkeetna Formation in the southeastern Talkeetna Mountains, Anchorage D-3 quadrangle (number 2 in Figures 3.5B, 3.6) representing the Kunae and Carlottense Zones of the Late Pliensbachian (Figure 4.4). Biostratigraphy of these two stratigraphic sections helped supplement previously known stratigraphic ranges of Late Pliensbachian ammonoids in western North America (Figures 2.2–2.4, 3.1–3.4) that were utilized in the extinction analysis, and also provided several new occurrences in the coeval strata of Alaska; including the species Fanninoceras (Charlotticeras) cf. maudense, Leptaleoceras? sp., and Lioceratoides (Lioceratoides) cf. L. involutum (Figures 4.4 and 4.5). 83  Figure 4.4 - Lithostratigraphy and biostratigraphy of the Late Pliensbachian Camp Creek (A) and Hicks Creek (B) sections in the Talkeetna Mountains, southern Alaska (after Caruthers and Smith, 2012). Ammonites from the Camp Creek section are from the Kunae Zone, while ammonites recovered from the section along Hicks Creek are indicative of the Carlottense Zone. f.s. = fine sand; m.s. = medium sand; c.s. = coarse sand; peb. = pebble; cg. = conglomerate; cob. = cobble; volc. = volcaniclastic; CCA-1 = volcanic ash sample (Figure 4.6, this study).  84  4.3.1.1.1 Camp Creek section This section is located along an unnamed tributary of Camp Creek on the southeastern side of Gunsight Mountain in the Talkeetna Mountains (Figure 3.5B). It is a ~250 m thick sequence of the Talkeetna Formation that is truncated above and below by faulting (Figure 4.4A). It consists predominantly of medium– coarse grained volcaniclastic sandstone with interbedded conglomerate and minor volcanic and pyroclastic rocks. Conglomerate facies include rounded and sub-rounded basalt and felsic magma clasts with rare occurrences of chert and sandstone. The basal 41 m consists of mostly siltstone with minor sandstone and conglomerate interbeds. Above 120 m, scoured and graded beds become more frequent, generally beginning with coarser-grained facies that grade into a fine sand or siltstone, suggesting deposition by turbidity currents. Low angle cross beds are recorded at ~37 and 89 m and oscillatory ripples at ~14 and 97 m.  85  86  Figure 4.5 - Distribution of Pliensbachian and Toarcian ammonite taxa that are systematically described in Chapter 7 of this study, according to locality number at the Hicks Creek and Camp Creek sections (Figure 4.4), the Whiteaves Bay and Yakoun River sections (Figures 5.4, 5.6) and the South Barrow #3 Core (Figure 5.11). Ch. = Chapter, A.A. = Arctic Alaska, and SB3 = South Barrow #3 well core. Previous work at this section suggested that the marine sedimentation that produced the Talkeetna Formation was restricted within the Sinemurian to Toarcian interval. This was based on the occurrence of the bivalves Weyla unca, Ostrea sp., and Cardinia sp., and several specimens of the brachiopod Callospiriferina tumida (Imlay, 1981; Sandy and Blodgett, 2000). Ammonites identified herein indicate that deposition of the Talkeetna Formation (at this section) occurred over a much shorter time interval than previously thought. Forty-five specimens of Late Pliensbachian ammonoid were identified, representing five genera that occur at ten localities distributed throughout the section (Figures 4.4A, 4.5). A Kunae Zone age determination is indicated by the presence of Leptaleoceras? sp., Fanninoceras (Charlotticeras) cf. maudense, and Amaltheus sp. from localities 6, 8, 13 and 14; these taxa are known to be restricted to this zone (Smith and Tipper, 1996). The presence of Arieticeras aff. domarense and Fanninoceras (Fanninoceras) fannini from localities 6–15 are also indicative of the Kunae Zone, however, these taxa also rarely occur in the lower Carlottense Zone (Smith and Tipper, 1996). Although a Kunae Zone age is determinable for this section, its stratigraphic relationship to the overlying Carlottense Zone is uncertain due to a lack of age-specific ammonites from the  87  top of the section. A U–Pb TIMS age date has been obtained from an ash bed occurring near the top of the section at 242.2 m (CCA-1 on Figure 4.4A). The age of the ash bed, derived from the youngest cluster of zircon grains (F–J), is calculated to be 184.12 + 0.17 Ma (Figure 4.6A). Based on current time scale calibrations (Pálfy et al., 2000), this indicates a late Kunae Zone age Figure 4.6B which is consistent with the ammonite fauna from the Camp Creek section. The older population of zircon grains (A–E), yielding a date of ~188 m.y., may consist of either xenocrysts or crystals with inherited cores. These zircon crystals are most likely related to arc-related magmatism of the Peninsular terrane that occurred over a time span of 202.1 to 181.4 Ma (Clift et al., 2005a,b; Rioux et al., 2007).  88  Figure 4.6 - A) U-Pb concordia diagram for ash sample CCA-1 in the Camp Creek section of the Talkeetna Mountains, Alaska (Figure 4.4A). Diagram shows 206 Pb/238U TIMS ages for ten zircon grains showing two distinct age populations, one at ~188 Ma and the other at ~184 Ma. The age of magmatism is calculated to be 184.12 + 0.17 Ma, based on five zircon grains (F–J). The population of five zircon grains A–E at ~188 Ma consists of xenocrysts or crystals containing inherited cores. B) A comparative diagram showing the relationship of the calculated age for grains F–J of CCA-1 to the calibrated geochronologic Time Scale for the Kunae / Carlottense zonal boundary (Pálfy et al., 2000). Figure from Caruthers and Smith (2012).  4.3.1.1.2 Hicks Creek section The section cropping out along Hicks Creek in the southern Talkeetna Mountains ~10 km from the Glenn Highway (Figure 3.5B) consists of ~205 m of fossiliferous marine volcaniclastic and volcanic rocks (Figure 4.4B). The lower 43 m consist of coarse-grained volcaniclastic sandstones, interbedded conglomerate lenses and minor siltstone. Conglomerate lenses contain rounded– sub-rounded, pebble-sized clasts of basalt and other igneous rocks. Siltstones and sandy siltstones occur from 43–87 m and are directly overlain by a thick (~60 m) unit of undifferentiated volcanic/volcaniclastic rocks. The succession is capped by ~58 m of coarse-grained sandstone and sandy siltstone. In general, the stratigraphic succession at Hicks Creek is lithology similar to that of Camp Creek but differs by the predominance of finer grained sedimentary rocks and the absence of scoured bedding surfaces and graded bedding. The fauna from the Hicks Creek section is taxonomically diverse and includes ammonoids, bivalves, brachiopods, gastropods, corals, plant fossils, fish scales and trace fossils. Bivalves and brachiopods are most abundant and diverse throughout the section, frequently occurring in coquinas but also as isolated individuals in siltstone intervals. The bivalves contain a mix of free-lying 89  and encrusting forms. Other taxonomic groups such as ammonoids, gastropods and corals are rare occurring mostly in the siltstone intervals. Two intervals in the lower 48 m of the section contain abundant fish scales that occur in a thinbedded (1 cm), light gray, volcanic ash (Figure 4.4B). The diverse benthic fauna from this section will be the subject of further research. Twenty specimens of Late Pliensbachian ammonoids were identified representing five genera that occur at five localities within this stratigraphic section (Figures 4.4B, 4.5). Fanninoceras (Fanninoceras) carlottense, and Lioceratoides (Lioceratoides) cf. involutum are identified from localities 1–4 and indicate the Carlottense Zone (Smith and Tipper, 1996). Fanninoceras (Fanninoceras) fannini, which is recorded from locality 3, is known primarily from the Kunae Zone of western North America but it also occurs rarely in the lower Carlottense Zone (Smith and Tipper, 1996). Lastly, Lytoceras sp. is identified from locality 5. Lytoceras is uncommon in the Lower Jurassic of North America and the specimen from Hicks Creek is the first figured specimen from the Pliensbachian. Ammonite biostratigraphy of this section indicates that deposition of the Talkeetna Formation at Hicks Creek occurred during the Carlottense Zone of the Late Pliensbachian, a slightly younger age than that of the Camp Creek section.  90  Chapter 5 Results 5.1 Extinction data and patterns of diversity Stratigraphic ranges of 206 ammonite and 242 foraminifera species were compiled and analyzed at an informally defined ‘sub zonal’ level for Pliensbachian–Toarcian strata in western North America. Compiled ammonite species were analyzed as one dataset and foraminiferal species were first analyzed according to region (e.g. Haida Gwaii, and northern Alaska) and then as a unified dataset. This is largely because the studied localities containing ammonite species were located in the paleo Pacific Ocean during the Early Jurassic, whereas the two areas containing benthic foraminiferal species are thought to have been located on the margins of two separate ocean basins during the Early Jurassic, namely the paleo Pacific Ocean and the paleo Arctic Ocean. Patterns in species diversity for ammonite and foraminiferal species are presented in Figure 5.1. Ammonite species diversity (Figure 5.1A) is measured as a total sum (Ntot), without singletons, and as an estimated mean. Regional foraminiferal species diversity data (Figure 5.1B) are only presented as Ntot values. In both cases, diversity levels in all three metrics generally follow the same overall pattern (within each taxonomic group independently) and show very similar values, which indicates that taxonomic singletons do not have a strong influence in this analysis. Results are therefore presented in terms of the total diversity (Ntot) with occasional reference to singleton-free data. Diversity data 91  pertaining to the singleton-free and estimated mean metrics for foraminiferal species are presented in Figure 5.2.  Figure 5.1 - Pliensbachian–Toarcian ammonite (A) and foraminiferal (B) species level diversity in western North America. Data is derived from compiled stratigraphic range charts (Figures 2.2–2.4; 3.1–3.4). Ammonite diversity data is compiled from many localities throughout western North America, whereas foraminiferal diversity data is from Haida Gwaii (BC) and the South Barrow #3 core (Alaska). Major declines in ammonite species diversity (A) are evident at six distinct intervals, whereas foraminiferal diversity data (B) only shows two major declines (Kanense Zone and in the later part of the Toarcian). Wh. = Whiteavesi, Fre. = Freboldi, Carl. = Carlottense, Kan. = Kanense, Plan. = Planulata, Crass. = Crassicosta, Hill. = Hillebrandti, Yak. = Yakounensis.  92  Results for the ammonite species diversity analysis (Figure 5.1A) show a steady increase in species diversity throughout the Imlayi–middle Whiteavesi Zones of the Early Pliensbachian. Diversity reaches a high point in the middle Whiteavesi Zone (19 species) and then declines steadily throughout the remaining part of the Early Pliensbachian, reaching a low point in the middle Freboldi Zone (12 species). From the late Freboldi Zone to the middle Kunae Zone, there is a sharp increase and diversity reaches its maximum value of 30 species. Also, over this time frame during the early Kunae Zone, there is an observed offset between the total diversity and singleton-free curves of ~5 species. Throughout the remaining part of the Late Pliensbachian and into the Early Toarcian, diversity undergoes a major decline that consists of two distinct steps. The initial step is an abrupt decline from the middle Kunae to the early Carlottense Zones when diversity drops from 30 to 14 species. The second step is a decline that begins after the middle Carlottense Zone, crosses the Pliensbachian / Toarcian boundary, and reaches a minimum in species diversity of 6 species in the middle Kanense Zone. From the late Kanense Zone and into the Middle Toarcian there is a gradual increase in species diversity that reaches a high point of 15 species in the middle Planulata Zone. Following this peak in diversity in the Middle Toarcian, there is a gradual decline from the late Middle Toarcian into the early Late Toarcian. During this interval, diversity falls to 8 species in the lower Crassicosta Zone and then gradually declines again,  93  reaching low levels (of ~5 species) throughout the Hillebrandti Zone of the Late Toarcian. Throughout the remaining part of the Late Toarcian there is an abrupt rise in diversity in the Yakounensis Zone that reaches a high point of 14 species, with a subsequent decline in the later part of the zone. Regional foraminiferal species diversity data (Figure 5.1B), analyzed independently according to geographic region, shows a much different pattern throughout the Pliensbachian–Toarcian time. The record from Haida Gwaii indicates an initial increase in diversity from ~10 species in the Imlayi Zone to ~40 in the Whiteavesi Zone and then increases, very gradually, to ~47 species throughout the rest of the Pliensbachian. Throughout the Early–Middle Toarcian there is a steady rise in diversity that reaches ~72 species in the late Planulatamiddle Crassicosta Zones. At this point, there is a steady decline to the Late Toarcian where species diversity in the late Yakounensis Zone drops to ~33. Data from the Alaska core (orange line in 5.1B) shows consistent foraminiferal species diversity of ~27 species throughout the Imlayi–late Freboldi Zones of the Early Pliensbachian. Throughout the Kunae Zone there is a distinct rise in diversity that reaches a maximum of 51 species in the late Kunae Zone. In the later part of the Pliensbachian extending across the Pliensbachian / Toarcian boundary and into the early Middle Toarcian, there is a major decline in diversity that is evident as two distinct steps. The first decline is into the Carlottense Zone, where diversity maintains fairly steady levels, decreasing only slightly from 43 to 39 species. The second decline is much more significant dropping from 39 in the late Carlottense Zone to 5 in the lower Planulata Zone. It should be noted that  94  above this interval in the SB #3 Alaska core there is a gap in the core-recovery and therefore it is not possible to obtain foraminiferal species diversity data from the middle Planulata Zone onward. The combined foraminiferal species diversity analysis (Figure 5.2A) constitutes the entire combined stratigraphic range chart from both British Columbia and Alaska. This analysis shows that foraminiferal species diversity maintains a steady increase throughout the Early Pliensbachian, reaching ~70 species in the late Freboldi Zone. Following this there is a small, sharp, decline in the early Kunae Zone to 65 species with a subsequent abrupt increase in diversity throughout the remaining part of the zone where diversity reaches a maximum of 88 species. In the later part of the Pliensbachian–Early Toarcian there is a gradual decline to 73 species in the middle part of the Kanense Zone. From the middle part of the Early Toarcian to end of the Middle Toarcian, species diversity fluctuates slightly but remains at this level. In the Late Toarcian diversity declines, gradually, and reaches ~30 species in the Late Toarcian.  95  Figure 5.2 - Combined Pliensbachian–Toarcian foraminiferal species diversity from Haida Gwaii and the South Barrow #3 core in Arctic Alaska, western North America. Data in (A) is the total combined species from the compiled stratigraphic range chart (Figures 2.2–2.4; 3.1–3.4). Data in (B) does not include species whose stratigraphic ranges are greater than 6 ammonite zones within the Pliensbachian–Toarcian interval. In (A) only small declines in biodiversity are noted in the Kunae, Kanense, and Planulata Zones and a gradual decline in species diversity that occurs throughout the later part of the Toarcian. In (B) declines in species diversity are evident over five potential intervals throughout the Pliensbachian–Toarcian time (red numbers). See Figure 5.1 for abbreviations.  96  An important factor of this diversity analysis is the proportion of longranging taxa compared with the single ended taxa (taxa that contain stratigraphic ranges that end or begin within the measured time interval). As mentioned, taxa that cross both measured boundaries (i.e. Nbt value in Figure 3.7B) are counted as a whole unit, single ended taxa counted as a half, and singletons counted as a third. If there is an over abundance of Nbt taxa, then changes in diversity within a particular interval may be diminished. In the combined foraminiferal species diversity analysis, there is an abundance of foraminifera whose stratigraphic ranges are exceedingly long in comparison with those of ammonite species (Figures 2.2–2.4; 3.1–3.4). Therefore the combined foraminiferal species stratigraphic range chart was re-analyzed without the long-ranging species (omitting species whose stratigraphic ranges are greater than six ammonite zones within the Pliensbachian–Toarcian interval, or are greater than half of the studied interval). The results are presented in Figure 5.2B. The combined foraminiferal species analysis without long-ranging species (Figure 5.2B) shows an increase in diversity from the Imlayi to Whiteavesi Zones, a very slight decrease across the transition between the Whiteavesi and Freboldi Zones, with a subsequent gradual increase in diversity from the Freboldi Zone to the early Kunae Zone. Throughout the Kunae Zone, there is a sharp rise in foraminifera, reaching a maximum of ~40 species in the late part of the zone. Following this apparent peak in foraminiferal diversity there is a large decline into the Early Toarcian that is divided into three phases, an initial drop in the early Carlottense Zone, a subsequent gradual decline throughout the remaining part of  97  the zone and a sharp decline across the Pliensbachian / Toarcian boundary where diversity reaches a low point of 24 species in the middle part of the Kanense Zone. From the late Early Toarcian, there is a steady rise in diversity with one very tenuous decline across the Planulata / Crassicosta zone boundary (number 4? in Figure 5.2). Diversity levels reach a Toarcian peak in the Crassicosta Zone at 37 species. Foraminiferal diversity then declines gradually throughout the Late Toarcian.  5.1.1 Extinction and Origination patterns Rate metrics established by Foote (2000) were used to assess patterns of extinction and origination in ammonoid and foraminiferal species throughout the Pliensbachian and Toarcian. For this analysis, long-ranging foraminiferal species were excluded for reasons previously discussed. Results show that, in general, the derived rates for all four metrics used (i.e. Van Valen with singletons, Van Valen without singletons, per-taxon rate, and estimated per capita rate) had similar values, except for the lower Kanense Zone. Consequently only the Van Valen metric with singletons is discussed (Figure 5.3). Ammonite and foraminiferal species had similar patterns of extinction and origination although the scales were quite different in that extinction rate among ammonites is much higher than in the foraminifera. Throughout much of the Pliensbachian, ammonites showed similar extinction and origination rates that were fairly low (blue line in Figure 5.3). There is an increase in origination in the lower Kunae Zone as well as a slight elevation in both rates that is offset in the late Kunae–early Carlottense Zonal interval. Across the Pliensbachian / Toarcian boundary, covering an interval from the middle of the Carlottense Zone to the 98  middle of the Kanense Zone, there is a very dramatic increase in the extinction rate. Throughout the Middle–Late Toarcian the extinction and origination rates do fluctuate, but generally maintain lower levels. Foraminiferal species data shows low extinction and origination rates throughout much of the Pliensbachian with offset increases in both metrics throughout the entire Kunae Zone interval (Figure 5.3B). Across the Pliensbachian / Toarcian boundary, covering an interval from the middle of the Carlottense Zone to the middle of the Kanense Zone, there is a sharp and marked increase in the extinction rate (green line in Figure 5.3B). Within this interval, origination rate is somewhat elevated but it lags behind the extinction rate considerably. Throughout the Middle Toarcian, rates are of similar magnitude as those experienced in the Late Pliensbachian. In the Late Toarcian, extinction rates are predominantly higher than origination and reach maximum values in the late Hillebrandti and middle Yakounensis zones.  99  Figure 5.3 - Extinction and Origination rates for Pliensbachian–Toarcian ammonite (A) and foraminiferal (B) species in western North America. Rate metrics for foraminifera do not contain species whose stratigraphic ranges are greater than 6 ammonite zones within the Pliensbachian–Toarcian time interval. Data shows accelerated extinction rates in the Kanense Zone for both taxonomic groups that correlates with major declines in species diversity at interval 3 (Figures 5.1, 5.2). See Figure 5.1 for abbreviations.  100  5.2 Geochemical data A total of 1596 samples were collected and analyzed for isotope geochemistry analysis from two temporally constrained stratigraphic sections on Haida Gwaii and the SB #3 core (Appendix B). Samples were analyzed for carbon-isotope (δ13Corg, reported vs. ‰ Vienna PeeDee Belemnite), total organic carbon (reported as wt%), nitrogen-isotope (δ15N, reported vs. ‰ air), and total nitrogen (reported as wt%) analyses. Previous and new work has established the temporal framework within which the geochemical data are presented.  5.2.1 Whiteaves Bay section The Whiteaves Bay section is located on the northern part of Moresby Island in the Skidegate Inlet area of Haida Gwaii (number 16 in Figure 3.5A2) and consists of approximately 115 m of strata, which includes the Ghost Creek, Fannin, and Whiteaves Formations of the Early Jurassic Maude Group. It was originally measured and described by Cameron and Tipper (1985) and subsequently re-measured and calibrated with the Early Pliensbachian–Middle Toarcian part of the zone scheme (Figure 5.4; using data in Smith et al., 1988; Jakobs et al., 1994; Smith and Tipper, 1996; Jakobs, 1997).  101  102  Figure 5.4 - Lithostratigraphy and ammonite biostratigraphy of the Whiteaves Bay section on Central Graham Island, Haida Gwaii (Modified from Smith and Tipper, 1996 and Jakobs, 1997). Fm. = Formation, C. = Carlottense, Pr. = Protogrammoceras, L. = Lioceratoides, P. = Pacificeras. In total, 375 samples were collected from the Whiteaves Bay section (Figure 5.5A, B). Of these samples, twelve from the base of the section were also analyzed for nitrogen-isotope (δ15N) and total nitrogen data (Figure 5.5C, D). From 0 to 5 m in the section δ13C values average –29‰ and then show a gradual shift to ~ –26‰ at 22 m in the section. From 22–28 m, δ13C values shift to ~ – 28‰ with a small positive shift to –25‰ at 25m in the section. Throughout this interval, organic carbon concentrations peak at 3.36% at 14.7 m in the section and average 1.5%. From 29 to 52 m in the section δ13C values fluctuate markedly, but generally show a positive excursion in carbon-isotopes throughout the interval. Throughout this interval there is often a noticeable difference between adjacent samples, where δ13C values have a difference that is greater than 1–3‰. Also within this interval there are several samples that contain carbon-isotope values that are anomalously high and range from –18 to –21‰ (red circles in Figure 5.5A, B). Organic carbon throughout this interval has lower variability, averaging 0.5% with higher amounts reaching ~2% at 30 m in the section. From 54 to 65 m in the section there is a positive excursion, where δ13C values shift from –28‰ (at 54 m) to approximately –26‰ (at 56–57 m) and then return to approximately –28‰ (at 61–64 m). From 66 to 76 m in the section there is a slight negative shift, but δ13C values generally average –27‰ in this interval.  103  From 76 to 77 m in the section δ13Corg values are slightly more positive and average –26‰. From 54 to 77 m in the section, TOC is consistent and averages 0.5% throughout the interval. From 77–83 m there is an abrupt –2‰ negative shift in δ13Corg to values that averages –28‰ for 6 m. This negative shift coincides with a slight increase in TOC to approximately 0.6% near the top of the excursion interval. From 83–102 m (δ13Corg values become more positive averaging –25‰ throughout the interval with a small negative excursion at ~92 m to –26 or –27‰. Organic carbon has lower variability and averages 0.4% throughout the interval. Low resolution δ15N and TN data from the basal 21 m of the section are shown in Figure 5.5C, D. In general δ15N values are positive and range between 0 and 1‰ throughout the interval. However three intervals (from 6–8 m, at 13.5 m and at 21.10 m) show signs of a negative trend where values are slightly negative and do not exceed –0.25‰. Total nitrogen (TN) throughout this interval ranged between 0.12% and 0.20% with two instances of lower TN values that were 0.08% and 0.07% at 1.8 and 17.4 m respectively.  104  105  Figure 5.5 - Geochemistry of the Whiteaves Bay section. (A) δ13Corg (‰ VPDB: Vienna PeeDee Belemnite) data; (B) Total Organic Carbon TOC (wt%) data; (C) δ15N (‰ Air) data; (D) Total Nitrogen TN (wt%) data. Red circles are samples that contain anomalous δ13Corg (greater than -22‰) values. CIE = Carbon isotope excursion, C. = Carlottense. 5.2.1.1 Temporal constraints The basal 27 m of section is assigned to the Ghost Creek Formation and contains species of Tropidoceras, Acanthopleuroceras and Metaderoceras, which are indicative of the Imlayi and Whiteavesi Zones (Figure 5.4). The Fannin Formation conformably overlies the Ghost Creek Formation and is ~51 m thick (Smith and Tipper, 1996). It contains species from many genera that are indicative of the Freboldi, Kunae, Carlottense, and lower Kanense Zones (Smith and Tipper, 1996; Jakobs, 1997). Dactylioceras kanense occurs at 71 m in the section (Figure 5.4) and, along with other species including Dactylioceras aff. comptum, Dactylioceras cf. alpestre, Lioceratoides allifordense, Lioceratoides propinquum, Protogrammoceras cf. paltum, Tiltoniceras antiquum (Figures 4.5, 5.4), indicates the earliest Toarcian (lowermost part of the Kanense Zone). As mentioned previously in Chapter 2.1, the incoming of Dactylioceras marks the base of the Toarcian generally and the species D. comptum occurs at this level in NE Russia (Dagys, 1968). Protogrammoceras cf. P. paltum, Lioceratoides propinquum and Tiltoniceras antiquum occur in Europe and the North America Cordillera, ranging from the Upper Pliensbachian to the Lower Toarcian (Howarth, 1992; Smith and Tipper, 1996). The co-occurrence of these taxa indicates the lower part of the Kanense Zone.  106  Although there are a number of negative excursions to lighter δ13Corg values throughout the section, there is a sustained interval at 77-83 m where values are approximately –29‰ for 6 m. Above the negative CIE interval, from 98–103 m in the section, several ammonite species have been identified which are indicative of the Planulata Zone (Figure 5.4). These include Rarenodia planulata, Leukadiella ionica, Leukadiella aff. helenae and an unidentified species of Cleviceras. At the top of the Whiteaves bay section the Whiteaves Formation is truncated by a fault, which separates it from the overlying Middle Jurassic Yakoun Group. It seems evident that the negative CIE at 77 m in the section occurs within the Early Toarcian at a correlative interval with the negative CIE in Europe.  5.2.2 Yakoun River section The Yakoun River section crops out along a stream-cut of the Yakoun River in Central Graham Island (number 10 in Figure 3.5A1) and is known as one of the most complete Toarcian sections in Haida Gwaii. This section, originally measured and described by Cameron and Tipper (1985) and re-measured by Jakobs (1997), extends from the lower part of the Whiteaves Formation through the Phantom Creek Formation and into Graham Formation (of the Middle Jurassic Yakoun Group). Jakobs et al. (1994) and Jakobs (1997) place zonelevel biostratigraphic constraint for the Whiteaves Formation and the lower part of the Phantom Creek Formation (Figure 5.6). Original measurement of the section indicates that the Whiteaves Formation is ~60 m thick (Jakobs, 1997). However, recent erosion has significantly increased exposure of the lower part (of the section), which now totals ~100 stratigraphic meters. 107  108  Figure 5.6 - Lithostratigraphy and ammonite biostratigraphy of the Yakoun River section on Central Graham Island, Haida Gwaii (Modified from Jakobs, 1997). Ph. = Phantom Creek Formation, Yak. = Yakounensis Zone, Pr. = Protogrammoceras, D. = Dactylioceras, Cl. = Cleviceras, H. = Harpoceras, P. = Peronoceras, Phy. = Phymatoceras, Den. = Denckmannia, Ps. = Pseudomercaticeras, Gr. = Grammoceras, Sph. = Sphaerocoeloceras, Pseu. = Pseudolioceras, Ha. = Hammotoceras. From the initial 300 samples that were collected in 2008 for geochemical analysis from the Yakoun River section, δ13Corg and TOC were determined on 274 samples, and δ13C data were obtained on 26 wood samples (Figure 5.7A, B). A sub-set of 169 samples was also used for δ15N and total nitrogen determinations (Figure 5.7C, D). From 0 to 32 m in the section, δ13Corg values average –24‰ with a small negative shift to ~ –27‰ at 17 m and a small positive shift to ~ – 22‰ at 23 m. From 0–22.5 m, TOC values average 0.9% (wt %) with two singlepoint maxima values of 2% and 2.3% at 7 m and 11 m respectively. From 22.5– 32 m, TOC values decrease, averaging ~0.4%. From 32 m to 43 m in the section there is an abrupt –7‰ negative shift in δ13Corg to values that average –31‰ for 11 m. This pronounced negative shift coincides with an increase in TOC values from 0.4% to 0.7% (from 32–38 m in the section), eventually reaching values that average 1.2% (from 39–41 m in the section), and subsequently show a sharp decrease in TOC to ~0.5% (from 41–43 m). After a return to more positive values of approximately –25‰ at 43 m in the section, values show a steady negative trend From 43 m to 105 m in the section; averaging –25‰ (from 43–55 m), –26‰ (from 55–70 m), –27‰ (from 70–90 m), and –29‰ (from 90–105 m). TOC values remain low throughout this interval, averaging 0.5%. Although δ13Cwood samples are sparse there is broad similarity to 109  the δ13Corg curve, including a large 8.5‰ negative shift to –30‰ at 33 m in the section and a similar pattern in carbon-isotope fluctuation from 47–53 m, 60–61 m, and 71–80 m in the section (Figure 5.7A, B). Nitrogen-isotope and total nitrogen data was generated for 169 samples of the initial sample-set that was collected in 2008 (Figure 5.7C, D). Throughout the entire section, δ15N values range from ~ –1.5 to 0.8‰ with the lightest values occurring in the negative CIE interval. From 0 to 32 m in the section δ15N values average –0.28‰ with occasional samples containing slightly positive values of 0.08–0.5‰. From 32 to 42 m in the section δ15N values are more negative, peaking at –1.5‰ at 32.8 m and averaging –0.6‰ for the interval. From 43–69 m δ15N values are closer to 0, averaging –0.2‰ and peaking at –0.6‰ at 45.4 m. From 70–103 m values average very close to 0.0‰ (at –0.001‰) with occasional small-scale positive and negative trends. Throughout the entire section, total nitrogen values consistently average 0.05%.  110  111  Figure 5.7 - Geochemistry of the Yakoun River section. (A) δ13Corg and δ13Cwood (‰ VPDB: Vienna PeeDee Belemnite) data; (B) Total Organic Carbon TOC (wt%) data; (C) δ15N (‰ Air) data; (D) Total Nitrogen TN (wt%) data; (E) 87Sr/86Sr ratios (from Gröcke et al. (2007). Ph. = Phantom Creek Formation, Yak. = Yakounensis Zone, CIE = Carbon isotope excursion. See Figure 5.6 for lithological legend. Averaged δ13Cwood data from similar stratigraphic horizons are reported with bars showing standard deviation. Hexagonal-shaped symbol denotes stratigraphic position of a zircon-bearing ash dated by Pálfy et al. (1997). Figure adapted from Caruthers et al. (2011). 5.2.2.1  Temporal constraints  Geochemical data for the Yakoun River stratigraphic section are temporally constrained by previous and new work (Figures 4.5; 5.6). In total, all five Toarcian ammonite zones are identified in this section (Jakobs, et al., 1994; Jakobs, 1997). At the base of the section, poorly preserved specimens of Hildaites and Cleviceras occur above, below and within the CIE interval which indicates a Kanense Zone (lowest Toarcian) age for this part of the sequence. Below the excursion interval, Dactylioceras cf. compactum (Dagys), and Protogrammoceras cf. paltum Buckman occur at the base of the section (Figure 5.6). Dactylioceras cf. compactum is known from the earliest Toarcian in NE Russia, at a similar stratigraphic level to D. kanense and D. comptum (Dagys, 1968). Its presence, along with Protogrammoceras cf. paltum, at the base of the Yakoun River section indicates the lowermost part of the Kanense Zone. Unfortunately, the ash bed below the CIE interval, at ~6 m in the section, did not yield enough zircons for U–Pb analysis. Above the negative CIE, from 45–70 m above the base of the section (Figure 5.6), Cleviceras cf. exaratum (Young and Bird), Cleviceras cf. chrysanthemum (Yokoyama), and Hildaites murleyi (Moxon) occur just below  112  Rarenodia planulata Venturi and Lukadiella ionica Renz and Renz (Jakobs, 1997). This is indicative of the Upper Kanense and Planulata Zones, which correlate approximately with the Serpentinum and Bifrons ammonite Zones of NW Europe (Figure 2.1). Previous analysis of belemnite rostra from the Yakoun River section (Figure 5.7E) show 87Sr/86Sr ratios ranging from 0.707138 to 0.707164 in three samples that were recovered from 53 m above the base of the section, ~ 10 m above the CIE (Gröcke et al., 2007). These values plot on the steeply rising (basal Serpentinum Zone) portion of the 87Sr/86Sr curve for NW Europe (star-shaped symbol in Figure 5.8), which is above the extinction event and in the waning part of the negative CIE (McArthur et al., 2000). A U–Pb age date of 181.4 + 1.2 Ma (2σ) from a zircon-bearing ash bed occurring ~29 m above the termination of the CIE interval in the Yakoun River section (hexagonalshaped symbol in Figure 5.7) has contributed to the geochronological time scale for the Planulata/Crassicosta Zone boundary (Pálfy et al., 1997). In summary, the large negative CIE from 32 m to 43 m within the Kanense Zone is concomitant with the negative CIE interval of NW Europe.  113  114  Figure 5.8 - Seawater 87Sr/86Sr for the Late Pliensbachian–Toarcian of Yorkshire, UK (after McArthur et al., 2000). Closed circles are data from McArthur et al. (2000) and grey squares are from Jones et al. (1994). The green star shows the position of three 87Sr/86Sr values at 53 m in the Yakoun River section (from Figure 5.7E). Dashed light-grey line shows stratigraphic position of the extinction interval documented by Caswell et al. (2009). Ammonite zones and subzones (vertical text) after Howarth (1992). Th. = Thouarsense Zone, Tenui. = Tenuicostatum Zone, tenuicostat. = Tenuicostatum Subzone.  5.2.2.2 Higher resolution sampling In 2010, an additional 318 samples were collected at a higher sampling resolution from the large negative CIE interval in the Yakoun River section and analyzed for δ13Corg (Figure 5.9). From 27 to 30.75 m in the section δ13Corg values show a steady gradual positive shift from approximately –26 to –24‰. From 30.80 to 31.15 m there is a large negative shift in carbon-isotopes, in which δ13Corg values decrease at three separate intervals (steps I, II and III in Figure 5.9B) of approximately 2‰, 3‰, and 1–2‰ respectively. This is potentially comparable with similar declines within the Early Toarcian successions of Europe, which are attributed to astronomical precession (Kemp et al., 2005; Hermoso et al., 2009a; Hesselbo and Pienkowski, 2011; Kemp et al., 2011). From 31.2 to 35.6 m in the section δ13Corg values remain very negative, ranging between –30 and –31‰. From 31.2–32.3 m δ13Corg values are the lowest, averaging –31‰ and reaching a minimum value of –32‰. At 35.7 m in the section there is an abrupt –6‰ positive shift in carbon isotope ratios, where δ13Corg values jump from –30‰ to –24‰ over 0.05 cm in the section. From 35.7 to 45.05 m in the section δ13Corg values remain more positive, averaging –25‰ with an occasional negative shift to –28 or –29‰.  115  In comparing the two sets of carbon-isotope data across the negative CIE interval in the Yakoun River section (Figure 5.9), it is now evident that the initial negative shift in carbon-isotopes occurred over three separate steps of –1 to – 3‰, as opposed to the –6‰ shift that is evident in the low resolution data (Figure 5.9A). It is also evident that the negative CIE interval in the high-resolution data is ~6 m thick, while the low-resolution data shows a thickness of ~11 m. The apparent truncation of the negative CIE interval at 35.7 m occurred, coincidentally, with a change in sampling strategy from the cut-bank of the Yakoun River to the riverbed. This retrospectively indicates that sampling occurred in section that was not in place. Within this ‘out of place’ portion of the 2010 data (36–45 m in Figure 5.9B), δ13Corg values are most similar to those in the uppermost Kanense Zone from the lower resolution data (Figure 5.9A), which suggests a small fault or bend in the section between the cut-bank and riverbed.  116  Figure 5.9 - Carbon-isotope data for the negative CIE interval of the Yakoun River comparing: A) the 2008 sample set (Figure 5.7) and B) samples collected at a higher-resolution sample interval in 2010. Higher-resolution sampling suggests an initial negative shift in δ13Corg of ~3‰ that occurs over three separate intervals (B), as opposed to the much larger shift of ~6‰ that is evident in (A). The entire negative CIE interval in (B) was not sampled, values in red are out of place. See Figure 5.6 for lithologic key.  5.2.3 South Barrow #3 core The SB #3 core (number 1 in Figures 3.5A, 3.6) is located on the North Slope of Alaska near the town of Barrow and is part of the National Petroleum Reserve. Drilling of the core commenced and was completed in 1949 reaching a total depth of ~884 m or 2,900 ft (Collins, 1961). Several lithostratigraphic units have been identified in the core, ranging from Triassic–Pleistocene in age. The  117  Lower Jurassic portion of the core is ~275 m (902 ft) thick and encompasses sediments of the Lower Kingak Formation. It is thought to be Hettangian– Toarcian in age (Collins, 1961). Samples collected for isotope geochemistry from a portion of the core that extends from ~675 to 530 m in depth has yielded several Late Pliensbachian and Early Toarcian ammonites and foraminifera (Imlay, 1955; 1981; Tappan, 1955). In total, 603 samples were collected for geochemical analysis from the SB #3 Alaska core and were analyzed for δ13Corg and TOC (Figure 5.10). From 676 to 652 m in the core there is a gradual negative excursion in δ13Corg, where values reach a minimum of –28‰ at 666 m and then steadily increase to –25‰ at 652 m in the core. TOC data for this interval shows heightened values that average 1.12% and reach a maximum of 2.8% at 667 m. At 652 m there is a gap interval of ~7 m where there was no recovery of core from the well. From 645– 629 m there is a steady decrease in δ13Corg with minimum values of –29‰ at 626.9 m and an average TOC of 0.7%. From 624–603 m δ13Corg values are slightly more positive and average –27‰. TOC is slightly higher throughout this interval and averages 1.06% from 624–620 m, 1.10% from 618–611 m, and 0.8% from 608–603 m in the core. At 603 m there is another gap of ~55 m where there was no recovery. From 548 to 534 m there is a slightly negative trend in carbonisotopes, but δ13Corg values average –27‰ throughout the interval. TOC is consistent throughout this interval, averaging 0.7%.  118  119  Figure 5.10 - A) Carbon-isotope chemostratigraphy and B) Total Organic Carbon (wt% TOC) for the U.S. Navy South Barrow #3 well core of Northern Alaska. See Figure 3.5 for location of the well core and Figure 5.6 for lithologic key. ? = ammonite zone-level constraint uncertain, but the earliest Toarcian foraminiferal zone JF9b is probable based on the presence of the species Triplasia kingakensis. Ku–Car = Kunae–Carlottense Zones, Kan–Plan = Kanense– Planulata Zones.  3.2.3.1 Temporal constraints Geochemical data for the SB #3 core are temporally constrained by previous work summarized in Figure 5.11. Imlay (1955; 1981) identifies the species Amaltheus cf. stokesi, Amaltheus cf. margaritatus, Amaltheus engelhardti and Amaltheus sp. from 669 to 630 m in the core, which indicates the Kunae and Carlottense Zones of the Late Pliensbachian (Figure 2.1). From 615 to 540 m in the core Dactylioceras sp., Dactylioceras kanense, and Dactylioceras cf. commune were identified. These species are known to occur in the Kanense and Planulata Zones of the Early and Middle Toarcian. This thereby leaves an interval of approximately 15 m in the core that could be either latest Pliensbachian or earliest Toarcian age. Within this interval, above the last occurrence of the Late Pliensbachian genus Amaltheus and before the first Early Toarcian genus Dactylioceras, there is a small negative CIE where δ13Corg reach a minimum value of –29‰ at 626 m in the core (Figure 5.10). At 628 m in the core the ammonite species Catacoeloceras? sp. juv. is identified and at 618–624 m the foraminifera Triplasia kingakensis is identified at (Imlay, 1955, 1981; Tappan, 1955). This foraminifera species is known as one of the earliest Toarcian foraminifera (Tappan, 1955; Nikitenko and Mickey, 2004) and therefore constrains this portion of the core  120  considerably, indicating that ~624 m in the core is stratigraphically just above the Pliensbachian / Toarcian boundary. Therefore, the small negative CIE at 626 m could potentially correlate with a small negative CIE at the Pliensbachian / Toarcian boundary from NW Europe (Littler et al., 2010). Above the large ~55 m gap in the core, Dactylioceras cf. commune is identified (Imlay, 1955; 1981). This species is known to occur in the uppermost Kanense Zone to Planulata Zone strata in western North America, which suggests that this interval is above the negative CIE interval of NW Europe, and Haida Gwaii (Figures 5.5, 5.7, 5.9). It therefore stands to reason that the Early Toarcian negative CIE interval, if present, would be located within the GAP interval from 603 to 548 m in the SB #3 core.  121  Figure 5.11 - Lithostratigraphy and biostratigraphy for the U.S. Navy South Barrow #3 well core of Northern Alaska. See Figure 3.5 for location of the well core. Stratigraphic position and identification of ammonites is from Imlay (1955; 1981) and Foraminiferal Zones are courtesy of B. Nikitenko (using data in Tappan, 1955). Ammonite Zone scheme is from Smith et al. (1988) and Jakobs et al. (1994).  122  Chapter 6 Discussion 6.1 A multi-phased Pliensbachian–Toarcian mass extinction One of the more pertinent issues that surround this extinction event is gauging its overall magnitude, duration and geographic extent on a global scale. It is currently suggested that the Pliensbachian–Toarcian extinction event in Europe and parts of the Arctic is multi-phased with the two most significant diversity declines occurring at the Pliensbachian / Toarcian boundary and at the Tenuicostatum / Serpentinum zonal boundary in the Early Toarcian (Dera et al., 2010). To date, this type of multi-phased scenario for the Pliensbachian– Toarcian extinction has only been assessed in ammonoid faunas of the northwest Tethys and Arctic domains and, aside from the main-phase of extinction, has not yet been demonstrated as occurring in other taxonomic groups. New paleontological data from western North America provides insight into the magnitude and duration of this multi-phased event. Ammonite and foraminiferal species diversity in western North America suggests a similar multi-phased event (Figure 6.1A). Species diversity in both taxonomic groups declined and reached low points in six separate intervals that correspond to the: 1) middle Whiteavesi–middle Freboldi Zones, 2) late Kunae– early Carlottense Zones, 3) late Carlottense–middle Kanense Zones, 4) late Planulata–early Crassicosta Zones, 5) middle Crassicosta–Hillebrandti Zones and 6) lower–middle Yakounensis Zone. These episodes of decreasing species  123  diversity correlate well with the multi-phased event recorded in combined ammonite data from the northwest European and Arctic domains as documented by Dera et al. (2010) (Figure 6.1B).  Figure 6.1 - Ammonite and foraminiferal species level biodiversity in (A) Western North America (this study) and (B) NW Tethys and Arctic domains (Dera et al., 2010). Figure shows a multi-phased event with major declines occurring over six correlative intervals. In both datasets (A and B), the main phase of extinction is a large progressive decline that begins just before the Pliensbachian / Toarcian boundary and extends into the Early Toarcian where diversity reaches its lowest levels at an interval coeval with the negative CIE. Note: the Middle Toarcian event in Dera et al. (2010) is illustrated here as two separate events (#4, #5). Intervals i & ii = approximate extinction intervals previously identified by Harries and Little (1999) and Caswell et al. (2009), Wh. = Whiteavesi, Fre. = Freboldi, Carl. = Carlottense, Kan., = Kanense, Plan. = Planulata, Crass. = Crassicosta, Hill. = Hillebrandti, Yak. = Yakounensis, Jam. = Jamesoni, Marg. = Margaritatus, Spin. = Spinatum, T. = Tenuicostatum, S. = Serpentinum, Var. = Variabilis, Th. = Thouarsense, D. = Dispansum, P. = Pseudoradiosa, A. = Aalensis, negative CIE = negative carbon-isotope excursion interval in Caruthers et al. (2011). 124  The new dataset from western North America shows a modest decline in ammonite diversity and a correlative, but smaller, decline in foraminiferal diversity in the Early Pliensbachian that begins in the middle of the Whiteavesi Zone and continues to the middle of the Freboldi Zone (number 1 in Figure 6.1A). Throughout this interval, the extinction rates in both taxonomic groups are slightly elevated in comparison with origination (Figure 5.3). This modest decline in species diversity occurs over a similar time frame, and is of similar magnitude, to the decline in ammonite species diversity across the Ibex–Davoei Zone boundary in the dataset from northwest Europe and the Arctic (number 1 in Figure 6.1B). This could therefore suggest a controlling mechanism that is global in extent. Of the five episodes of diversity decline that constitute the multi-phased Pliensbachian–Toarcian extinction described by Dera et al. (2010), all are distinguishable in western North America (numbers 2 to 6 in Figure 6.1). In both datasets, the main phase of extinction consists of two successive phases of decline (number 3 in Figure 6.1): one at the Pliensbachian / Toarcian boundary and the other within the middle part of the Kanense Zone (at a correlative interval with the Tenuicostatum / Serpentinum Zone boundary of northwest Europe; Figure 2.1). At the onset of the main-phase of extinction in western North America, ammonite and foraminiferal diversity reached maximum values of ~30 and 40 species respectively in the Kunae Zone and then began to decline gradually into the Early Toarcian (numbers 2 and 3 in Figure 6.1A). In the combined northwest European and Arctic data (Figure 6.1B), species diversity  125  rebounds in the latest Pliensbachian following the decline in the Margaritatus– Spinatum Zones and then subsequently collapses across the Pliensbachian– Toarcian boundary into the Early Toarcian where ammonite species diversity reaches a minimum value. During this main-phase of extinction in western North America, ammonite and foraminiferal diversity reach minimum values of 5 and 25 species respectively in the middle part of the Kanense Zone (Figure 6.1A). These low levels of Early Toarcian diversity probably resulted from extraordinarily high extinction rates observed in both taxonomic groups in the lowest part of the Kanense Zone (Figure 5.3). One of the major discrepancies related to a global control mechanism for this event is the suggestion that the negative CIE only occurred in the Tethys Ocean area at an interval that is above the main extinction horizon (Wignall et al., 2005). However, more recent work has shown that the negative CIE is recognizable in many correlative Early Toarcian successions that are farremoved from the Tethys Ocean (Al-Suwaidi et al., 2010; Caruthers et al., 2011; Suan et al., 2011; Gröcke et al., 2011; Izumi et al., 2012). This suggests that this phenomenon was most likely global in extent and strongly supports influence by the methane hydrate reservoir. Furthermore, our study shows that the main phase of extinction is a progressive decline in species diversity that begins just before the Pliensbachian / Toarcian boundary and extends into the Early Toarcian where diversity reaches minimum values in three ocean basins (namely the paleo Pacific, Arctic and Tethys oceans) at an interval that is precisely correlative with the negative CIE (Figure 6.1). This apparent synchronicity  126  between diversity minimums during the main-phase of extinction and the negative CIE is a good indication that methane release played an important role in this extinction event, working to escalate its effects. Following the main phase of extinction in the Kanense Zone, ammonite species diversity in both datasets shows a rise into the Planulata Zone of the Middle Toarcian and its European correlative with a subsequent decline into the Hillebrandti Zone of the Late Toarcian. This decline in ammonite species is composed of two separate phases (numbers 4 and 5 in Figure 6.1). Although the extinction rate during phase number 4 is somewhat higher than origination (Figure 5.3), the observed values are considerably lower than those in the Kanense Zone. Within the later phase (number 5 in Figure 6.1), ammonite species diversity maintains lower levels throughout the Hillebrandti Zone of the Late Toarcian with low levels also observed in the correlative zone of Europe. During this phase, extinction and origination rates were nearly identical which resulted in no observable change in diversity. Lastly in the Late Toarcian, the ammonite dataset from northwest Europe and the Arctic shows an abrupt rise in diversity in the lower part of the Dispansum Zone with a final decline in the upper part of the zone; maintaining lower levels throughout the Pseudoradiosa and Aalensis zones (number 6 in Figure 6.1). In western North America, this final event is recognizable and is constrained based on the appearance of two ammonite species, Hammatoceras insigne ranging from the upper Hillebrandti–lower Yakounensis zones and Dumortieria cf. levesquei in the upper part of the Yakounensis Zone (Figure 2.4).  127  In the zone scheme of Northwest Europe (Page, 2003), these two species are noted as the index for the Insigne and Levesquei Subzones encompassing the lower parts of the Dispansum and Pseudoradiosa zones respectively (Figure 2.1). Therefore, the diversity peak occurring in the Yakounensis Zone (Figure 6.1A) is potentially correlative with the peak in the lower part of the Dispansum Zone (Figure 6.1B). The subsequent decline in diversity in the Yakounensis Zone is correlative with the decline in the lower part of the Pseudoradiosa Zone. This thereby constitutes a sixth potential correlative phase of diversity decline evident in North America, Europe and parts of the Arctic in the Late Toarcian. Foraminiferal species diversity, following the main-phase of extinction, shows a similar steady increase from the late Kanense to middle Crassicosta zones with a very small (negligible?) decline across the Planulata–Crassicosta Zone boundary that is correlative with phase number 4 in Figure 6.1. Diversity then declines gradually throughout the Late Toarcian. During the small decline in diversity at the Planulata / Crassicosta Zone boundary there is no observable trend in extinction rate (Figure 5.3) and therefore this does not appear to be a recognizable event (in the foraminiferal data). However, in the later part of the Toarcian, extinction rates are generally higher than origination which could account for the observed decline in diversity.  6.1.1 Correlation with the Karoo–Ferrar magmatism Currently this multi-phased extinction event is attributed to a variety of paleoenvironmental changes that could be related to the Volcanic Greenhouse Scenario but, except for the main-phase of extinction (number 3 in Figure 6.1), are believed to be restricted to the Tethys Ocean area (Dera et al., 2010). These 128  include: sea-level fluctuation affecting restricted basins, saline stratification, warming (or cooling) of seawater, changing water current dynamics and variation in geochemical cycles (McArthur et al., 2000; 2008; Bailey et al., 2003; Suan et al., 2008; 2010; Dera et al., 2010; Dera and Donnadieu, 2012). However, if the six extinction phases are global in extent, it is plausible that the eruption of the Karoo–Ferrar LIP is the underlying controlling factor. When comparing the approximate timing of this multi-phased event to the previously established, and calibrated, time scale for the Pliensbachian and Toarcian Stages (Figure 2.1), the middle Whiteavesi–middle Freboldi Zone event occurred at 186 Ma, the late Kunae–early Carlottense Zone event occurred at 184 Ma, the late Carlottense–middle Kanense Zone event occurred at 183 Ma, the late Planulata–early Crassicosta Zone event occurred between 182 and 181 Ma, the middle Crassicosta–Hillebrandti Zone event occurred between 181 and 180 Ma, and the Yakounensis Zone event occurred between 179–178 Ma. A study by Jourdan et al. (2008) shows eruption ages for magmatism in the Karoo Basin that occur over ~10 My, between 186 and 176 Ma, with the main pulses of magmatism occurring over a window of ~3 Ma from 184 to 181 My (Figure 6.2); with the main volume of basalt emplacement taking place over a 3 to 4.5 Ma window around 180 Ma. A comparison of eruption ages from the Karoo Basin and the multi-phased extinction event reveals a correlation, in that: 1) four of the six pulses of extinction occur within the main-phase of magmatism, 2) the initial Early Pliensbachian decline in species diversity occurs within error range of the onset of Karoo magmatism and 3) the latest Toarcian decline correlates with the  129  later stages of Karoo magmatism (Figure 6.2). This suggests that the eruption of the Karoo–Ferrar LIP was a major underlying factor in the long-term environmental change that resulted in the multi-phased extinction. During the Early Toarcian, a number of compounding factors could have simultaneously created a significant amount of environmental stress which induced episodes of methane release and marine anoxia, and escalated this (main) phase of extinction (as indicated in Hesselbo et al., 2000; 2007; Beerling et al., 2002; Jenkyns et al., 2001; Jenkyns, 2003; 2010; Jourdan et al., 2008; Dera et al., 2010 and references therein).  130  131  Figure 6.2 - Diagram comparing the timing of the multi-phased Pliensbachian– Toarcian extinction with emplacement of the Karoo volcanic province. The 40 Ar/39Ar age probability density distribution diagram and frequency histogram of volcanic rocks in the Karoo Basin are modified from Jourdan et al., (2008) to show its correlation in timing with the known phases of extinction. Figure suggests that all six phases of diversity decline occur within the duration of Karoo magmatism. Baj. = Bajocian.  132  6.2 The negative CIE interval & the long-term carbonisotope record The principal line of evidence that supports the suspected global release of methane hydrate along continental margins is a ~5–7‰ δ13C change, which points to a large perturbation in the global carbon reservoir, identifiable as a negative CIE (Hesselbo et al., 2000; 2007; Kemp et al., 2005; Sabatino et al., 2009; Hermoso et al., 2009a; Suan et al., 2008). At the onset of this study presented herein, there was considerable debate surrounding the global extent and temporal correlation of this excursion. This debate is centered around whether or not the negative CIE, observable in parts of the NW European and Mediterranean Tethys Ocean, is suggestive of: 1) the sudden and global release of methane hydrate within a runaway greenhouse scenario (Hesselbo et al., 2000; 2007), or 2) regional upwelling of 12C-rich bottom water in various parts of the Tethys Ocean (van de Schootbrugge et al., 2005; Wignall et al., 2005; McArthur et al., 2008). The data presented herein from western North America will contribute to this debate by directly comparing the coeval geochemical records from an ocean basin that is geographically far removed from the Tethys Ocean. Haida Gwaii is part of the tectonically displaced Wrangellia terrane and as such its geographic position has varied through time. Ammonite data suggest that during the Late Pliensbachian, Wrangellia was located in the northeastern part of the paleo-Pacific Ocean (number 7 in Figure 1.11A) near the present Canada/US border (at roughly 49º N latitude) at an unknown distance west of the North American margin (Smith, 2006). In the study presented herein, large  133  negative CIEs are identified within the Early Toarcian stratigraphy from the two sampled sections in Haida Gwaii, along the Yakoun River and at Whiteaves Bay (Figures 5.5, 5.7, 6.3). The discovery of chronologically well-constrained negative CIEs from the paleo-Pacific Ocean with similar profiles and magnitude to correlative European successions clearly demonstrates a global control. Of the two postulated mechanisms for the negative CIE, only methane release could reasonably cause such a global response. A defining characteristic, in the European data, is a broad positive shift in δ13C that begins near the Pliensbachian / Toarcian boundary and extends into the Serpentinum Zone, but is interrupted by a pronounced negative excursion at the Tenuicostatum / Serpentinum Zone boundary (Figure 6.3C–E). Data from Haida Gwaii show a somewhat similar perturbation, particularly in the Whiteaves Bay section where there is a more definitive positive trend in carbon-isotope values that begins at the Pliensbachian / Toarcian boundary and is interrupted, with a large negative CIE, about midway through the Kanense Zone strata (Figure 5.5). In the Yakoun River section the Pliensbachian / Toarcian boundary is covered and therefore the onset of a possible positive carbon isotope excursion cannot be determined.  134  135  Figure 6.3 - Correlative Early to Middle Toarcian carbon-isotope data from western North America, NW Europe, and the Mediterranean (modified from Caruthers, 2011). (A and B) Kanense–Planulata Zone data from Whiteaves Bay and the Yakoun River sections on Haida Gwaii, British Columbia Canada; (C) Spinatum–Bifrons Zone data from the Mochras Borehole, Wales; (D) Spinatum–Falciferum Zone data from Sancerre–Couy Borehole, France; (E) Spinatum–Levisoni Zone data from Peniche, Portugal (after Hesselbo et al., 2007); (F) Spinatum–Falciferum Zone data from Yorkshire, UK. Pl. = Pliensbachian, Car. = Carlottense Zone, Spin. = Spinatum Zone. Furthermore, Haida Gwaii isotope data support the broad correlation of the Kanense Zone with the Tenuicostatum and Serpentinum ammonite Zones of the standard NW European scheme and suggests the approximate correlative position of the Tenuicostatum / Serpentinum boundary-equivalent within the Kanense Zone. Ammonites in western North America generally include endemic, pandemic, Tethyan, and Boreal taxa. However Boreal taxa, upon which the NW European ammonite zonal scheme is based, are somewhat rare and are confined to either the North American craton or the northern parts of displaced terranes (Smith and Tipper, 1986; Smith et al., 1988; Jakobs et al., 1994; Smith, 2006). This can cause difficulties when correlating between NW Europe and western North America. The presence of what appears to be an isochronous CIE in both areas is therefore an extremely useful aid in calibrating the North American ammonite zonal scheme with the primary standard scheme of NW Europe. Since publication of the initial Yakoun River carbon-isotope and TOC data from the current project (Caruthers et al., 2011), the negative CIE interval has been identified from other areas that are far-removed from the Tethys Ocean. To date, the negative CIE interval is now identifiable from correlative successions in  136  South America (Al-Suwaidi et al., 2010), Siberia (Suan et al., 2011), and Japan (Gröcke et al., 2011; Izumi et al., 2012). Together, these data strongly support the hypothesis of a sudden and catastrophic perturbation in the global carbon reservoir during the Early Toarcian (Hesselbo et al., 2000); most likely caused by the release of methane hydrate from continental margins.  6.2.1 The long-term carbon-isotope record In order to better assess the correlation of this extinction event with the eruption of the Karoo–Ferrar LIP, it is necessary to consider the effect that this eruption had on the carbon-isotope record. This is partially addressed by the methane release hypothesis in Hesselbo et al. (2000). In this work, Hesselbo et al. (2000) argue that the magnitude of the Early Toarcian negative CIE is too great to have been derived solely from volcanogenic outgassing of CO2 (with a δ13C signature ≈ –7‰), instead they suggest that this negative CIE more likely resulted from an additional release from the continental margin methane clathrate reservoir (δ13C ≈ –60‰) due to volcanogenic CO2-driven warming. However, in a more recent study, Deines (2002) suggests that mantle derived CO2 contains a bimodal δ13C signature that peaks at –5 and –25‰, thereby reaffirming the possibility that volcanogenic outgassing of CO2 could create large perturbations in the δ13C record. Furthermore, as we now understand, there are multiple events of species level diversity decline throughout the Pliensbachian– Toarcian stages that also correlate (temporally) very well with the eruption of the Karoo–Ferrar LIP (Figures 6.1, 6.2). In order to gain a better perspective on the hypothesized environmental change, it is necessary to expand the scope of the  137  carbon-isotope record by observing its trend throughout the entire duration of the Karoo–Ferrar magmatism. To help understand the potential long-term affect of Karoo–Ferrar magmatism, new carbon-isotope data in western North America was compiled along with pre-existing published data (also from Haida Gwaii in western North America) in order to generate a long-term carbon-isotope profile that spans much of the Late Triassic (Norian) to Toarcian time (Figures 6.4, 6.5). Composite data from Williford et al. (2007) shows that in the Late Triassic, δ13C is mostly constant at ~ –29‰, at the Triassic / Jurassic boundary there is a large ~2‰ negative shift that is immediately followed by a large positive ~6‰ excursion that extends throughout the middle part of the Hettangian (Figure 6.4). Following this large positive excursion in carbon-isotope ratios, from the later part of the Hettangian to the Sinemurian, δ13C returns to more negative values that are stabilized at ~ – 31‰ (Figure 6.4).  138  139  Figure 6.4 - Late Triassic to Early Jurassic (Sinemurian) carbon-isotope data for a section at Kennecott Point, Haida Gwaii (after Williford et al., 2007). In the Early Pliensbachian, at the bottom part of the composite Pliensbachian–Toarcian carbon-isotope profile (Figure 6.5), δ13C values are similarly more negative (to Sinemurian values in Figure 6.4) at –29‰ before showing a steady shift to heavier values. From the middle part of the Early Pliensbachian a very interesting pattern develops, in that δ13C values fluctuate abruptly and frequently with a somewhat inconsistent magnitude that ranges ~3– 7‰ (Figure 6.5). This pattern of fluctuation in carbon-isotope values continues throughout the majority of the Pliensbachian, into the middle part of the Toarcian (upper part of the Planulata Zone) before returning to more negative values of ~ –29‰. Interestingly this type of perturbation appears to be recognizable at a global scale. A recent study by Silva et al. (2011) shows a carbon-isotope profile from the Lusitanian Basin of Portugal that contains a similar and correlative pattern of isotopic perturbation within the Early to Late Pliensbachian stratigraphy (Figure 6.6). When comparing this dataset from Portugal to the correlative compiled dataset from western North America (Figure 6.5), a series of small positive and negative shifts are evident. Correlative data shows three positive and five negative CIEs throughout the studied interval (Figure 6.6). Positive CIEs occur within the: Whiteavesi (Ibex) Zone, lower part of the Freboldi (Davoei) Zone, and upper part of the Freboldi Zone. Negative CIEs occur at the Whiteavesi / Freboldi Zonal boundary, within the Freboldi Zone and within the  140  Kunae Zone (Figure 6.6). During the Pliensbachian Portugal was located in the western part of the Tethys Ocean, Haida Gwaii was located in the eastern part of the paleo Pacific Ocean and northern Alaska was located in the paleo Arctic Ocean (Figure 1.11). Therefore this remarkably similar pattern of perturbation in the carbon reservoir is recorded in the coeval strata from three distant and geographically far-removed ocean basins and should be considered a global phenomenon. Perturbation, such as this, in the global carbon reservoir would consequently require a global controlling mechanism.  141  142  Figure 6.5 - Combined δ13Corg values for the Pliensbachian and Toarcian from Haida Gwaii and Arctic Alaska, western North America. Data shows a series of small-scale excursions that extend throughout much of the interval, with a welldefined negative CIE in the Kanense Zone of the Early Toarcian.  Figure 6.6 - A comparison of Early to Late Pliensbachian δ13C values between Peniche, Portugal (from Silva et al., 2011) and western North America. Data shows similar correlative trends in carbon-isotope values between the western Tethys, paleo Pacific Ocean and paleo Arctic Ocean carbon reservoirs. Whit. = Whiteavesi, Car. = Carlottense, Sp. = Spinatum, N.A. = North America, Eur. = Europe. Payne and Kump (2007) examine such large perturbations in δ13C with respect to the severe Permian–Triassic extinction event. In their study they assess the influence of LIP volcanism, oxidation of organic carbon, and oxidation of biogenic methane on the Late Permian–Middle Triassic global carbon record. This record shows a series of large shifts (of ~8‰ δ13Ccarbonate in Figure 6.7)  143  through the Late Permian and Early Triassic that are similar to the variations throughout the Pliensbachian–Toarcian reported here (Figures 6.5, 6.6). This comparison is justified because both intervals have been associated with erupting flood basalt provinces, namely the Siberian Traps in the Late Permian– Middle Triassic and the Karoo–Ferrar LIP in the Pliensbachian–Toarcian, which are hypothesized to have initiated events of enhanced global warming and environmental change (Hallam and Wignall, 1997; Pálfy and Smith, 2000). In their carbon cycle box model, Payne and Kump (2007) used a computer simulation to investigate the potential effect that various environmental disturbances have on the Late Permian–Middle Triassic carbon-isotope record. Throughout this interval, the carbon-isotope record shows a series of large-scale shifts, whereby small negative CIEs are followed by larger positive excursions (Figure 6.7). They argue that, although the release of methane (in the Volcanic Greenhouse Scenario in Figure 1.2) cannot be ruled out for any one excursion, this record of repeated negative CIEs is incompatible with this causal factor for all of the observed negative CIEs (Payne and Kump, 2007). This is because there is insufficient time available between events with which to regenerate the clathrate reservoir (Payne and Kump, 2007). Instead they look at other factors within the Volcanic Greenhouse Scenario (Figure 1.2) as potential causes for these shifts in the carbon-isotope record. They suggest that volcanogenic outgassing of 30,000 GT CO2 during Siberian Trap volcanism would produce the observed pattern (Figure 6.7). As mentioned above, methane hydrate has a δ13C value of approximately –60‰ whereas volcanogenic outgassing of CO2 carries a bimodal  144  δ13C signature of approximately –5 and –25‰ (Deines, 2002). Therefore a negative CIE from volcanogenic outgassing of CO2 (during LIP eruption) would have less of an effect on the carbon-isotope record, and would produce a smaller magnitude negative CIE, than would a release of methane hydrate. Furthermore, Siberian Trap volcanism is thought to occur over a similar interval of time (Figure 1.3B), as with the observed disturbance in the carbon-isotope record, and therefore supports this theory. The Volcanic Greenhouse Scenario also helps to explain the multiple positive excursions that are observed within the Early Triassic record. Regarding these large positive excursions, Payne and Kump (2007) suggest that they were caused by an increase in productivity within the marine environment. As discussed in Chapter 1.1, global warming (from increased volcanogenic CO2 during LIP eruption) causes: 1) increased continental weathering and runoff, which increases the supply of nutrients into the marine environment and, in turn, increases primary productivity (Wignall, 2005), 2) anoxic marine water which enhances the regeneration of phosphate and, this too, intensifies the amount of primary productivity in the water column (Payne and Kump, 2007). This corresponding production of increased biomass preferentially absorbs C12 from the water column and causes the observed positive excursion in the carbonisotope record (Payne and Kump, 2007). However, in their model, they were unable to produce multiple positive and negative excursion couplets, as observed in the Late Permian–Middle Triassic carbon-isotope record (Figure 6.7). Therefore they conclude that the observed pattern in carbon-isotope profile can  145  only be accounted for through several pulses of carbon release from volcanogenic sources.  Figure 6.7 - Composite Late Permian–Middle Triassic carbon-isotope record from southern China (generalized from Payne and Kump, 2007). Data shows a series of excursions, thought to have been largely the result of large-scale outgassing of volcanogenic CO2 during Siberian Trap magmatism. The eruption of the Karoo–Ferrar LIP, on the other hand, is thought to have produced ~9,000 GT CO2 (Beerling and Brentnall, 2007), which is much less than the estimated 30,000 GT CO2 (Payne and Kump, 2007) produced during Siberian Trap volcanism. However the change in isotope ratios in the Early Pliensbachian to Middle Toarcian successions is much less than those recorded in the Late Permian–Middle Triassic (Figures 6.5, 6.6, 6.7). Therefore it is suggested that volcanogenic outgassing of CO2 during the eruption of the Karoo–Ferrar LIP could have produced a series of small perturbations in the  146  carbon-isotope record, reflecting a smaller volume of CO2 release, that began in the middle of the Early Pliensbachian and continued until the Planulata Zone of the Middle Toarcian (Figure 6.5). The large negative CIE occurring in the Early Toarcian Kanense Zone is much larger than the other, smaller, fluctuations in δ13C that occur before and after this event (Figure 6.5). As previously discussed, this negative CIE is thought to have resulted from a catastrophic release from the methane hydrate reservoir. Therefore, a suspected mechanism chain for this event would infer that volcanogenic outgassing of CO2 from the eruption of the Karoo–Ferrar LIP initiated greenhouse conditions and global warming during the Pliensbachian– Toarcian interval. In the Early Toarcian a tipping point was reached, whereby ~5,000 GT of methane hydrate was released along the continental margins. This large injection of methane hydrate greatly accelerated greenhouse conditions and further escalated the extinction of marine organisms.  6.3 Global vs. regional marine anoxia The sudden release of methane hydrate along continental margins is also suspected to have contributed toward a globally extensive body of anoxic marine water, dubbed the T–OAE (Jenkyns, 1988). As discussed previously in Chapter 1, the primary line of evidence supporting the T–OAE stems from geochemical studies in northwest Europe and in parts of the Mediterranean. In these studies, primarily in northwest Europe, isotopic perturbation (Figures, 1.9, 1.10) typically co-occurs with organically enriched shale that, in places, reach as high as 15% TOC (Jenkyns, 1985; 1988; Baudin et al., 1990a,b) and is largely thought to be  147  the result of marine anoxia (Jenkyns, 1988; 2010). In particular, it has been argued that excess burial of organic carbon, resulting from marine anoxia (T– OAE), created a broad positive excursion in the carbon-isotope profile throughout the correlative Tenuicostatum / Serpentinum Zones (Jenkyns and Clayton, 1986; Jenkyns, 1988; Jenkyns et al., 2001). This broad positive excursion is interrupted at a level that is correlative with the Tenuicostatum / Serpentinum zone boundary by the large negative CIE interval (Hesselbo et al., 2000), which is also correlative with the main-phase of marine extinction (number 3 in Figure 6.1). Interestingly, when looking at the averaged carbon-isotope values from sections in western North America (Figure 6.5), a similar trend emerges with a broad positive excursion extending from latest Pliensbachian to the middle of the Planulata Zone that is interrupted by a large negative CIE in the Early Toarcian. However, as previously discussed, a positive excursion in carbon-isotopes may also be derived through a rebounding effect from increased productivity during volcanogenic outgassing of CO2, derived primarily from outgassing during LIP eruption (Payne and Kump, 2007). Therefore, with respect to this particular trend in the carbon-isotope record, the ultimate cause may be difficult to differentiate between the observed effects from outgassing during LIP eruption versus the effects from marine anoxia. Another, more definitive, indication of changing redox conditions in the Tethys Ocean area (to a more anoxic environment) is perturbation in isotopic systems such as nitrogen, manganese, sulfate and molybdenum, also discussed in Chapter 1 (Figures 1.2, 1.10). However, a comprehensive record for these  148  isotope systems outside the Tethys Ocean area is not yet available. Moreover, a clear record of Early Toarcian carbon-isotope perturbation co-occurring with high TOC in areas outside the Tethys Ocean has not yet been established, except for two sections from northern Siberia where TOC reached 2% and 6% within the negative CIE interval (Suan et al., 2011). Consequently, questions remain concerning the actual geographic extent of the T–OAE. Total organic carbon, nitrogen-isotope and TN data compiled for the Pliensbachian and Toarcian of western North America sheds some light on this question.  6.3.1 Total organic carbon Pederson and Calvert (1990) address the importance of organic matter supply to the sea floor in the context of bottom water anoxia, suggesting that increased primary productivity is the first-order controlling factor for the accumulation of organic-rich facies in the modern sediments of the Black Sea and that these sediments are not particularly enriched with organic carbon. This is particularly relevant with respect to the hypothesized global geographic extent of the T-OAE, which is evidenced primarily by a suspected coeval record showing elevated TOC within the Early Toarcian stratigraphy (Jenkyns, 1988; 2010 and references therein). Throughout much of the Pliensbachian and Toarcian (excluding the Early Pliensbachian) in western North America, TOC values are consistently low and range generally between 0.5% and 1.5% (Figure 6.8). These values are more in line with modern hemipelagic shelf sediments (McIver, 1975; Table 1 in Tyson and Pearson, 1991). However, at certain intervals throughout the Pliensbachian– Toarcian, TOC values appear elevated and reach concentrations of ~1.5% to 149  3.5%. These include the: 1) Early Pliensbachian (Imlayi and Whiteavesi Zones), 2) Late Pliensbachian (Kunae Zone), 3) the earliest Toarcian (below the negative CIE interval), and 4) within the negative CIE interval (Figure 6.8). Within the negative CIE interval, concentrations of TOC are much lower than the coeval data from northwest Europe (at ~15% TOC, as referenced above), and are also much lower than the other intervals (of elevated TOC) in western North America. Additionally, many of these intervals with elevated TOC (in Figure 6.8) are often only comprised of a couple data points that contain higher concentrations of organic carbon. This contrasts with the TOC data from the Early Pliensbachian of western North America where TOC values are consistently higher (Figure 6.8). Sabatino et al. (2009) report TOC values in pelagic shales of the western Tethys that are elevated, but much lower (< 2%) than those from northern Europe. They suggest that in deeper water environments a variety of factors could have led to a degradation of organic matter in the water column prior to deposition, resulting in lower organic carbon values. Furthermore, a recent comparison of carbon-isotope data from the Paris (France) and Lusitanian (Portugal) basins demonstrates a negative CIE that is decoupled from black shale deposition (Hermoso et al., 2009b). These studies along with this new TOC data from the paleo-Pacific Ocean therefore suggest that organic carbon concentration seems to be a factor that is controlled locally and should not be used to characterize global marine anoxia during the Early Toarcian. Data from Haida Gwaii does not show significant organic carbon enrichment occurring in  150  the same interval as the negative CIE. While this does not rule-out marine anoxia in the paleo-Pacific Ocean, it also does not support it.  151  152  Figure 6.8 - Combined Total Organic Carbon for Pliensbachian and Toarcian samples from Haida Gwaii and Arctic Alaska, western North America.  6.3.2 Nitrogen-isotope data Early Toarcian geochemical data from northern Europe (British Isles) and southern Europe (Italy) show a pronounced positive excursion in δ15N that generally ranges in magnitude ~3–6‰ (Figure 1.10) and correlates with elevated TOC in the negative CIE interval (Jenkyns et al., 2001). This enrichment in nitrogen-isotope ratios is interpreted by Jenkyns et al. (2001) to represent upwelling of a deoxygenated water mass that had undergone partial denitrification. As discussed previously in Chapter 1.1, degradation (oxidation) of organic matter is normally accomplished through the consumption of oxygen, which produces abundances of carbon dioxide, nitrate, and phosphate (Froelich et al., 1979). However, in times of environmental stress such as in an event of global warming, oxygen becomes less abundant or is absent from the environment (Figure 1.2). The microbially mediated system then begins to consume nitrate in order to continue this process (Froelich et al., 1979; Jenkyns et al., 2001). Its consumption from the water column is referred to as ‘denitrification’. Jenkyns et al. (2001) provides further evidence of the advancement of this process by arguing that in the correlative Exaratum subzone of northwest Europe and Italy, denitrification occurred just before the presence of isorenieratane (Schouten et al., 2000), which is a biomarker compound derived from phototrophic sulfur bacteria. The presence of small pyrite framboids within laminated carbon-rich shales is also indicated from Yorkshire (UK) and the 153  Belluno Trough (Italy), which is further indicative of euxinic conditions in the water column (Wilkin et al., 1996; Bellanca et al., 1999; Jenkyns et al., 2001; Wignall et al., 2005; Jenkyns et al., 2007; Jenkyns, 2010). Jenkyns et al. (2001) therefore conclude that denitrification (and subsequent sulfate reduction) probably resulted from enhanced productivity and that the observed positive excursion in nitrogen-isotopes required the upwelling of 15N-rich water, which was then consumed by organic-walled phytoplankton. This suggestion of euxinic conditions in the northern European epeiric and marginal Tethyan carbonateplatform setting is further supported by sulfate-S isotope data in Gill et al. (2011); showing increases in sulfur isotope ratios of sulfate that are coeval with positive excursions in carbon-isotope values across the suspected anoxic interval. In comparison with European data, nitrogen-isotope data from western North America does not show a positive excursion in δ15N within the Early Toarcian negative CIE interval (Figure 6.9). In fact, δ15N values show a general fluctuation between approximately –1.0‰ and 0.25‰ throughout much of the Toarcian with occasional shifts of +/- 1‰ magnitude at various intervals. Throughout the negative CIE interval, δ15N values show two successive (albeit of small magnitude) negative excursions that reach minimum values of –1.50‰ and –1.27‰ respectively (Figure 6.9). As discussed by Jenkyns et al. (2001), typical Jurassic and Cretaceous marine shale has δ15N values that range between –2.5 to 4‰, which encompasses the observed values from western North America for the Toarcian. In modern marine sediments, δ15N ranges approximately from 4 to 10‰ (Altabet  154  and Francois, 1994; Figure 2 in Robinson et al., 2012) with denitrification (δ15N of NO3-) producing mean values within anoxic water masses that reach 19‰ in the eastern equatorial Pacific Ocean (Tesdal et al., 2013). Therefore it seems to be evident with respect to the Toarcian δ15N data that the northeast paleo-Pacific Ocean was not experiencing denitrification. If denitrification was not occurring, then it is possible that the Toarcian sediments were potentially deposited under more normal redox conditions, which is somewhat contrary to hypotheses suggesting a globally extensive T–OAE. If correct then it could be surmised that the release of methane hydrate and the subsequent acceleration in greenhouse conditions (during the Early Toarcian) created anoxic water in only certain marine environments (Figure 6.10), through mechanisms related to the basin-restriction model of McArthur et al. (2008). These restricted basins, containing anoxic marine water, could have been distributed globally and were most-likely not solely tied to the Tethys Ocean area.  155  156  Figure 6.9 - Combined δ15N values for the Early Pliensbachian and Toarcian from Haida Gwaii, western North America. Figure shows nitrogen-isotope values that are more positive in the Early Pliensbachian and values that are slightly negative to near 0.0‰ throughout much of the Toarcian.  Figure 6.10 - Adapted flow chart for the Volcanic Greenhouse Scenario (modified from Wignall, 2001 and 2005) showing the potential effects of global warming in restricted vs. unrestricted basins (orange diamond). 	
   6.3.2.1 Regional denitrification during the Early Pliensbachian Cameron and Tipper (1985) suggest that the Ghost Creek Formation (including the Early Pliensbachian stratigraphy in Whiteaves Bay) was deposited in a euxinic deeper water basin. This is supported by its black colour, the  157  abundance of pyrite and organic carbon, as well as the presence of oil stains and bitumen. They also noted that benthic fossils and bioturbation were uncommon. Near the top of the Ghost Creek Formation, biodiversity and bioturbation increase, organic carbon seems to decrease (evidenced by a lighter coloured lithology), and it is therefore thought that sediments were deposited under more normal redox conditions (Cameron and Tipper, 1985). These interpretations are further corroborated by the low-resolution geochemical data from the Ghost Creek Formation at Whiteaves Bay. In contrast to the previously discussed data from the Toarcian stratigraphy, in the Early Pliensbachian: 1) the nitrogen-isotope dataset yield δ15N values that are more positive and range generally between 0.0 and 1‰ (Figure 6.9), 2) TOC values are also consistently high and decrease substantially throughout the remaining part of the Pliensbachian–Toarcian interval (Figure 6.8) and 3) organic nitrogen concentrations (measured as Wt % total nitrogen or TN) are also consistently much higher, ranging ~0.075 and 0.2%, in comparison with the Toarcian values that are consistent at ~0.05% (Figure 6.11). Although these Early Pliensbachian δ15N values are more positive than those recorded in the Early Toarcian, it should be noted that they are not as light as Toarcian values in European successions which are thought to indicate denitrification (~3‰ in Jenkyns et al., 2001). However these more positive Early Pliensbachian δ15N values, in conjunction with the lithologic interpretations of Cameron and Tipper (1985), may infer that the sediments of the Ghost Creek Formation were potentially deposited in a suboxic water column.  158  As currently understood, there do not seem to be any accounts of global anoxic marine water in the Early Pliensbachian and therefore this instance might only be regionally extensive within part of the Wrangellia composite terrane. It seems reasonable to hypothesize that this episode of marine anoxia could have resulted from a stratified body of water within a locally restricted basin within the terrane. Perhaps, within this basin, the dynamics were such that water currents were weak or absent and could not adequately circulate the water column, thus producing denitrification and anoxia. Gradually, throughout the Early Jurassic, this basin was filled with coarser sedimentation of the Fannin Formation, during the later part of the Early Pliensbachian, and seawater stratification was no longer possible. The water column then returned to a normal oxidizing environment.  159  160  Figure 6.11 - Combined total nitrogen (TN) values for the Early Pliensbachian and Toarcian from Haida Gwaii, western North America. Data shows much greater levels of nitrogen in the Early Pliensbachian, in comparison with the Toarcian data.  161  Chapter 7 Systematic paleontology The following ammonoid taxa are described using the open nomenclature following Bengston (1988) and Smith and Tipper (1996) and are used to supplement: 1) the previously known biostratigraphy from the Talkeetna Mountains and Haida Gwaii and 2) the known stratigraphic ranges of Pliensbachian–Toarcian ammonoids in western North America, which are used in the extinction analysis of this study. See Figure 4.5 (reproduced below) for the distribution of ammonite species, identified here, according to geographic locality. Many of these taxa are well known to western North America and have been previously described in detail with extensive synonymy lists. Therefore the synonymy list, provided herein for each species, is abbreviated with reference for a more complete list given. Specimens described herein are curated at the University of Montana Paleontology Center (UMPC) and the Pacific Museum of the Earth (part of the Beaty Biodiversity Museum, University of British Columbia).  162  163  Figure 4.5 - Distribution of Pliensbachian and Toarcian ammonite taxa that are systematically described in Chapter 7 of this study, according to locality number at the Hicks Creek and Camp Creek sections (Figure 4.4), the Whiteaves Bay and Yakoun River sections (Figures 5.4, 5.6) and the South Barrow #3 Core (Figure 5.11). Ch. = Chapter, A.A. = Arctic Alaska, and SB3 = South Barrow #3 well core. Class Cephalopoda CUVIER, 1797 Order Ammonoidea ZITTEL, 1884 Suborder Ammonitina HYATT, 1889 Family Oxynoticeratidae HYATT, 1875 Genus Fanninoceras MCLEARN, 1930 Subgenus Fanninoceras MCLEARN, 1930 Type species: Fanninoceras fannini MCLEARN, 1930 Fanninoceras (Fanninoceras) carlottense MCLEARN, 1930 Pl. 7.1, figs 1–6 1884. Sphenodiscus requienianus? D’ORBIGNY.– WHITEAVES, p. 248, pl. 22, fig. 4. 1996. Fanninoceras (Fanninoceras) carlottense MCLEARN.– SMITH & TIPPER, pl. 2, figs. 3–7, text-figs. 30j, 31d (and synonymy therein). Description.– Involute, rapidly expanding form with a very narrow umbilicus, compressed whorl section and acute venter without a keel. Inner whorls have gently sinuous prorsiradiate ribs that project onto the venter and become weak and disappear as shell diameter increases. Material.– Nine specimens from localities 1, 2, & 4 in the Hicks Creek section, Talkeetna Mountains Alaska.  164  Discussion.– F. carlottense is the most involute and stratigraphically highest ranging species of Fanninoceras in western North America and is distinguishable by its characteristically narrow umbilicus. F. carlottense has previously been reported from two isolated localities in the Talkeetna Mountains by IMLAY (1981) but it was not figured or placed within stratigraphic context. Occurrence.– Reported from the East Pacific Realm from South America (Argentina and Chile) to western North America (Nevada, Oregon, British Columbia Canada, and Alaska). Age.– Carlottense Zone (Late Pliensbachian).  Fanninoceras (Fanninoceras) fannini MCLEARN, 1930 Pl. 7.1, figs 7–19 1930. Fanninoceras fannini MCLEARN, p. 4, pl. 1, fig. 3. 1996. Fanninoceras (Fanninoceras) fannini MCLEARN.– SMITH & TIPPER, pl. 3, figs. 1–12; pl. 5, figs. 1, 2, text-figs. 27, 30d–e, 31a–c (and synonymy therein). Description.– An involute form with a moderately wide umbilicus for the genus. Whorls are depressed with a broadly arched venter early in ontogeny becoming more compressed with an acute venter in later whorls. Ribs are well spaced, simple–slightly sinuous, projected onto the venter in early whorls and becoming weak with growth. Outer whorls are smooth. Material.– Twenty five specimens from localities 6, & 8–14 in the Camp Creek section and three specimens from locality 3 in the Hicks Creek section, Talkeetna Mountains Alaska.  165  Discussion.– As discussed by SMITH and TIPPER (1996) there is a strong resemblance between F. fannini and F. carlottense but F. fannini has a larger umbilicus. Occurrence.– F. fannini is common in the East Pacific Realm from South America to southern Alaska. Age.– Primarily known from the Kunae Zone with rare occurrences in the Carlottense Zone (Late Pliensbachian).  Subgenus Charlotticeras SMITH & TIPPER, 1996 Type species: Fanninoceras (Charlotticeras) carteri SMITH & TIPPER, 1996 Fanninoceras (Charlotticeras) cf. maudense SMITH & TIPPER, 1996 Pl. 7.1, figs 20–22 cf.  1996. Fanninoceras (Charlotticeras) maudense SMITH & TIPPER, p. 32, pl.  6, figs. 6–11, text-figs. 30a–b. Description.– Midvolute form with a compressed whorl section, flat flanks, and incipient keel. Umbilical wall is low, sharp, and vertical. Ornamentation is well defined with coarse, sinuous, ribs that are prorsiradiate and project onto the venter to the incipient keel. Material.– Two specimens from the Camp Creek section, one from locality 8 and another from locality 13, Talkeetna Mountains Alaska. Occurrence.– Only known from British Columbia Canada (Haida Gwaii, formerly the Queen Charlotte Islands), and Alaska (Talkeetna Mountains). Age.– Kunae Zone (Late Pliensbachian).  166  Family Dactylioceratidae HYATT, 1867 Genus Dactylioceras HYATT, 1867 Type species: Ammonites communis SOWERBY, 1815 Dactylioceras cf. compactum (DAGIS, 1968) Pl. 7.2, figs 1, 2 cf.  1968. Kedonoceras compactum DAGIS, pl. XI, figs 1–3.  Description.– An evolute cadicone form with a depressed sub-circular whorl and a broad venter baring no keel. Rectiradiate ribs project from the umbilicus to the ventrolateral shoulder and become paired across the venter. Material.– One specimen from Locality #2 in the Yakoun River section, Haida Gwaii British Columbia. Discussion.– This species is originally identified from the Propinquum Zone of northeast Russia. Its occurrence in the Yakoun River section is just above the species Protogrammoceras cf. paltum, which suggests a similar age to the Russian specimen. Occurrence.– Currently this species is only known from Russia and western Canada. Age.– Kanense Zone (Early Toarcian).  Family Amaltheidae HYATT, 1867 Genus Amaltheus DE MONTFORT, 1808  167  Type species: Amaltheus margaritatus DE MONTFORT, 1808 Amaltheus sp. Pl. 7.1, figs 23–25 Description.– An involute form with a compressed whorl section and convex flanks. The acute venter bears a low keel that consists of distinct forwardly directed chevrons. A sharp umbilical wall is slightly undercut. Sinuous primary ribs are prorsiradiate and fade toward the venter. Secondary ribs are much less defined and fade quickly toward the venter. Material.– Two poorly preserved specimens from Locality #14 in the Camp Creek section, Talkeetna Mountains Alaska. Two poorly preserved specimens from Localities #1 & #2 in the South Barrow #3 well core. Discussion.– Specimens resemble the well known Boreal Realm species Amaltheus stokesi (SOWERBY), discussed in detail by HOWARTH (1958). However poor preservation precludes positive identification. IMLAY (1981) also identifies A. stokesi from an isolated locality in the Talkeetna Mountains without stratigraphic context as well as in other intervals within the South Barrow #3 well core. Occurrence.– Amaltheus is distributed throughout the Boreal Realm including: Canada, Russia, northwest Europe, and Alaska. It is also known from localities with mixed Tethyan/Boreal faunas including: Canada, Alaska, Italy, Hungary, Japan, and Russia. Age.– Kunae Zone (Late Pliensbachian).  Family Hildoceratidae HYATT, 1867  168  Genus Arieticeras SEGUENZA, 1885 Type species: Ammonites algovianus OPPEL, 1862. Lectotype and paralectotype (the latter refigured from SCHRÖDER, 1927) designated and figured by WIEDENMAYER (1977). Arieticeras aff. domarense (MENEGHINI, 1867) Pl. 7.1, figs 26–31 1867. Ammonites (Harpoceras) domarensis MENEGHINI, p. 7. 1996. Arieticeras aff. domarense (MENEGHINI).– SMITH & TIPPER, pl. 20, fig. 4; text-fig. 37a (and synonymy therein). 2007. Arieticeras gr. domarense (MENEGHINI).– MOUTERDE, DOMMERGUES, MEISTER & ROCHA, p. 88, pl. 6, fig. 5. Description.– Evolute and slowly expanding with a wide umbilicus, compressed, subquadrate, whorl section and sharp keel. Simple to slightly flexuous rectiradiate ribs are most prominent along the flank and disappear at the shoulder. Material.– Seventeen specimens from localities 7, 8, 14, & 15 in the Camp Creek section, Talkeetna Mountains Alaska. Occurrence.– Known from the northern Mediterranean area, Oregon, Nevada, British Columbia Canada (Haida Gwaii), and Alaska (Talkeetna Mountains). Age.– Kunae Zone (Late Pliensbachian).  Genus Leptaleoceras BUCKMAN, 1918 Type species: Leptaleoceras leptum BUCKMAN, 1918  169  Leptaleoceras? sp. Pl. 7.1, figs 32, 33 Description.– Midvolute to evolute with a compressed whorl section bearing a keel. Ribs are rectiradiate, sinuous and are densely spaced along the flank. The specimen figured on Pl. 7.1 fig. 33, shows incipient bundling of ribs near the umbilicus. Material.– Three specimens from locality 6 in the Camp Creek section, Talkeetna Mountains Alaska. Discussion.– All the specimens are small. The densely spaced ribbing suggests an assignment of Leptaleoceras but the bundling of ribs on one specimen suggests that Canavaria may be represented. Occurrence.– Only known from British Columbia (Haida Gwaii), and Alaska (Talkeetna Mountains). Age.– Kunae Zone (Late Pliensbachian).  Genus Protogrammoceras SPATH, 1913 Subgenus Protogrammoceras SPATH, 1913 Type species: Grammoceras bassanii FUCINI, 1901 Protogrammoceras cf. paltum (Buckman, 1922) Pl. 7.2, figs 3–6  cf.  1922. Paltarpites paltus BUCKMAN, pl. 362A.  1981. Protogrammoceras cf. paltum (Buckman) IMLAY, pl. 12, figs 11, 12.  170  cf.  1992. Protogrammoceras (Protogrammoceras) paltum (Buckman)  HOWARTH, pl. 1, figs 1–3; pl. 2, fig 1 (and synonymy therein). 1996. Protogrammoceras (Protogrammoceras) cf. paltum (Buckman), SMITH & TIPPER, pl. 24, figs 1–4 (and synonymy therein).  Description.– A poorly preserved, evolute (midvolute?) form with a compressed ellipsoidal or ogivale whorl section, wide umbilicus and a keel. Evenly spaced falcoid ribs project along the flank to the ventrolateral shoulder, where they disappear before reaching the keel. Ribs are well pronounced and evenly spaced. Material.– Four specimens in total, three from Localities #5 & #6 in the Whiteaves Bay section and one from Locality #1 in the Yakoun River section on Haida Gwaii. Discussion.– The small size and poor preservation of these specimens precludes positive identification. Occurrence.– This species occurs throughout much of Europe and parts of North America where it has been identified from British Columbia (Frebold, 1970; Thompson and Smith, 1992; Smith and Tipper, 1996), Arctic Canada (Hall and Howarth, 1983) and Alaska (Imlay, 1981). Age.– Carlottense Zone (Late Pliensbachian) to Kanense Zone (Early Toarcian).  Genus Lioceratoides SPATH, 1919  171  Subgenus Lioceratoides SPATH, 1919 Type species: Leioceras? Grecoi FUCINI, 1901, p. 91, pl. 11, figs. 4, 5, by original designation (SPATH, 1919, p. 174). Lioceratoides (Lioceratoides) cf. involutum SMITH & TIPPER, 1996 Pl. 7.1, figs 34–36 cf.  1996. Lioceratoides (Lioceratoides) involutum SMITH and TIPPER, pl. 26,  figs. 2–4; text-fig. 39p.  Description.– A midvolute to involute, rapidly expanding form with a prominent keel, compressed whorl section with flat flanks, and an abrupt umbilical wall. Volution tends to be more midvolute on outer whorls. Ribs are densely spaced, sinuous, and occasionally bundled forming irregularities in appearance. Material.– Two specimens from locality 3 in the Hicks Creek section, Talkeetna Mountains Alaska. Occurrence.– British Columbia Canada (Haida Gwaii) and Alaska (Talkeetna Mountains). Age.– Carlottense Zone (Late Pliensbachian).  Lioceratoides (Lioceratoides) cf. allifordense (MCLEARN, 1930) Pl. 7.2, fig 7 cf.  1930. Harpoceras allifordense, MCLEARN, pl. 2, fig 1.  cf.  1932. Harpoceras allifordense, MCLEARN, pl. 5, figs 1–3.  cf.  1964. Harpoceras allifordense, FREBOLD, pl. 8, fig 5 (holotype refigured).  172  cf.  1996. Lioceratoides (L.) allifordense, SMITH & TIPPER, pl. 26, figs 5, 6.  Description.– Moldic impression of a midvolute form with an oval whorl, wide umbilicus, convex flanks, and narrow venter baring a keel. Fasciculate ribs are coarse with rounded tops on the inner whorls and become more sinuous and flat topped with fine intercalated secondary ribs on the outer whorl. Material.– One specimen from Locality #4 in the Whiteaves Bay, Haida Gwaii. Discussion.– This specimen resembles L. allifordense from a section on Maude Island, Haida Gwaii, in SMITH & TIPPER (1996 pl. 26, fig. 5) with respect to the sinuous fasciculate ribs that become flat topped and intercalated with growth. However, this specimen from Whiteaves Bay is poorly preserved and incomplete and precludes a positive identification. Occurrence.– This species is only known from Haida Gwaii, British Columbia Canada. Age.– Carlottense Zone (Late Pliensbachian) to Kanense Zone (Early Toarcian).  Subgenus Pacificeras REPIN, 1970 Type species: Schloenbachia propinqua WHITEAVES, 1884, p. 274, pl. 33, fig. 2. Lioceratoides (Pacificeras) cf. propinquum (WHITEAVES, 1884) Pl. 7.2, figs 8–10 cf.  1884. Schloenbachia propinqua WHITEAVES, pl. 33, figs 2, 2a.  173  cf.  1996. Lioceratoides (Pacificeras) propinquum (Whiteaves) SMITH &  TIPPER, pl. 28, figs 1–11; pl. 29, fig. 1 (and synonymy therein).  Description.– Evolute with an ogivale whorl section, flat or slightly convex flanks, abrupt umbilical shoulder and a simple keel. Coarse widely spaced ribs are sparse and occasionally paired on inner whorls, becoming more evenly distributed and slightly sinuous with growth. Material.– Two specimens from Localities #1 & #2 in the Whiteaves Bay section, Haida Gwaii. Discussion.– The coarse widely spaced ribs on the inner whorls distinguishes this species, as indicated by SMITH & TIPPER (1996; p. 72). This feature is also evident on similar specimens from Haida Gwaii (SMITH & TIPPER, 1996; pl. 28 figs 5b, 7b, 9a, and 11a). Occurrence.– As indicated in SMITH & TIPPER (1996), this species is only known from the northeast Pacific. Age.– Carlottense Zone (Late Pliensbachian) to Kanense Zone (Early Toarcian).  Genus Tiltoniceras BUCKMAN, 1913 Type species: Tiltoniceras costatum BUCKMAN, 1913, p. 8. Tiltoniceras cf. antiquum (WRIGHT, 1882) Pl. 7.2, figs 11, 12 cf.  1882. Harpoceras antiquum WRIGHT, pl. 57, figs 1, 2 (non figs 3, 4)  174  cf.  1992. Tiltoniceras antiquum (Wright) HOWARTH, Text-fig. 13, pl. 6, figs 1–8  (and synonymy therein). cf.  1996. Tiltoniceras antiquum (Wright) SMITH & TIPPER, pl. 30, figs 1–4;  Text-fig. 39 l, m (and synonymy therein).  Description.– An evolute form with a compressed shell, oval whorl, flat flanks and a narrow venter baring a prominent keel that extends to the ventrolateral shoulder. Inner whorls have coarse ribs that are slightly sinuous and evenly spaced, and become very fine to non-existent quickly with growth. Material.– Two specimens from Localities #3 & #6 in the Whiteaves Bay section, Haida Gwaii. Discussion.– Poor preservation of these specimens precludes positive identification. However in the larger specimen (Pl. 7.2, Fig. 12), the compressed whorl and prominent keel is similar to specimens in HOWARTH (1992; text-fig. 13, pl. 6, figs. 1, 7) and SMITH & TIPPER (1996; pl. 30, fig. 1). The sutures that appear above the outer whorl fragment are thought to be from another poorly preserved ammonite and not part of this specimen. Occurrence.– This species has been described from northwest Europe, Russia, and Haida Gwaii. Age.– Carlottense Zone (Late Pliensbachian) to Kanense Zone (Early Toarcian).  Genus Harpoceras WAAGEN, 1869  175  Type species: Ammonites falcifer SOWERBY, 1820, p. 99, pl. 254, fig. 2.  Harpoceras cf. nitescens (SCHLEGELMILCH, 1976) Pl. 7.8, figs 1 cf.  1976. Harpoceras nitescens SCHLEGELMILCH, pl. 46, fig 1.  Description.– An evolute, compressed whorl fragment with an ellipsoidal whorl section, flat flanks and a narrow carinate venter baring a keel. Ornamentation consists of coarse, gently flexuous and evenly spaced, primary ribs that are surrounded on either side by fine secondaries. Ribs project along the flank, coarsening slightly toward the venter and terminating at the ventrolateral shoulder. Material.– One specimen from Locality #10 in the Yakoun River section, Haida Gwaii. Discussion.– This specimen also resembles H. eseri in SCHLEGELMILCH (1976; pl. 46, fig. 2). However H. eseri has ribs that are occasionally paired at the umbilical shoulder, which is not evident in this specimen and in the material in SCHLEGELMILCH (1976). Occurrence.– This species is known from northwest Europe and Haida Gwaii in western North America. Age.– Planulata Zone (Middle Toarcian).  Genus Cleviceras HOWARTH, 1992  176  Type species: Ammonites exaratus YOUNG & BIRD, 1828, p. 266. Cleviceras exaratum (YOUNG & BIRD, 1828) Pl. 7.3, figs 1, 4, 5 See HOWARTH (1992) and JAKOBS (1997) for a complete list of synonymies. 1828. Ammonites exaratus YOUNG & BIRD, p. 266. 1992. Cleviceras exaratum (Young & Bird) HOWARTH p. 90, pl. 9, figs 2–6; pl. 10, figs 1–8; pl. 11, figs 1–17; pl. 12, figs 1–5; pl. 13, figs. 1, 2; text figs 10, 16, 18C, 19C, 20, 21 (and synonymy therein). 1997. Cleviceras cf. exaratum (Young & Bird) JAKOBS pl. 3, figs. 6, 7, 12, 13; pl. 4, figs. 3, 4 (and synonymy therein).  Description.– An evolute (midvolute?) form with an ellipsoidal whorl section, rounded umbilical and ventrolateral shoulders, wide umbilicus and convex flanks. Falcoid ribs are evenly distributed and project onto the venter fading just prior to the sharp keel. Ribs are prorsiradiate and fine close to the umbilical shoulder and then change direction at approximately mid-flank and become gently flexuous and coarse with rounded tops as they project toward the venter. Material.– Two specimens from Localities #8 & #9 in the Yakoun River section, Haida Gwaii. Discussion.– These specimens strongly resemble C. cf. exaratum in JAKOBS (1997) which (in part) was also collected from the same section. The falcoid ribbing that changes direction, abruptly, at approximately mid-flank strongly resembles that of C. exaratum in HOWARTH (1992, pl. 10, fig. 1–8) and differs  177  from Harpoceras serpentinum and H. falciferum (HOWARTH, 1992, pl. 19, figs 1– 4; pl. 20, figs. 1–11) in that the mid-flank flexure in rib direction (in C. exaratum) is not as pronounced as in H. serpentinum and H. falciferum (e.g. falcoid ribbing as opposed to falcate ribbing). Occurrence.– This species is known from northwestern Europe, Siberia and Japan. In western North America it is known from the Canadian Rockies as well as parts of the Cordilleran region including Haida Gwaii. Age.– upper part of the Kanense Zone (Early Toarcian) to Planulata Zone (Middle Toarcian).  Cleviceras spp. Pl. 7.3, figs 2, 3, 6, 7  Description.– A midvolute form with an ellipsoidal–sub-quadrate whorl, a wide umbilicus, abrupt umbilical shoulder, convex flanks and a broad venter baring a sharp keel. Ribs on inner whorls are coarse, evenly spaced, and sinuous. On the outer whorl ribs become falcoid and are fine from the umbilical shoulder to ¼ flank, where they abruptly change direction become flexuous and coarse with round tops, projecting towards the venter and terminating at the shoulder. Material.– Two specimens from Localities #3 & #6 in the Yakoun River section, Haida Gwaii. Discussion.– The fragmentary nature and poor preservation of this specimen precludes further identification within the genus. This specimen is similar to C.  178  exaratum with respect to the falcoid ribbing on the outer whorl, but differs with respect to the position of the bend in rib direction along the flank. In C. exaratum, this occurs at approximately ½ flank and in this specimen the flexure in rib direction occurs at approximately ¼ flank. When comparing with specimens in HOWARTH (1992, pl. 10, figs 1–8), specimens from Haida Gwaii seem to have a wider umbilicus with coarser ribs on inner whorls. Age.– Kanense Zone (Early Toarcian).  Genus Hildaites BUCKMAN, 1921 Type species: Hildaites subserpentinus BUCKMAN, 1921, p. 217. Hildaites murleyi (MOXON, 1841) Pl. 7.4, figs 1–4 1841. Ammonites murleyii MOXON, pl. 24, figs. 6. 1992. Hildaites murleyi (Moxon) HOWARTH, pl. 30, figs 9, 10; pl. 31, figs 1–8; pl. 32, fig 4 (and synonymy therein). 1994. Hildaites murleyi (Moxon) JAKOBS, SMITH & TIPPER, pl. 1, figs 15–18. 1997. Hildaites murleyi (Moxon) JAKOBS, pl. 5, figs 1–9.  Description.– An evolute–midvolute (?) form with a sub-rectangular to ellipsoidal whorl, round umbilical shoulder, convex flanks, and a broad carinate venter baring a large keel. Coarse primary ribs are sinuous with round tops and extend from the umbilical shoulder across the flank and terminate at the  179  ventrolateral shoulder. Fine sinuous secondary ribs also occur between the primaries, and project onto the venter to the keel. Material.– Three specimens from Localities #5 & #7 in the Yakoun River section, Haida Gwaii. Occurrence.– This species has been described from the Mediterranean area, northwest Europe and South America. In western North America it is known from several localities in the Cordilleran region including Haida Gwaii. Age.– upper part of the Kanense Zone (Early Toarcian).  Hildaites spp. Pl. 7.2, figs 13, 14  Description.– Evolute with a sub-rectangular to sub-quadrate whorl section, convex flanks and a broad carinate or tricarinate (?) venter baring a keel. Ribs are gently sinuous, become more prominent and coarse across the flank toward the venter and project onto the ventrolateral shoulder. Material.– Two specimens from Locality #4 in the Yakoun River section, Haida Gwaii. Discussion.– Poor preservation of these specimens precludes a more positive identification. These specimens resemble H. murleyi with their coarse ribs that are gently sinuous, however, there does not seem to be evidence of fine secondary ribs that project to the keel. Age.– Kanense Zone (Early Toarcian).  180  Hildaites? sp. Pl. 7.2, fig 15  Description.– An evolute form with a sub-quadrate whorl section, convex flank, and a broad carinate venter that bears a keel. Ribs are coarse, well spaced and sinuous along the flank, terminating at the ventrolateral shoulder. Material.– One specimen from Locality #8 in the Whiteaves Bay section, Haida Gwaii. Discussion.– The poor preservation of this specimen preclude a positive identification. The coarse, well-spaced and gently sinuous ribs, suggest its placement within this genus. Age.– Kanense Zone (Early Toarcian).  Genus Leukadiella RENZ, 1913 Type species: Leukadiella helenae RENZ, 1913, p. 587–590. Leukadiella ionica RENZ & RENZ, 1947 Pl. 7.4, figs 5, 6 1947. Leukadiella ionica RENZ & RENZ, pl. 12, figs 5, 7. 1966. Leukadiella ionica Renz & Renz WENDT, pl. 2, fig 3. 1966. Leukadiella ionica Renz & Renz KOTTEK, pl. 13, fig 1. 1969. Leukadiella ionica Renz & Renz GALLITELLI–WENDT, pl. 3, fig 8. 1994. Leukadiella ionica Renz & Renz JAKOBS, SMITH & TIPPER, pl. 2, figs 9, 10.  181  1995. Leukadiella ionica Renz & Renz JAKOBS, figs. 5.1–5.5, 6.1–6.6, 6.9, 6.10 (and synonymy therein). 1997. Leukadiella ionica Renz & Renz JAKOBS, pl. 6, figs 1, 2, 4, 7, 8.  Description.– Evolute with a sub-quadrate whorl section, round umbilical shoulder containing prorsiradiate bullae, slightly convex flank, and a broad venter baring a keel with tubercles lining the ventrolateral shoulder. Ribs are coarse, gently sinuous, and well spaced along the flank. Prorsiradiate ribs extend from the umbilical shoulder at poorly preserved tubercles and bend gently rursiradiately at ¼ to ½ flank, and possibly join at the ventrolateral tubercles. Material.– One specimen from Locality #7 in the Whiteaves Bay section, Haida Gwaii. Discussion.– A characterizing feature of Leukadiella ionica is the presence of forked ribs that rejoin at the ventrolateral tubercles (RENZ and RENZ, 1947; JAKOBS, 1995), which is present in this specimen from Haida Gwaii (first noted in other specimens from Haida Gwaii in JAKOBS, 1995). Occurrence.– This species is primarily known from the Mediterranean area, and is also known from Haida Gwaii. Age.– Planulata Zone (Middle Toarcian).  Genus Podagrosites GUEX, 1973 Type species: Pseudogrammoceras podagrosum MONESTIER, 1921, neotype figured by Guex, 1973, pl. 1, fig. 9.  182  Podagrosites latescens (SIMPSON, 1843) Pl. 7.4, figs 7, 8 1843. Ammonites latescens SIMPSON, p. 54, 55. 1997. Podagrosites latescens (Simpson) JAKOBS, pl. 7, figs 1–4, 7, 8; pl. 8, figs 1, 2, 8–15 (and synonymy therein). Description.– Evolute with a wide umbilicus and an oval whorl containing gently rounded umbilical shoulder, flat to slightly convex flanks, and a carinate venter baring a sharp prominent keel. Ribs are coarse, gently flexuous, and evenly spaced with fine secondary ribs evident on the outer two whorls. Material.– One specimen from Locality #14 in the Yakoun River section, Haida Gwaii. Occurrence.– This species is known from northwest Europe and Haida Gwaii. Age.– Hillebrandti Zone (Late Toarcian).  Family Phymatoceratidae HYATT, 1867 Genus Phymatoceras HYATT, 1867 Type species: Phymatoceras robustum HYATT, 1867, p. 88. Phymatoceras cf. erbaense (HAUER, 1856) Pl. 7.4, figs 11, 12 cf.  1856. Ammonites erbaensis HAUER, pl. 11, fig. 10–14.  cf.  1976. Phymatoceras erbaense (Hauer) SCHLEGELMILCH, pl. 44, fig 6.  cf.  1997. Phymatoceras cf. erbaense (Hauer) JAKOBS, pl. 13, figs 10, 11.  183  Description.– An evolute whorl fragment with a sub-quadrate whorl section and broad venter baring a keel. Gently prorsiradiate ribs are very coarse and well spaced along the flank, projecting slightly onto the venter. Material.– One specimen from Locality #13 in the Yakoun River section, Haida Gwaii. Discussion.– Poor preservation precludes positive identification, however, the very coarse prorsiradiate ribs that project onto the venter is similar to specimens in JAKOBS (1997). Occurrence.– This species is known from northwest Europe and the Cordilleran region of western North America including Haida Gwaii. Age.– Crassicosta Zone (Middle Toarcian) to Hillebrandti Zone (Late Toarcian).  Phymatoceras hillebrandti JAKOBS, 1992 Pl. 7.4, figs 9, 10 1992. Phymatoceras hillebrandti JAKOBS, pl. 13, fig 4; pl. 14, figs 1–5; pl. 15, figs 1–2. 1994. Phymatoceras hillebrandti Jakobs JAKOBS, SMITH & TIPPER, pl. 4, figs. 13, 14, 18–23. 1997. Phymatoceras hillebrandti Jakobs JAKOBS, pl. 14, figs 1–6; pl. 15, figs 1– 12; pl. 16, figs 7–8.  Description.– An evolute fragment with an ellipsoidal whorl, flat flanks with a gently rounded ventrolateral shoulder and a carinate venter baring a keel. Ribs  184  are paired, gently flexuous, and consist of strong primaries with weaker secondaries that are joined at a set of weak tubercles near the umbilical shoulder. Paired ribs project prorsiradiately from the umbilicus, curve gently at mid-flank and diminish in size as they project across the flank, terminating at the ventrolateral shoulder. Material.– One specimen from Locality #15 in the Yakoun River section, Haida Gwaii. Discussion.– The paired ribbing and presence of tubercles in this specimen matches well with material illustrated in JAKOBS (1992; 1997). Occurrence.– This species is only described from Haida Gwaii. Age.– Hillebrandti Zone (Late Toarcian).  Phymatoceras sp. Pl. 7.5, figs 1, 2  Description.– A slowly expanding evolute form with an ellipsoidal–oval whorl section, convex flanks and a broad venter without a keel. Ribs are simple and very coarse. They project rectiradiately along the flank and terminate abruptly at a set of poorly preserved ventrolateral tubercles located at the shoulder. Material.– One specimen from Locality #10 in the Yakoun River section, Haida Gwaii. Discussion.– This specimen is similar to Phymatoceras cf. erbaense in JAKOBS (1997; pl. 13, figs 10–11) with respect to whorl shape and size, as well as  185  the coarse rectiradiate ribs. However, this specimen differs in that it does not possess a keel and has ribs that do not project onto the venter. It also occurs in strata of a slightly older age than previous descriptions of P. cf. erbaense from Haida Gwaii (described from the Crassicosta Zone of the Middle Toarcian by JAKOBS, 1997). Age.– Planulata Zone (Middle Toarcian).  Phymatoceras? sp. Pl. 7.6, figs 1, 2  Description.– An evolute form with a wide umbilicus, compressed ellipsoidal whorl, flat flanks and carinate venter. Poorly preserved rursiradiate bullae project inward from the ventrolateral shoulder. Fine sinuous ribs are well spaced with rounded tops and project from the umbilicus along the flank and terminate at the ventrolateral shoulder. Material.– One specimen from Locality #12 in the Yakoun River section, Haida Gwaii. Discussion.– Poor preservation in this specimen precludes a more definitive identification. Age.– Crassicosta Zone (Middle Toarcian).  Genus Rarenodia VENTURI, 1975 Type species: Rarenodia planulata VENTURI, 1975, p. 13, pl. 1, fig. 7.  186  Rarenodia planulata (VENTURI, 1975) Pl. 7.6, figs 3–6; Pl. 7.7, fig 1; Pl. 7.8, fig. 2 1975. Rarenodia planulata VENTURI, pl. 1, figs 3, 6–9, text-figs 3–5. 1992. Rarenodia planulata Venturi JAKOBS, pl. 3, fig. 9; pl. 4, fig. 1; pl. 5, fig. 1; pl. 7, fig. 1. 1994. Rarenodia planulata Venturi JAKOBS, SMITH & TIPPER, pl. 2, figs 21, 22. 1997. Rarenodia planulata Venturi JAKOBS, pl. 10, figs 1, 2; pl. 11, figs. 1, 2; pl. 12, figs. 1, 2, 5–8 (and synonymies therein).  Description.– An evolute, slowly expanding form with an ellipsoidal–sub rectangular whorl section, an abrupt umbilical shoulder, convex flanks and a broad carinate venter containing a sharp and prominent keel. Simple ribs are paired and project rectiradiately along the flank from a set of tubercles, located along the umbilical shoulder. Tubercles are more pronounced along the inner whorls. Ribs terminate on the ventrolateral shoulder. Material.– Three specimens from Localities #3, #10 & #11 in the Yakoun River section, Haida Gwaii. Occurrence.– This species is known from Italy in the Mediterranean area and the Cordilleran region of western North America, including Haida Gwaii. Age.– Planulata Zone (Middle Toarcian).  Genus Lytoceras SUESS, 1865 Type species: Ammonites fimbriatus SOWERBY 1817.  187  Lytoceras sp. Pl. 7.1, figs 37–39  Description.– An evolute form with a depressed sub-quadrate whorl section, a flattened venter, flat flanks and a well-defined umbilical wall. Prorsiradiate constrictions are well defined along the flank and become rectiradiate to gently rursiradiate across the venter. Ribbing is not present. A poorly preserved but complex suture is evident on the adapical part of the specimen. Material.– One fragment from locality 5 in the Hicks Creek section, Talkeetna Mountains Alaska. Discussion.– The prorsiradiate constrictions that become gently rursiradiate across the venter resemble L. apertum in GEYER (1893; Pl. 8, figs. 4, 5) and L. fimbriatum (SOWERBY) in RAKUS & GUEX (2002; Pl. 3, fig. 2), however the whorl shape is not the same. Our specimen also bears slight resemblance in whorl shape to L. fuggeri (GEYER, 1893; Pl. 8, fig. 8 & 9) but it has a depressed subquadrate whorl shape rather than the more compressed whorl shape of L. fuggeri. Occurrence.– This constitutes the first figured occurrence of Lytoceras from the Pliensbachian of western North America. Age.– Carlottense Zone (Late Pliensbachian).  188  Plate 7.1  189  Plate 7.1 – All figures in (Plates I–VIII) are natural size unless otherwise indicated. UMPC = University of Montana Paleontology Center, PME = Pacific Museum of the Earth (UBC), loc. = ammonite locality. Figs. 1-6:  Fanninoceras (Fanninoceras) carlottense MCLEARN 1-2, UMPC 13356; Carlottense Zone; loc. 4, Hicks Creek section. 3, 5, UMPC 13353; Carlottense Zone; loc. 2, Hicks Creek section. 4, UMPC 13351; Carlottense Zone; loc. 1, Hicks Creek section. 6, UMPC 13352; Carlottense Zone; loc. 2, Hicks Creek section.  Figs. 7-19:  Fanninoceras (Fanninoceras) fannini MCLEARN 7, UMPC 13347; 8, UMPC 13360; 9-10, 13361; 17-18, UMPC 13362; Carlottense Zone; loc. 3, Hicks Creek section. 11, UMPC 13345; Kunae Zone; loc. 8, Camp Creek section. 12, UMPC 13342; Kunae Zone; loc. 8, Camp Creek section. 13-14, UMPC 13343; Kunae Zone; loc. 13, Camp Creek section. 15-16, UMPC 13340; Kunae Zone; loc. 8, Camp Creek section. 19, UMPC 13341; Kunae Zone; loc. 8, Camp Creek section (latex cast).  Figs. 20-22: Fanninoceras (Charlotticeras) cf. maudense SMITH & TIPPER 20-21, UMPC 13331; Kunae Zone; loc. 13, Camp Creek section (21 is X2). 22, UMPC 13332; Kunae Zone; loc. 8, Camp Creek section. Figs. 23-25: Amaltheus sp. 23, UMPC 13308; Kunae Zone; loc. 14, Camp Creek section. 2425, UMPC 13309; Kunae Zone; loc. 14, Camp Creek section (25 is X2).  190  Figs. 26-31: Arieticeras aff. domarense (MENEGHINI) 26-27, UMPC 13318; Kunae Zone; loc. 7, Camp Creek section. 2829, UMPC 13319; Kunae Zone; loc. 8, Camp Creek section. 30, UMPC 13310; Kunae Zone; loc. 15, Camp Creek section. 31, UMPC 13313; Kunae Zone; loc. 14, Camp Creek section (latex cast). Figs. 32-33: Leptaleoceras? sp. 32-33, UMPC 13330; Kunae Zone; loc. 6, Camp Creek section (specimen in Fig. 33 is a cast of the specimen in Fig. 32). Figs. 34-36: Lioceratoides (Lioceratoides) cf. involutum SMITH & TIPPER 34-35, UMPC 13357; Carlottense Zone; loc. 3, Hicks Creek section. 36, UMPC 13358; Carlottense Zone; loc. 3, Hicks Creek section. Figs. 37-39: Lytoceras sp. 37-39, UMPC 13359; Carlottense Zone; loc. 5, Hicks Creek section.  191  Plate 7.2  192  Plate 7.2 Figs. 1-2:  Dactylioceras cf. compactum (DAGIS) 1-2, PME 000111; Kanense Zone; loc. 2, Yakoun River section. (x2)  Figs. 3-6:  Protogrammoceras cf. paltum (BUCKMAN) 3-4, PME 000112; Kanense Zone; loc. 1, Yakoun River section. 5, PME 000113; Carlottense Zone; loc. 6, Whiteaves Bay section. 6, PME 000114; Carlottense Zone; loc. 6, Whiteaves Bay section. (4 is x1.10)  Fig. 7:  Lioceratoides (Lioceratoides) cf. allifordense (MCLEARN) 7, PME 000115; Carlottense Zone; loc. 4, Whiteaves Bay section. (X2)  Figs. 8-10:  Lioceratoides (Pacificeras) cf. propinquum (WHITEAVES) 8-9, PME 000116; Carlottense Zone; loc. 1, Whiteaves Bay section. 10, PME 000117; Carlottense Zone; loc. 2, Whiteaves Bay section. (8-9 are X2)  Figs. 11-12: Tiltoniceras cf. antiquum (WRIGHT) 11, PME 000118; Carlottense Zone; loc. 3, Whiteaves Bay section. 12, PME 000119; Carlottense Zone; loc. 6, Whiteaves Bay section. Figs. 13-14: Hildaites spp. 13-14, PME 000120; Kanense Zone; loc. 4, Yakoun River section. (x2) Figs. 15:  Hildaites? sp.  193  15, PME 000121; Kanense Zone; loc. 8, Whiteaves Bay section.  194  Plate 7.3  195  Plate 7.3 Figs. 1, 4-5: Cleviceras exaratum (YOUNG & BIRD) 1, PME 000122; Kanense Zone; loc. 9, Yakoun River section. 4-5, PME 000123; Kanense Zone; loc. 8, Yakoun River section. (1, 4 are x0.75) Figs. 2-3, 6-7:  Cleviceras spp.  2-3, PME 000124; Kanense Zone; loc. 3, Yakoun River section. 6-7, PME 000125; Kanense Zone; loc. 6, Yakoun River section.  196  Plate 7.4  197  Plate 7.4 Figs. 1-4:  Hildaites murleyi (MOXON) 1, PME 000126; Kanense Zone; loc. 5, Yakoun River section. 2-3, PME 000127; Kanense Zone; loc. 7, Yakoun River section. 4, PME 000128; Kanense Zone; loc. 7, Yakoun River section.  Figs. 5-6:  Leukadiella ionica (RENZ & RENZ) 5-6, PME 001090; Planulata Zone; loc. 7, Whiteaves Bay section.  Fig. 7-8:  Podagrosites latescens (SIMPSON) 7-8, PME 001091; Hillebrandti Zone; loc. 14, Yakoun River section.  Figs. 9-10:  Phymatoceras hillebrandti (JAKOBS) 9-10, PME 001092; Hillebrandti Zone; loc. 15, Yakoun River section.  Figs. 11-12: Phymatoceras cf. erbaense (HAUER) 11-12, PME 001093; Crassicosta Zone; loc. 13, Yakoun River section.  198  Plate 7.5  199  Plate 7.5 Figs. 1-2:  Phymatoceras sp. 1-2, PME 001094; Planulata Zone; loc. 10, Yakoun River section.  (1-2 are x0.75)  200  Plate 7.6  201  Plate 7.6 Figs. 1-2:  Phymatoceras? sp. 1-2, PME 001095; Crassicosta Zone; loc. 12, Yakoun River section. (2 is x2)  Figs. 3-6:  Rarenodia planulata (VENTURI) 3-4, PME 001096; Planulata Zone; loc. 11, Yakoun River section. 5-  6, PME 001097; Planulata (?) Zone; float loc. near 3, Yakoun River section.  202  Plate 7.7  203  Plate 7.7 Fig. 1:  Rarenodia planulata (VENTURI) 1, PME 001098; Planulata Zone; loc. 10, Yakoun River section.  204  Plate 7.8  205  Plate 7.8 Fig. 1:  Harpoceras cf. nitescens (SCHLEGELMILCH) 1, PME 001099; Planulata Zone; loc. 10, Yakoun River section.  Fig. 2:  Rarenodia planulata (VENTURI) 2, PME 002000; Planulata Zone; loc. 10, Yakoun River section.  206  Chapter 8 Conclusions This study uses paleontology and isotope geochemistry to investigate the causes and effects of a well-known second order extinction of marine organisms during the Pliensbachian–Toarcian stages of the Early Jurassic. A primary objective of this work is to compare new data from western North America with previously established records in European successions in order to test hypotheses related to its duration, magnitude and controlling mechanisms. Analysis of the stratigraphic ranges of 206 ammonite and 242 foraminiferal species in North America indicates declines in diversity during six separate intervals throughout Pliensbachian–Toarcian time. These intervals of extinction can be correlated with declines in ammonite species and generic data from a combined dataset of the northwest European and Arctic domains (Dera et al., 2010). This suggests that the multi-phased species level extinction during the Pliensbachian–Toarcian interval was potentially global. Phases of extinction correspond with the: 1) middle of the Early Pliensbachian (middle Whiteavesi– middle Freboldi Zones), 2) middle of the Late Pliensbachian (late Kunae–early Carlottense Zone), 3) Pliensbachian / Toarcian boundary into the Early Toarcian (late Carlottense–middle Kanense Zones), 4) Middle Toarcian (late Planulata– early Crassicosta Zones) and 5) late Middle–early Late Toarcian (middle Crassicosta–Hillebrandti Zones), and 6) Late Toarcian (lower–middle Yakounensis Zone).  207  Species level data from western North America also support previous conclusions suggesting that the main-phase of extinction is a global event during which diversity declined in a gradual fashion beginning at the Pliensbachian / Toarcian boundary and extending into the Early Toarcian (Harries and Little, 1999; Caswell et al., 2009). In western North America, species diversity in both ammonite and foraminifera reach minimum values in the middle part of the Early Toarcian Kanense Zone which coincides with the negative CIE interval described by Caruthers et al. (2011). Recognition of a multi-phased extinction across the paleo Pacific, Arctic and Tethys Ocean basins greatly expands the known geographic extent of this event. Previously, at the onset of this study, it was suggested that only the mainphase of extinction might be global but it is now apparent that other phases of decline could potentially be global requiring a controlling mechanism that is also global in its reach. Previously it was argued that volcanogenic outgassing of CO2 during the eruption of the Karoo–Ferrar LIP initiated greenhouse conditions and caused the extinction of marine organisms (Pálfy and Smith, 2000). Data presented herein support this interpretation. There is a good overlap between the eruption ages of main volume of basalt emplacement in the Karoo Basin (Jourdan et al., 2008) and calibrated ages of declines in species diversity that constitute this multiphased event. Four of the six pulses of extinction occur within the main-phase of magmatism, the Early Pliensbachian decline occurs within error range of the  208  onset of Karoo magmatism, and the Late Toarcian decline coincides with the later stages of magmatism. New carbon-isotope data provided in this research strongly supports the Volcanic Greenhouse hypothesis, which implicates global warming (from volcanogenic outgassing of CO2 during Karoo–Ferrar magmatism) as a primary control mechanism for the Pliensbachian–Toarcian extinction (Pálfy and Smith, 2000). The long-term carbon-isotope record (using a comprehensive dataset of 1596 samples analyzed herein as well as a previously published dataset by Williford et al., 2007) shows isotopic fluctuations over much of the Late Triassic (Norian) to Toarcian interval on Haida Gwaii. Within this record, there is a distinct pattern that develops from the Early Pliensbachian to the Middle Toarcian. Within the entire profile, δ13C values are consistent (at approximately –31‰) throughout much of the Hettangian and into the Sinemurian. Throughout the middle part of the Early Pliensbachian into the middle part of the Middle Toarcian, δ13C values become much more positive and fluctuate abruptly and frequently with a magnitude that ranges ~3–7‰. The long-term carbon-isotope profile is consistent with the carbon cycle box model originally proposed by Payne and Kump (2007) to demonstrate the influence of Siberian Trap magmatism during the Late Permian–Middle Triassic. Payne and Kump’s (2007) model suggests that volcanogenic outgassing of CO2 during LIP eruption produces small negative CIEs that are followed by larger (higher magnitude) positive excursions brought on by increased primary productivity. Several pulses of magmatism would  209  therefore produce the observed pattern in the carbon-isotope profile of the Early Pliensbachian–Middle Toarcian interval. Carbon- and nitrogen-isotope geochemistry is also utilized in this study to address two other mechanisms that are thought to have intensified the extinction during part of its main-phase. It is argued that during the Early Toarcian, correlative with the diversity minimum at the Tenuicostatum / Serpentinum Zone boundary, there was a rapid release from the global methane hydrate reservoir which created a large negative CIE in marine and terrestrial carbon reservoirs (Hesselbo et al., 2000). Previously, the geographic extent of this large negative CIE was uncertain, as it was only known from parts of Northwest Europe and Mediterranean area of the Tethys Ocean. New data from two sections on Haida Gwaii demonstrates an Early Toarcian negative CIE that is of similar magnitude and duration to the excursion evident in Europe. During the Early Toarcian, Haida Gwaii was located in the northeast part of the paleo-Pacific Ocean and therefore this data provides good evidence of a globally controlled perturbation in carbon-isotopes. The two datasets from Haida Gwaii show a similar pattern of perturbation, in that a broad positive shift in δ13Corg is interrupted by a sharp and pronounced negative excursion of a similar magnitude and duration as in the European data. At the onset of the negative CIE, high-resolution data in the Yakoun River section shows a large decline in δ13Corg that is composed of three distinct pulses (ranging 1–3‰). The entire negative CIE interval averages –31‰ for 11 m within the Early Toarcian Kanense Zone. This negative CIE in the Yakoun River section also  210  coincides with TOC values that increase from ~0.4% to ~1.2% and subsequently decrease to ~0.5% for the remainder of the section. In the Whiteaves Bay section, the initial part of the negative CIE occurs in at least two separate pulses of ~ –1 to 2‰, and the entire negative CIE interval averages –28‰ for 6 m in the section. This interval also has TOC values that increase slightly to ~ 0.6%. The apparent difference in stratigraphic thickness of the negative CIE interval between these two sections could potentially have resulted from condensing in the Whiteaves Bay section. Carbon-isotope data was also collected from the South Barrow #3 core of northern Alaska. However, there was an unfortunate gap in core recovery that spanned ~55 m within the Early Toarcian stratigraphy where it is believed that the negative CIE (if present) would occur. The new carbon-isotope data from Haida Gwaii: 1) supports a global control rather than local upwelling mechanism for carbon-isotope perturbation, 2) provides further evidence of cyclic pulsing within the negative CIE interval that has been attributed to astronomical precession (Kemp et al., 2005) and 3) also supports calibration of the Early Toarcian ammonite zonal schemes of western North America, NW Europe and parts of the Mediterranean area. Lastly, the study presented herein uses TOC, nitrogen-isotope and TN data to test the highly debated global extent of an anoxic body of marine water during the Early Toarcian (Jenkyns, 1988), dubbed the T–OAE. The T–OAE is thought to have greatly escalated the observed effects of the main-phase of extinction at the level that is correlative with the Tenuicostatum / Serpentinum zone boundary (Caswell et al., 2009 and references therein). Evidence for the T–  211  OAE stems primarily from northwest and parts of southern Europe where isotopic perturbations in many systems, including nitrogen- (Jenkyns et al., 2001), typically co-occur with organically enriched shale with TOC levels as high as 15% (Jenkyns, 1985; 1988; Baudin et al., 1990a,b). A positive excursion in nitrogenisotope values is thought to signify denitrification in an anoxic water column. High concentrations of TOC in European successions support this by indicating that organic matter was not being degraded or oxidized at the time of deposition. In the Pliensbachian–Toarcian of western North America, concentrations of TOC did not exceed 2% within the negative CIE interval, values that are more typical of modern hemipelagic shelf sediments (McIver, 1975). Also, in the CIE interval and throughout the majority of the Toarcian, δ15N values show no positive excursion but are generally consistent, ranging between 0.0 and –1.0‰. In fact, δ15N values seem to show two successive, small-scale, negative excursions within the negative CIE interval. Therefore this TOC and nitrogen-isotope data does not suggest denitrification within the water column for this part of northeast paleo-Pacific Ocean during the Early Toarcian. Unlike the carbon-isotope excursion, which seems to be globally isochronous, TOC concentration is dependent on local factors and therefore should not be used to characterize global marine anoxia. An alternative model for the T–OAE suggests that the anoxic water was locally restricted to silled basins within northwest Europe (McArthur et al., 2008). This hypothesis is based on the over-abundance of organic carbon that cooccurs with perturbations in many isotopic systems within the Early Toarcian  212  stratigraphy (McArthur et al., 2008 and references therein). New TOC, nitrogenisotope and TN data from western North America support this hypothesis, in that there does not seem to be an over abundance of organic carbon nor evidence of denitrification within the Early Toarcian stratigraphy. Therefore, it is unlikely that there was an anoxic water mass within this part of the paleo Pacific Ocean. It is possible that during this time of intense climate change, anoxic water of the T– OAE could only develop under certain conditions and environments (e.g. mechanisms described in the restricted basin model by McArthur et al., 2008). These restricted basins, containing anoxic masses of water were not necessarily restricted to the Tethys Ocean area. Curiously, however, there does seem to be evidence for denitrification and marine anoxia in the Early Pliensbachian Ghost Creek Formation on Haida Gwaii. When comparing with the entire Toarcian interval, the Ghost Creek Formation at Whiteaves Bay has TOC and TN concentrations that are much higher and also has nitrogen-isotope values that are more positive. This new geochemical data corroborates previous interpretations, which suggest a euxinic basin during deposition of this formation. However, currently there are no accounts of global marine anoxia in the Early Pliensbachian and therefore this instance most likely record a regional occurrence. It is possible that a body of anoxic marine water could have developed from a stratified water column in a restricted basin setting within the Wrangellia composite terrane. Gradually this basin was in-filled and the water column then returned to more normal redox conditions by late Early Pliensbachian time.  213  Therefore, to summarize the results of this work from western North America, several major conclusions can be drawn concerning the magnitude, duration and potential controlling mechanisms of the Pliensbachian–Toarcian mass extinction: 1) It now seems evident that species-level diversity in two taxonomic groups declined globally over six intervals throughout Pliensbachian– Toarcian time; 2) There is a strong correlation in timing between this multiphased event, long-term carbon-isotope perturbation and eruption ages of the Karoo magmatic province. This suggests that volcanogenic outgassing of CO2 during LIP eruption was the preeminent factor driving global warming and mass extinction; 3) The most severe, main-phase, extinction interval begins at the Pliensbachian / Toarcian boundary and continues into the Early Toarcian where global species diversity reach minimum values in the middle part of the Kanense Zone; 4) This main-phase (middle Kanense Zone) extinction and diversity minimum was reinforced significantly by a global release from the methane hydrate reservoir, evidenced by a large coeval negative excursion in carbonisotope values from two successions on Haida Gwaii (the first well-documented record of this excursion outside of the Tethys Ocean area); and 5) The northeast paleo Pacific Ocean was not experiencing denitrification and significant organic carbon enrichment during the main-phase of extinction in the middle part of the Kanense Zone, and therefore cannot be considered anoxic at this time. 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Cephalopoda, Munich, Germany, 1, 329–522.  244  Appendices Appendix A: Diversity and Rate Metrics  245  246  247  248  Appendix B: Geochemical Data  Haida Gwaii Identifier YR 01 YR 02 YR 03 YR 04 YR 05 YR 06 YR07 YR 08 YR 09 YR 10 YR 11 YR 12 YR 13 YR 14 YR 15 YR 16 YR 17 YR 18 YR19 YR 20 YR 21 YR 22 YR 23 YR 24 YR 25 YR 26 YR 27 YR 28 YR 29 YR 30 YR 31 YR 32 YR 33 YR 34 YR 35 YR 36  Amt. (Mg) 4.36 4.41 4.68 4.25 4.35 4.76 4.45 4.07 4.49 4.12 4.26 4.74 4.26 4.87 4.75 4.18 4.38 4.50 4.46 4.70 4.32 4.68 4.28 4.75 4.43 4.32 4.70 4.89 4.27 4.84 4.35 4.62 4.38 4.55 4.29 4.33  Yakoun River 2008 data TOC% δ 13C adj 0.79 -24.02 0.84 -24.05 0.88 -24.10 1.00 -24.34 0.82 -23.42 0.81 -24.71 0.45 -24.30 0.81 -24.88 0.80 -24.64 0.85 -24.44 0.79 -24.27 0.84 -24.49 0.90 -24.27 0.74 -24.54 0.68 -24.87 0.78 -24.57 0.96 -24.59 2.09 -24.92 0.70 -25.21 1.06 -24.09 0.62 -24.44 0.73 -24.89 0.86 -24.66 1.16 -24.16 0.90 -24.49 0.90 -24.56 2.31 -24.76 0.82 -24.76 0.82 -24.93 0.73 -24.95 0.36 -24.66 0.40 -24.29 0.67 -24.61 0.96 -24.23 0.90 -23.89 0.76 -24.19  TN% 0.05  δ 15N -0.74  0.05  -0.17  0.06  -0.85  0.04  -0.14  0.05  -0.06  0.06  -0.08  0.06  0.06  0.05  -0.36  0.05  -0.15  0.04  -0.59  0.05  -0.64  0.05  -0.39  0.05  -0.67  0.04  -0.08  0.05  0.15  0.05  0.15  0.06  -0.04  0.06  -0.15  Ht (m) 0.00 0.40 0.80 1.20 1.60 2.00 2.40 2.80 3.20 3.60 4.00 4.40 4.80 5.20 5.60 6.00 6.40 6.80 7.60 8.00 8.40 8.80 9.20 9.60 10.00 10.40 10.80 11.20 11.60 12.00 12.40 12.80 13.20 13.60 14.00 14.40  249  Haida Gwaii Yakoun River 2008 data TOC% 0.65 0.81 0.50 0.92  δ 13C adj -24.14 -24.87 -25.26 -25.39  Identifier YR 37 YR 38 YR 39 YR 40 YR 41 YR42 YR43 YR 44 YR 45 YR 46 YR 47 YR 48 YR 49 YR 50 YR 51 YR 52 YR 53 YR 54 YR 055 YR 56 YR 57 YR 58 YR 59  Amt. (Mg) 4.62 4.32 4.71 4.63 4.46 4.49 4.86 4.84 4.65 4.27 4.65 4.77 4.28 4.29 4.22 4.72 4.46 4.36 4.29 4.47 4.39 4.33  1.11 0.79 0.92 0.94 0.82 0.91 0.84 0.86 0.80 0.74 0.94 0.31 0.26 0.37 0.37 0.39 0.46 0.44  -26.16 -26.71 -24.60 -24.23 -23.63 -24.06 -24.37 -25.09 -24.00 -24.55 -23.99 -24.68 -24.57 -22.07 -25.10 -23.95 -24.44 -24.63  YR 61 YR 62 YR 63 YR 64 YR 65 YR 66 YR 67 YR 68 YR 69 YR 70 YR 71 YR 72 YR 73 YR 74  4.55 4.67 4.71 4.35 4.56 4.44 4.54 4.64 4.91 4.52 4.52 4.48 4.81 4.83  0.57 0.36 0.34 0.32 0.31 0.44 0.47 0.34 0.30 0.39 0.30 0.49 0.62 0.71  -24.44 -24.73 -24.73 -25.04 -25.24 -25.36 -25.86 -25.52 -25.15 -25.23 -24.81 -24.90 -24.94 -24.77  TN% 0.06  δ 15N 0.18  0.04  -0.44  0.05  -0.02  0.03  -0.39  0.06  -0.26  0.06  -0.37  0.07 0.05 0.05  -0.31 -0.35 -0.92  0.08  -0.38  0.08  -0.02  0.05  -0.48  0.05  -0.26  0.06  0.54  0.05  -0.04  0.04 0.05 0.06 0.05 0.05  -0.29 -0.31 -0.14 -0.69 -0.32  0.05 0.04 0.06  -0.38 -0.46 -0.28  Ht (m) 14.80 15.20 15.60 16.00 16.40 16.80 17.20 17.60 18.00 18.40 18.80 19.20 19.60 21.60 22.00 22.40 22.80 23.20 23.60 24.00 26.00 26.40 26.80 27.00 27.20 27.60 28.00 28.40 28.80 29.00 29.20 29.60 30.00 30.40 30.80 31.20 31.60 32.00 250  Haida Gwaii Yakoun River 2008 data Identifier YR 75 YR 76 YR 77 YR 78 YR 79 YR 80 YR 81 YR 82 YR 83 YR 84 YR 85 YR 86 YR 87 YR 88 YR 89 YR 90 YR 91 YR 92 YR 93 YR 94 YR 95 YR 96 YR 97 YR 98 YR 99 YR 100 YR 101 YR 102 YR 103 YR 104 YR 105 YR 106 YR 107 YR 108 YR 109 YR 110 YR 111 YR 112  Amt. (Mg) 4.66 4.52 4.43 4.40 4.49 4.69 4.68 4.84 4.27 4.57 4.39 4.34 4.30 4.44 4.57 4.58 4.48 4.23 4.81 4.90 4.39 4.64 4.25 4.54 4.55 4.65 4.80 4.65 4.45 4.71 4.29 4.37 4.73 4.60 4.31 4.28 4.43 4.72  TOC% 0.69 0.38 0.99 0.83 0.44 0.65 1.04 0.69 0.64 0.98 0.62 0.97 0.61 0.66 0.47 0.53 0.61 1.26 1.33 1.96 0.91 0.85 0.95 1.28 0.69 0.51 0.60 0.50 0.45 0.41 0.56 0.61 0.49 0.44 0.57 0.45 0.48 0.46  δ 13C adj -24.59 -31.39 -31.24 -31.18 -30.77 -30.64 -31.08 -30.27 -31.56 -31.67 -31.67 -31.77 -31.50 -31.58 -31.37 -31.80 -31.61 -30.31 -29.51 -29.64 -29.89 -30.24 -30.39 -29.42 -31.39 -31.03 -31.22 -31.20 -31.27 -31.22 -31.11 -31.29 -31.14 -29.74 -25.71 -25.57 -24.63 -24.84  TN% 0.06 0.04 0.05 0.05  δ 15N -0.49 -1.50 -0.28 -1.08  0.04 0.06 0.04 0.05 0.06 0.06  -0.88 -0.95 -0.67 -0.05 -0.81 0.08  0.06 0.06 0.05 0.05 0.05 0.05 0.04 0.05 0.05 0.05 0.04 0.04 0.05 0.06  -0.05 -0.62 0.00 -0.72 -0.24 -1.08 -0.85 -1.27 -0.71 -1.10 -0.64 -0.76 0.06 -0.51  0.05 0.05 0.05 0.05 0.04 0.05 0.07 0.06 0.04 0.06  -0.76 -0.46 -0.55 -0.26 -0.87 -0.09 -0.52 -0.07 -0.35 -0.06  Ht (m) 32.40 32.80 33.20 33.60 34.00 34.40 34.80 35.20 35.60 36.00 36.40 36.80 37.20 37.60 38.00 38.40 38.80 39.20 39.60 40.00 40.40 40.80 41.20 41.40 41.60 41.80 42.00 42.20 42.40 42.60 42.80 43.00 43.20 43.40 43.60 43.80 44.00 44.20 251  Haida Gwaii Yakoun River 2008 data Identifier YR 113 YR 114 YR 115 YR 116 YR 117 YR 118 YR 119 YR 120 YR 121 YR 122 YR 123 YR 124 YR 125 YR 126 YR 127 YR 128 YR 129 YR 130 YR 131 YR 132 YR 133 YR 134 YR 135 YR 136 YR 137 YR 138 YR 139 YR 140 YR 141 YR 142 YR 143 YR 144 YR 145 YR 146 YR 147 YR 148 YR 149 YR 150  Amt. (Mg) 4.48 4.83 4.43 4.67 4.36 4.47 4.27 4.65 4.36 4.51 4.36 4.32 4.33 4.28 4.62 4.75 4.55 4.45 4.62 4.29 4.30 4.95 4.65 4.50 4.50 4.38 4.43 4.73 4.51 4.64 4.83 4.72 4.84 4.63 4.63 4.68 4.79 4.28  TOC% 0.51 1.13 0.73 0.94 0.84 0.59 0.59 0.79 0.84 0.88 0.94 0.79 0.81 0.78 0.38 0.52 0.69 0.77 0.65 0.73 0.65 0.76 0.83 0.80 0.75 0.86 0.59 0.48 0.74 0.75 0.57 0.65 0.49 0.47 0.57 0.65 0.51 0.68  δ 13C adj -24.67 -24.75 -24.92 -24.43 -25.11 -25.67 -26.15 -24.96 -25.11 -25.23 -24.93 -25.00 -25.04 -25.00 -26.71 -26.58 -25.70 -26.44 -26.25 -25.75 -25.59 -25.85 -25.65 -25.51 -26.15 -26.74 -27.04 -26.61 -25.75 -26.03 -25.83 -26.06 -26.16 -26.85 -26.41 -24.84 -24.98 -25.30  TN% 0.06 0.05 0.05 0.05 0.04 0.04 0.04 0.05 0.05  δ 15N -0.01 -0.11 -0.01 -0.57 0.15 -0.66 -0.52 -0.50 -0.62  0.05  0.24  0.05  -0.56  0.03  -0.37  0.05  -0.06  0.05  0.02  0.05  0.01  0.05  0.01  0.05  0.04  0.03  -0.07  0.05  -0.01  0.05  -0.44  0.04  -0.79  0.04  0.09  0.05  -0.39  Ht (m) 44.40 44.60 44.80 45.00 45.20 45.40 45.60 45.80 46.00 46.20 46.40 46.60 46.80 47.00 47.20 47.40 47.60 47.80 48.00 48.20 48.40 48.60 48.80 49.00 49.20 49.40 49.60 49.80 50.00 50.20 50.40 50.60 50.80 51.00 51.20 51.60 52.00 52.40 252  Haida Gwaii Yakoun River 2008 data Identifier YR 151 YR 152 YR 153 YR 154 YR 155 YR 156 YR 157 YR158 YR 159 YR 160 YR 161 YR 162 YR 163 YR 164 YR 165 YR 166 YR 167 YR 168 YR 169 YR 170 YR 171 YR 172 YR 173 YR 174 YR 175 YR 176 YR 177 YR 178 YR 179 YR 180 YR 181 YR 182 YR 183 YR 184 YR 185 YR 186 YR 187 YR 188  Amt. (Mg) 4.31 4.84 4.87 4.75 4.25 4.64 4.85 4.51 4.17 4.46 4.40 4.47 4.40 4.54 4.81 4.76 4.77 4.37 4.35 4.38 4.35 4.82 4.67 4.38 4.14 4.95 4.77 4.61 4.29 4.84 4.29 4.64 4.39 4.65 4.83 4.21 4.66 4.30  TOC% 0.73 0.61 0.51 0.80 0.64 0.43 0.64 0.23 0.48 0.60 0.69 0.67 0.56 0.56 0.71 0.51 0.45 0.53 0.65 0.44 0.50 0.32 0.44 0.38 1.46 0.37 0.37 0.47 0.42 0.41 0.47 0.59 0.49 0.35 0.46 0.40 0.36 0.52  δ 13C adj -25.14 -25.40 -24.88 -25.28 -24.81 -26.69 -25.04 -27.02 -26.31 -26.16 -25.63 -25.96 -25.35 -25.10 -25.90 -25.52 -26.31 -26.11 -26.35 -26.61 -27.09 -27.08 -27.46 -27.12 -25.24 -26.48 -26.85 -27.19 -26.25 -26.87 -26.80 -26.77 -26.97 -26.84 -27.66 -26.33 -26.69 -27.04  TN% 0.04  δ 15N -0.61  0.05  -0.20  0.05  0.20  0.05  0.31  0.05  -0.09  0.05  0.14  0.05  -0.25  0.05  -0.04  0.04  -0.50  0.05  -0.49  0.04  -0.11  0.04  -0.77  0.04  -0.50  0.04  0.12  0.04  0.03  0.04  0.32  0.05  -0.06  0.05  0.17  0.04  -0.31  Ht (m) 52.80 53.20 53.60 54.00 54.40 54.80 55.20 55.60 56.00 56.40 56.80 57.20 57.60 58.00 58.40 58.80 59.20 59.60 60.00 60.40 60.80 61.20 61.60 62.00 62.40 62.80 63.20 63.60 64.00 64.40 64.80 65.20 65.60 66.00 66.40 66.80 67.20 67.60 253  Haida Gwaii Yakoun River 2008 data Identifier YR 189 YR 190 YR 191 YR 192 YR 193 YR 194 YR 195 YR 196 YR 197 YR 198 YR 199 YR 200 YR 201 YR 202 YR 203 YR 204 YR 205 YR 206 YR 207 YR 208 YR 209 YR 210 YR 211 YR 212 YR 213 YR 214 YR 215 YR 216 YR 217 YR 218 YR 219  Amt. (Mg) 4.55 4.28 4.72 4.01 4.58 4.80 4.91 4.24 4.34 4.61 4.32 4.44 4.71 4.46 4.32 4.83 4.54 4.38 4.72 4.60 4.59 4.81 4.84 4.17 4.90 4.51 4.83 4.66 4.66 4.29 4.31  TOC% 0.41 0.56 0.88 0.84 0.51 0.43 0.47 0.44 0.47 0.51 0.72 0.69 0.61 0.57 0.71 0.64 0.54 0.51 0.59 0.55 0.73 0.52 0.56 0.51 0.45 0.65 0.56 0.52 0.44 0.44 0.38  δ 13C adj -26.49 -26.05 -25.44 -24.91 -26.97 -26.66 -26.84 -28.95 -27.12 -26.62 -26.85 -26.93 -26.41 -26.49 -26.38 -26.32 -26.33 -26.67 -26.28 -26.44 -28.11 -27.31 -28.30 -27.93 -27.65 -27.78 -27.28 -27.47 -27.15 -28.89 -29.34  YR 220 YR 221 YR 222 YR 223 YR 224 YR 225  4.48 4.54 4.72 4.54 4.68 4.68  0.25 0.51 0.54 0.40 0.41 0.49  -28.97 -27.91 -27.96 -28.13 -28.54 -27.79  TN% 0.05  δ 15N -0.06  0.05  -0.62  0.04  -0.27  0.05  -0.08  0.05  -0.26  0.05  -0.18  0.05  -0.21  0.06  0.37  0.05  0.30  0.06  0.22  0.05  0.13  0.05  -0.01  0.06  -0.15  0.06  0.18  0.05  -0.30  0.06  -0.30  0.05  -0.21  0.05  -0.09  Ht (m) 68.00 68.40 68.80 69.20 69.60 70.00 70.40 70.80 71.20 71.85 72.25 72.65 73.05 73.45 73.85 74.25 74.65 75.05 75.45 75.85 76.25 76.65 77.05 77.45 78.45 78.85 79.25 79.65 80.05 80.45 80.85 82.85 82.88 83.28 83.68 84.08 84.48 254  Haida Gwaii Yakoun River 2008 data Identifier YR 226 YR 227 YR 228 YR 229 YR 230 YR 231 YR 232 YR 233 YR 234 YR 235 YR 236 YR 237  Amt. (Mg) 4.37 4.55 4.51 4.29 4.46 4.47 4.34 4.33 4.40 4.85 4.49 4.71  TOC% 0.49 0.54 0.55 0.37 0.63 0.38 0.37 0.35 0.32 0.38 0.53 0.55  δ 13C adj -28.03 -27.73 -28.05 -28.43 -28.73 -28.54 -27.62 -27.60 -27.79 -27.80 -27.55 -27.88  TN%  δ 15N  0.06  -0.14  0.05  0.24  0.06  -0.01  0.05  0.05  0.05  -0.04  0.05  0.31  YR 239 YR 240 YR 241 YR 242 YR 243 YR 244 YR 245 YR 246 YR 247 YR 248 YR 249 YR 250 YR 251 YR 252 YR 253 YR 254 YR 255 YR 256 YR 257 YR 258 YR 259 YR 260 YR 261 YR 262 YR 263  4.53 4.53 4.53 4.82 4.68 4.45  0.30 0.37 0.41 0.39 0.19 0.27  -28.49 -29.54 -29.79 -29.84 -28.90 -29.37  0.05  -0.02  0.03  -0.53  0.04  -0.54  0.05  -0.07  0.05  -0.03  0.05  0.05  0.05  0.15  0.05  -0.19  0.06  0.37  0.06  -0.14  0.05  -0.07  0.06  0.23  0.06  0.31  4.75 4.62 4.57 4.73 4.51 4.46 4.69 4.54 4.52 4.25 4.77 4.31 4.43 4.66 4.26 4.83 4.41 4.35  0.52 0.43 0.47 0.57 0.52 0.38 0.52 0.37 0.46 0.58 0.54 0.62 0.38 0.67 0.52 0.64 0.45 0.42  -29.74 -28.98 -29.57 -28.96 -29.07 -29.02 -28.97 -29.86 -29.56 -29.27 -29.33 -29.74 -29.32 -29.30 -29.14 -29.20 -29.45 -29.21  Ht (m) 84.88 85.28 85.68 86.08 86.48 86.88 87.28 87.68 88.08 88.48 88.88 89.18 89.58 89.98 90.38 90.78 91.18 91.58 91.98 92.38 92.78 93.18 93.58 93.98 94.38 94.78 95.18 95.58 95.98 96.38 96.78 97.18 97.58 97.98 98.38 98.78 99.18 99.58 255  Haida Gwaii Yakoun River 2008 data Identifier YR 264 YR 265 YR 266 YR 267 YR 268 YR 269 YR 270 YR 271 YR 272 YR 273 YR 274  Amt. (Mg) 4.41 4.75 4.49 4.68 4.61 4.29 4.49 4.84 4.20 4.64 4.70  TOC% 0.54 0.35 0.41 0.52 0.61 0.59 0.60 0.46 0.52 0.41 0.50  δ 13C adj -29.10 -29.19 -29.17 -29.32 -29.27 -29.24 -29.06 -29.36 -28.60 -29.12 -29.21  TN%  δ 15N  0.06  0.33  0.06  0.31  0.06  0.30  0.05  -0.28  0.06  -0.07  Ht (m) 99.98 100.38 100.78 101.18 101.58 101.98 102.38 102.78 103.18 103.58 103.98  256  Haida Gwaii Yakoun River 2008 wood data Identifier YR41W YR41W YR72W YR77W YR77W YR01W YR02W YR08W YR03W YR04W YR05W YR06W YR07W YR10W YR09W YR11W YR12W YR13W YR14W YR15W YR16W YR17W YR20W YR18W YR19W YR21W  Amt. (Mg) 0.046 0.061 0.053 0.059 0.050 0.048 0.048 0.045 0.043 0.046 0.043 0.048 0.046 0.057 0.052 0.054 0.050 0.045 0.045 0.044 0.046 0.043 0.046 0.048 0.044 0.046  δ 13Cadj -21.19 -20.53 -29.10 -29.74 -29.94 -26.20 -22.62 -22.13 -25.20 -22.37 -24.90 -22.16 -24.46 -23.02 -23.80 -23.72 -20.55 -22.41 -28.84 -25.56 -23.52 -25.79 -26.09 -23.67 -25.72 -24.30  Ht (m) 16.40 16.40 31.20 33.20 33.20 47.80 49.20 49.20 49.90 50.00 51.00 51.00 51.40 51.40 53.40 60.50 60.80 61.50 71.20 71.67 71.72 71.85 79.05 80.25 80.50 90.38  257  Haida Gwaii Yakoun River 2010 data Identifier YR 275 YR 276 YR 277 YR 278 YR 279 YR 280 YR 281 YR 282 YR 283 YR 284 YR 285 YR 286 YR 287 YR 288 YR 289 YR 290 YR 291 YR 292 YR 293 YR 294 YR 295 YR 296 YR 297 YR 298 YR 299 YR 300 YR 301 YR 302 YR 303 YR 304 YR 305 YR 306 YR 307 YR 308 YR 309 YR 310  Amt. (Mg) 3.454 3.532 3.739 3.088 3.078 3.732 3.001 3.221 3.418 3.181 3.555 3.182 3.057 3.149 3.582 3.811 3.825 3.802 3.758 3.239 3.419 3.167 3.259 3.331 3.761 3.131 3.561 3.306 3.529 3.549 3.557 3.666 3.292 3.355 3.342 3.262  TOC% 0.35 0.41 0.30 0.35 0.44 0.36 0.35 0.35 0.35 0.35 0.33 0.32 0.31 0.36 0.33 0.32 0.35 0.34 0.41 0.34 0.38 0.39 0.48 0.52 0.51 0.66 0.58 0.35 0.59 0.57 0.53 0.53 0.51 0.58 0.65 0.47  δ 13Cadj -25.00 -25.41 -25.13 -26.11 -25.47 -25.80 -25.90 -25.40 -25.27 -25.20 -25.40 -25.04 -25.66 -25.23 -25.18 -25.47 -24.89 -24.81 -25.05 -25.35 -25.06 -24.52 -24.37 -24.85 -24.93 -24.54 -24.59 -25.20 -24.17 -24.39 -24.65 -24.63 -24.70 -24.40 -26.20 -25.00  Ht (m) 27.20 27.30 27.40 27.50 27.60 27.70 27.80 27.90 28.00 28.10 28.20 28.30 28.40 28.50 28.60 28.70 28.80 28.90 29.00 29.10 29.20 29.30 29.40 29.50 29.60 29.70 29.80 29.90 30.00 30.05 30.10 30.15 30.20 30.25 30.30 30.35 258  Haida Gwaii Yakoun River 2010 data Identifier YR 311 YR 312 YR 313 YR 314 YR 315 YR 316 YR 317 YR 318 YR 319 YR 320 YR 321 YR 322 YR 323 YR 324 YR 325 YR 326 YR 327 YR 328 YR 329 YR 330 YR 331 YR 332 YR 333 YR 334 YR 335 YR 336 YR 337 YR 338 YR 339 YR 340 YR 341 YR 342 YR 343 YR 344 YR 345 YR 346  Amt. (Mg) 3.749 3.551 3.957 3.663 4.149 3.53 3.744 3.937 3.773 6.359 4.264 3.453 6.155 6.616 5.232 5.971 3.632 4.185 3.724 3.043 3.994 3.564 3.643 3.123 3.875 3.418 3.432 3.142 3.614 3.763 3.504 3.879 3.936 3.381 3.768 3.374  TOC% 0.40 0.47 0.44 0.42 0.47 0.43 0.40 0.39 0.40 0.12 0.16 0.24 0.21 0.22 0.13 0.19 0.27 0.24 0.28 0.56 0.17 0.94 0.23 0.60 0.72 0.46 0.51 0.69 1.07 0.99 0.88 0.78 0.50 0.95 0.68 0.76  δ 13Cadj -25.23 -26.29 -26.32 -25.79 -25.89 -25.53 -26.37 -25.95 -26.23 -29.73 -29.47 -29.65 -29.71 -29.72 -29.50 -29.83 -30.21 -30.85 -31.15 -31.48 -30.12 -32.07 -30.85 -31.69 -31.95 -31.12 -31.51 -31.56 -31.24 -31.21 -31.00 -31.06 -30.89 -32.07 -31.66 -31.94  Ht (m) 30.40 30.45 30.50 30.55 30.60 30.65 30.70 30.75 30.80 30.85 30.90 30.95 31.00 31.05 31.10 31.15 31.20 31.25 31.30 31.35 31.40 31.45 31.50 31.55 31.60 31.65 31.70 31.75 31.80 31.85 31.90 31.95 32.00 32.05 32.10 32.15 259  Haida Gwaii Yakoun River 2010 data Identifier YR 347 YR 348 YR 349 YR 350 YR 351 YR 352 YR 353 YR 354 YR 355 YR 356 YR 357 YR 358 YR 359 YR 360 YR 361 YR 362 YR 363 YR 364 YR 365 YR 367 YR 368 YR 369 YR 370 YR 371 YR 372 YR 373 YR 374 YR 375 YR 376 YR 377 YR 378 YR 379 YR 380 YR 381 YR 382 YR 383  Amt. (Mg) 3.879 3.721 3.619 4.056 3.538 3.435 3.561 3.473 3.735 3.117 3.382 3.775 3.967 3.967 4.041 3.374 3.666 3.356 3.518 3.46 3.449 3.063 3.249 3.652 3.67 3.439 3.435 3.171 4.095 3.165 3.661 3.311 3.51 3.658 3.557 3.774  TOC% 0.76 0.59 0.90 0.56 0.61 0.67 0.90 0.73 0.62 0.75 0.55 0.52 0.62 0.43 0.51 0.49 0.52 0.36 0.72 0.58 0.55 0.56 0.61 0.76 0.82 0.91 0.98 0.63 0.45 0.55 0.55 0.37 0.30 0.29 0.41 0.39  δ 13Cadj -31.62 -31.39 -31.59 -30.75 -30.85 -30.39 -31.09 -30.96 -30.83 -30.33 -30.67 -30.95 -30.09 -31.40 -31.23 -31.32 -31.38 -31.03 -31.26 -31.26 -31.47 -31.24 -31.13 -31.31 -30.76 -31.43 -31.22 -31.51 -31.34 -31.46 -31.72 -30.22 -31.02 -31.23 -31.16 -31.47  Ht (m) 32.20 32.25 32.30 32.35 32.40 32.45 32.50 32.55 32.60 32.65 32.70 32.75 32.80 32.85 32.90 32.95 33.00 33.05 33.10 33.15 33.20 33.25 33.30 33.35 33.40 33.45 33.50 33.55 33.60 33.65 33.70 33.75 33.80 33.85 33.90 33.95 260  Haida Gwaii Yakoun River 2010 data Identifier YR 384 YR 385 YR 386 YR 387 YR 388 YR 389 YR 390 YR 391 YR 392 YR 393 YR 394 YR 395 YR 396 YR 397 YR 398 YR 399 YR 400 YR 401 YR 402 YR 403 YR 404 YR 405 YR 406 YR 407 YR 408 YR 409 YR 410 YR 411 YR 412 YR 413 YR 414 YR 415 YR 416 YR 417 YR 418 YR 419  Amt. (Mg) 3.583 3.73 3.685 3.903 3.028 3.524 3.324 3.24 3.217 3.872 3.673 3.543 3.281 3.776 3.705 3.477 3.794 3.465 3.247 3.948 6.019 3.637 3.73 3.137 3.554 3.112 3.237 3.728 3.474 3.406 3.406 3.098 3.746 3.703 3.586 3.493  TOC% 0.58 0.48 0.36 0.51 0.49 0.53 0.71 0.59 0.54 0.63 0.32 0.60 0.65 0.49 0.38 0.35 0.37 0.32 0.32 0.26 0.29 0.38 0.44 0.42 0.41 0.31 0.74 0.49 0.36 0.38 0.50 0.57 0.54 0.47 0.68 0.81  δ 13Cadj -31.69 -31.45 -31.12 -31.31 -31.10 -31.14 -30.06 -31.50 -31.20 -30.34 -30.81 -31.71 -30.99 -30.67 -31.21 -30.88 -30.61 -30.54 -30.12 -30.06 -30.04 -30.50 -30.24 -30.39 -30.39 -30.26 -30.50 -30.66 -30.65 -30.67 -30.46 -30.90 -31.37 -30.51 -24.21 -24.03  Ht (m) 34.00 34.05 34.10 34.15 34.20 34.25 34.30 34.35 34.40 34.45 34.50 34.55 34.60 34.65 34.70 34.75 34.80 34.85 34.90 34.95 35.00 35.05 35.10 35.15 35.20 35.25 35.30 35.35 35.40 35.45 35.50 35.55 35.60 35.65 35.70 35.75 261  Haida Gwaii Yakoun River 2010 data Identifier YR 420 YR 421 YR 422 YR 423 YR 424 YR 425 YR 426 YR 427 YR 428 YR 429 YR 430 YR 431 YR 432 YR 433 YR 434 YR 435 YR 436 YR 437 YR 438 YR 439 YR 440 YR 441 YR 442 YR 443 YR 444 YR 445 YR 446 YR 447 YR 448 YR 449 YR 450 YR 451 YR 452 YR 453 YR 454 YR 455  Amt. (Mg) 3.898 3.521 3.591 3.872 3.792 3.721 3.141 3.885 3.836 3.66 3.452 3.732 3.06 3.021 3.631 3.152 3.059 3.724 3.384 3.567 3.742 3.266 3.685 3.669 3.737 3.031 3.112 3.674 3.278 6.123 6.029 3.528 3.734 5.861 6.333 5.519  TOC% 0.79 0.86 0.98 1.00 0.92 0.89 0.65 0.58 0.66 0.70 0.84 0.74 0.85 0.86 0.81 0.94 0.96 0.90 0.91 0.90 0.88 0.90 0.81 0.81 0.76 0.80 0.85 0.84 0.57 0.79 0.70 0.67 0.96 0.81 0.26 0.21  δ 13Cadj -24.15 -24.16 -23.95 -24.47 -24.40 -24.29 -25.31 -25.51 -25.94 -25.79 -24.67 -24.32 -24.11 -24.16 -25.10 -24.01 -24.21 -24.37 -24.09 -24.55 -24.55 -24.64 -24.68 -24.67 -24.49 -25.34 -25.01 -24.57 -25.25 -25.09 -24.21 -24.72 -25.33 -24.48 -28.48 -27.25  Ht (m) 35.80 35.85 35.90 35.95 36.00 36.05 36.10 36.15 36.20 36.25 36.30 36.35 36.40 36.45 36.50 36.55 36.60 36.65 36.70 36.75 36.80 36.85 36.90 36.95 37.00 37.05 37.10 37.15 37.20 37.25 37.30 37.35 37.40 37.45 37.50 37.55 262  Haida Gwaii Yakoun River 2010 data Identifier YR 456 YR 457 YR 458 YR 459 YR 460 YR 461 YR 462 YR 463 YR 464 YR 465 YR 466 YR 467 YR 468 YR 469 YR 470 YR 471 YR 472 YR 473 YR 474 YR 475 YR 476 YR 477 YR 478 YR 479 YR 480 YR 481 YR 482 YR 483 YR 483 YR 484 YR 485 YR 486 YR 487 YR 488 YR 489 YR 490  Amt. (Mg) 3.333 3.812 3.872 3.926 3.568 4.086 3.282 3.809 5.766 3.468 3.502 3.701 4.423 2.937 3.668 3.588 3.24 3.669 4.034 3.624 3.988 3.889 3.443 3.691 3.526 3.449 3.799 3.39 4.875 4.42 3.408 4.292 3.326 3.787 3.907 3.484  TOC% 0.70 0.67 0.74 0.73 0.57 0.52 0.59 0.53 0.54 0.55 0.49 0.50 0.57 0.96 0.85 0.88 0.81 0.80 0.59 0.75 0.82 0.92 0.93 0.88 0.93 0.95 0.81 0.77 0.49 0.81 0.79 0.77 0.82 0.90 0.85 0.86  δ 13Cadj -23.74 -23.86 -23.98 -25.50 -25.77 -26.49 -25.64 -25.70 -25.53 -25.53 -25.76 -25.05 -25.17 -28.99 -24.94 -25.25 -24.76 -25.49 -25.88 -24.90 -25.24 -25.91 -24.79 -25.10 -25.22 -25.33 -25.30 -24.51 -25.85 -24.85 -24.41 -24.77 -24.28 -24.71 -24.62 -25.02  Ht (m) 37.60 37.65 37.70 37.75 37.80 37.85 37.90 37.95 38.00 38.05 38.10 38.15 38.20 38.25 38.30 38.35 38.40 38.45 38.50 38.55 38.60 38.65 38.70 38.75 38.80 38.85 38.90 38.95 39.00 39.05 39.10 39.15 39.20 39.25 39.30 39.35 263  Haida Gwaii Yakoun River 2010 data Identifier YR 491 YR 492 YR 493 YR 494 YR 495 YR 496 YR 497 YR 498 YR 499 YR 500 YR 501 YR 501 YR 502 YR 502 YR 503 YR 503 YR 504 YR 505 YR 506 YR 507 YR 508 YR 509 YR 510 YR 511 YR 512 YR 513 YR 514 YR 515 YR 516 YR 517 YR 518 YR 519 YR 520 YR 522 YR 523 YR 524  Amt. (Mg) 3.284 3.736 3.55 3.785 3.808 3.355 3.298 3.39 3.274 3.116 3.962 3.804 4.388 4.639 3.638 3.462 4.398 3.26 3.298 3.676 3.14 3.077 3.644 3.433 3.206 3.662 3.492 3.395 3.707 3.168 3.209 3.603 3.836 3.287 3.418 3.448  TOC% 0.80 0.67 0.84 0.79 0.83 0.85 0.69 0.49 0.36 0.55 0.58 0.50 0.48 0.42 0.83 0.72 1.50 0.96 0.83 0.86 0.64 0.75 0.52 0.58 0.67 0.79 0.80 0.64 0.66 0.52 0.56 0.66 0.75 0.71 0.72 0.65  δ 13Cadj -24.66 -25.17 -25.34 -25.32 -25.75 -25.30 -25.43 -26.25 -26.85 -27.23 -24.98 -25.37 -27.84 -27.97 -24.94 -25.29 -25.62 -26.03 -25.47 -25.96 -25.86 -26.09 -26.29 -26.05 -25.76 -25.36 -24.99 -28.16 -28.34 -25.60 -25.50 -25.49 -25.73 -25.59 -25.29 -25.51  Ht (m) 39.40 39.45 39.50 39.55 39.60 39.65 39.70 39.75 39.80 39.85 39.90 39.95 40.00 40.05 40.10 40.15 40.20 40.25 40.30 40.35 40.40 40.45 40.50 40.55 40.60 40.65 40.70 40.75 40.80 40.85 40.90 40.95 41.00 41.05 41.10 41.15 264  Haida Gwaii Yakoun River 2010 data Identifier YR 525 YR 526 YR 527 YR 528 YR 529 YR 530 YR 531 YR 532 YR 533 YR 534 YR 535 YR 536 YR 537 YR 538 YR 539 YR 540 YR 541 YR 542 YR 543 YR 544 YR 545 YR 546 YR 547 YR 548 YR 549 YR 550 YR 551 YR 552 YR 553 YR 554 YR 555 YR 556 YR 557 YR 558 YR 559 YR 560  Amt. (Mg) 3.811 3.298 3.514 3.649 3.811 3.587 3.799 3.604 3.496 3.736 3.326 3.449 3.354 3.779 3.987 3.098 3.524 3.382 3.811 3.624 3.796 3.625 3.77 3.241 3.702 5.303 3.533 3.372 3.673 3.976 3.03 5.285 3.689 3.679 3.578 3.765  TOC% 0.69 0.75 0.56 0.71 0.65 1.10 0.68 0.70 0.68 0.68 0.63 0.70 0.86 0.69 0.63 0.62 0.64 0.59 0.68 0.61 0.68 0.64 0.68 0.77 0.56 0.63 0.53 0.61 0.63 0.58 0.64 0.28 0.61 0.51 0.51 0.45  δ 13Cadj -25.38 -25.30 -27.28 -25.77 -25.72 -26.99 -25.46 -25.60 -25.58 -25.26 -25.39 -25.25 -25.07 -25.26 -24.89 -24.74 -25.22 -24.86 -25.65 -25.89 -26.26 -25.71 -26.15 -25.69 -25.55 -25.82 -25.93 -25.59 -25.17 -25.64 -25.65 -25.86 -25.95 -26.03 -26.44 -26.14  Ht (m) 41.20 41.25 41.30 41.35 41.40 41.45 41.50 41.55 41.60 41.65 41.70 41.75 41.80 41.85 41.90 41.95 42.00 42.05 42.10 42.15 42.20 42.25 42.30 42.35 42.40 42.45 42.50 42.55 42.60 42.65 42.70 42.75 42.80 42.85 42.90 42.95 265  Haida Gwaii Yakoun River 2010 data Identifier YR 561 YR 562 YR 563 YR 564 YR 565 YR 566 YR 567 YR 568 YR 569 YR 570 YR 571 YR 572 YR 573 YR 574 YR 575 YR 576 YR 577 YR 578 YR 579 YR 580 YR 581 YR 582  Amt. (Mg) 3.42 3.328 3.246 3.931 3.536 4.272 3.342 3.623 4.042 3.959 4.077 3.196 3.835 3.38 3.298 3.73 3.996 3.168 3.386 3.639 3.526 6.273  TOC% 0.47 0.62 0.62 0.60 0.81 0.72 0.62 0.56 0.71 0.60 0.76 0.71 0.66 0.66 0.68 0.70 0.62 0.59 0.58 0.61 0.79 0.30  δ 13Cadj -25.86 -25.26 -25.25 -24.38 -24.00 -23.98 -24.14 -24.48 -24.34 -24.72 -24.92 -24.42 -24.55 -24.45 -24.89 -24.95 -25.09 -25.04 -25.13 -24.89 -24.52 -23.55  YR 584 YR 585 YR 586 YR 587 YR 588 YR 589 YR 590 YR 591 YR 592  3.808 3.453 3.72 4.184 3.061 3.593 3.099 3.735 6.65  0.55 0.83 0.62 0.65 0.59 0.57 0.64 0.59 0.62  -25.74 -25.55 -25.65 -25.20 -25.58 -25.16 -24.75 -25.48 -24.74  Ht (m) 43.00 43.05 43.10 43.15 43.20 43.25 43.30 43.35 43.40 43.45 43.50 43.55 43.60 43.65 43.70 43.75 43.80 43.85 43.90 43.95 44.00 44.05 44.15 44.25 44.35 44.45 44.55 44.65 44.75 44.85 44.95 45.05  266  Haida Gwaii Whiteaves Bay data Identifier WB 079 WB 080 WB 081 WB 082 WB 083 WB 084 WB 085 WB 086 WB 087 WB 088 WB 089 WB 090 WB 091 WB 092 WB 093 WB 094 WB 095 WB 096 WB 097 WB 098 WB 099 WB 100 WB 101 WB 102 WB 103 WB 104 WB 105 WB 106 WB 107 WB 108 WB 109 WB 110 WB 111 WB 112 WB 113 WB 114  Amt. (Mg) 0.90 2.71 6.51 5.64 1.01 1.45 3.25 3.29 1.02 1.36 1.46 1.94 0.88 0.80 1.05 0.82 1.08 1.80 1.31 0.96 0.83 2.66 2.53 2.74 2.61 2.17 0.69 0.65 1.88 0.77 0.82 0.73 0.73 0.68 0.82 0.63  TOC% 2.15 0.74 0.40 0.33 1.33 1.42 0.74 0.66 1.47 1.12 2.03 1.77 2.90 2.46 1.83 2.06 0.81 0.72 1.32 1.63 1.31 0.85 0.66 0.91 1.02 1.24 1.62 2.22 1.42 2.51 2.15 2.37 2.52 2.17 3.00 2.84  δ 13Cadj -29.36 -29.15 -28.72 -28.31 -29.08 -29.40 -28.39 -29.23 -29.39 -29.41 -29.45 -28.59 -28.65 -28.82 -29.27 -28.84 -28.56 -27.70 -28.36 -28.14 -28.57 -28.75 -28.50 -26.68 -28.84 -28.33 -28.37 -28.36 -28.27 -28.52 -27.63 -27.99 -28.06 -28.10 -27.56 -27.98  TN% 0.19  δ 15N 0.00  0.08  0.67  0.16  0.70  0.16  -0.25  0.14  -0.27  0.14  0.97  Ht (m) 0.00 0.30 0.60 0.90 1.20 1.50 1.80 2.10 2.40 2.70 3.00 3.30 3.60 3.90 4.20 4.50 4.80 5.10 6.30 6.60 6.90 7.20 7.50 7.80 8.10 8.40 8.70 9.00 9.30 9.60 9.90 10.20 10.50 10.80 11.10 11.40 267  Haida Gwaii Whiteaves Bay data Identifier WB 115 WB 116 WB 117 WB 118 WB 119 WB 120 WB 121 WB 122 WB 123 WB 124 WB 125 WB 126 WB 127 WB 128 WB 129 WB 130 WB 131 WB 132 WB 133 WB 134 WB 135 WB 136 WB 137 WB 138 WB 139 WB 140 WB 141 WB 142 WB 143 WB 144 WB 145 WB 146 WB 147 WB 148 WB 149 WB 150  Amt. (Mg) 0.71 0.62 0.90 0.91 0.74 0.78 0.62 0.79 0.66  TOC% 2.10 2.28 1.61 1.68 1.80 2.03 1.87 2.45 2.38  δ 13Cadj -27.83 -28.12 -28.18 -28.22 -27.86 -28.69 -28.00 -28.10 -27.64  0.69 0.61 0.75 0.77 0.93 0.82 0.62 3.10 0.84 3.00 0.66 0.90 0.66 3.46 0.69 0.92 1.07 1.29 6.01 1.86 0.78 0.66 1.10 0.93 0.86 0.90  3.36 1.79 1.90 1.87 2.34 1.56 1.39 1.14 1.97 0.70 1.57 1.71 1.62 0.38 1.79 1.70 1.20 1.13 0.18 0.65 1.64 3.33 1.37 2.10 1.59 2.08  -28.18 -27.79 -27.67 -27.50 -27.84 -26.57 -27.54 -25.22 -27.20 -26.51 -27.22 -26.82 -25.76 -25.66 -27.72 -27.06 -27.36 -26.74 -27.60 -26.16 -26.96 -26.85 -26.66 -26.46 -26.51 -26.48  TN% 0.15  δ 15N 0.09  0.20  -0.18  0.17  0.21  0.07  0.88  0.12  0.12  0.16  -0.18  Ht (m) 11.70 12.00 12.30 12.60 12.90 13.20 13.50 13.80 14.10 14.40 14.70 15.00 15.30 15.60 15.90 16.20 16.50 16.80 17.10 17.40 17.70 18.00 18.30 18.60 18.90 19.30 19.60 19.90 20.20 20.50 20.80 21.10 21.40 21.70 22.00 22.30 268  Haida Gwaii Whiteaves Bay data Identifier WB 151 WB 152 WB 153 WB 154 WB 155 WB 156 WB 157 WB 158 WB 159 WB 160 WB 161 WB 162 WB 163 WB 164 WB 165 WB 166 WB 167 WB 168 WB 169 WB 170 WB 171 WB 172 WB 173 WB 174 WB 175 WB 176 WB 177 WB 178 WB 179 WB 180 WB 181 WB 182 WB 183 WB 184 WB 185 WB 186  Amt. (Mg) 0.84 0.65 3.71 3.69 3.44 3.28 3.41 0.88 3.63 3.38 2.90 3.62 3.73 3.38 0.71 1.36 3.32 4.15 6.39 0.90 0.73 0.78 1.08 1.06 0.62 3.07 0.75 3.50 3.33 3.71 3.61 3.67 3.57 4.50 3.28 3.14  TOC% 1.32 2.03 0.34 0.51 0.47 0.71 0.74 1.71 0.51 0.45 0.55 0.48 0.36 0.59 1.38 1.13 0.43 0.52 0.16 0.85 1.56 1.70 2.05 1.83 1.99 0.88 0.92 0.40 0.30 0.49 0.30 0.33 0.44 0.45 0.30 0.46  δ 13Cadj -26.57 -27.08 -27.04 -28.16 -27.98 -27.62 -28.07 -27.47 -25.81 -27.50 -27.68 -27.31 -27.56 -28.90 -28.40 -28.39 -28.59 -27.83 -28.45 -28.09 -28.35 -27.59 -26.60 -27.85 -28.28 -23.56 -28.03 -26.74 -27.65 -27.34 -24.82 -28.12 -24.33 -22.39 -28.08 -24.76  TN%  δ 15N  Ht (m) 22.60 23.13 23.43 23.73 24.03 24.33 24.63 24.93 25.23 25.53 25.83 26.13 26.43 26.83 27.13 27.43 27.73 27.83 28.13 28.43 28.73 29.03 29.33 29.63 29.93 30.23 30.53 30.83 31.13 31.43 31.73 33.03 33.33 33.63 33.93 34.23 269  Haida Gwaii Whiteaves Bay data Identifier WB 187 WB 188 WB 189 WB 190 WB 191 WB 192 WB 193 WB 194 WB 195 WB 196 WB 197 WB 198 WB 199 WB 200 WB 201 WB 202 WB 203 WB 204 WB 205 WB 206 WB 207 WB 208 WB 209 WB 210 WB 211 WB 212 WB 213 WB 214 WB 215 WB 216 WB 217 WB 218 WB 219 WB 220 WB 221 WB 222  Amt. (Mg) 3.05 2.98 3.69 3.47 3.46 3.67 3.13 3.96 3.20 3.53 3.70 2.95 3.53 3.42 3.82 3.73 3.28 3.53 6.03 3.54 3.39 4.08 4.15 4.04 6.13 3.32 3.63 3.19 5.25 3.20 3.27 3.44 5.94 3.03 6.55 3.98  TOC% 1.00 0.35 0.64 0.62 0.80 0.62 0.61 0.51 0.68 0.84 0.56 0.83 0.27 0.49 0.27 0.46 0.45 0.40 0.34 0.51 0.58 0.44 0.52 0.54 0.34 0.79 0.82 0.61 0.94 0.37 0.58 0.29 0.25 0.46 0.33 0.18  δ 13Cadj -20.33 -24.91 -21.70 -25.84 -24.07 -26.66 -27.50 -26.92 -26.78 -27.61 -28.24 -23.20 -26.00 -26.29 -25.73 -25.93 -26.61 -25.21 -25.86 -21.32 -21.56 -27.35 -27.70 -27.00 -26.23 -18.91 -22.03 -28.60 -25.96 -28.03 -26.76 -27.77 -27.41 -27.26 -23.18 -25.44  TN%  δ 15N  Ht (m) 34.53 34.83 35.13 35.43 35.73 36.03 36.33 36.63 36.93 37.23 37.53 37.83 38.13 38.43 38.73 39.03 39.33 39.63 39.93 40.23 40.53 40.83 41.13 41.43 41.73 42.03 42.33 42.63 42.93 43.25 43.55 43.85 44.15 44.45 44.75 45.05 270  Haida Gwaii Whiteaves Bay data Identifier WB 223 WB 224 WB 225 WB 226 WB 227 WB 228 WB 229 WB 230 WB 231 WB 232 WB 233 WB 234 WB 235 WB 236 WB 237 WB 238 WB 239 WB 240 WB 241 WB 242 WB 243 WB 244 WB 245 WB 246 WB 247 WB 248 WB 249 WB 250 WB 251 WB 252 WB 253 WB 254 WB 255 WB 256 WB 257 WB 258  Amt. (Mg) 3.23 3.00 3.65 6.51 3.71 3.60 6.70 3.59 5.84 5.96 6.18 3.65 3.47 6.32 3.12 3.43 3.41 3.15 3.84 3.45 4.07 5.97 6.36 6.72 3.19 6.62 3.19 3.07 6.60 4.23 3.80 1.69 2.30 3.25 3.24 3.42  TOC% 0.33 0.54 0.61 0.23 0.45 0.36 0.76 0.62 0.25 0.84 0.37 0.39 0.29 0.08 0.64 0.74 0.57 0.56 0.65 0.78 0.55 0.34 0.36 0.18 0.50 0.14 0.29 0.31 0.28 0.43 0.50 0.86 1.29 0.74 0.73 0.66  δ 13Cadj -27.52 -27.45 -27.28 -27.86 -28.24 -25.46 -18.29 -25.54 -26.70 -19.86 -27.72 -24.18 -26.88 -27.25 -27.69 -26.50 -26.17 -28.08 -28.24 -25.84 -28.03 -22.14 -22.72 -27.57 -22.61 -26.71 -26.42 -27.86 -28.94 -21.73 -27.35 -28.43 -27.92 -28.62 -27.22 -28.30  TN%  δ 15N  Ht (m) 45.35 45.65 45.95 46.25 46.55 46.85 47.15 47.45 47.75 48.05 48.35 48.65 48.95 49.25 49.55 49.85 50.15 50.45 50.75 51.05 51.35 51.65 51.95 52.25 52.55 52.85 53.15 53.45 53.75 54.05 54.35 54.65 54.95 55.25 55.55 55.85 271  Haida Gwaii Whiteaves Bay data Identifier WB 259 WB 260 WB 261 WB 262 WB 263 WB 264 WB 265 WB 266 WB 267 WB 268 WB 269 WB 270 WB 271 WB 272 WB 273 WB 274 WB 275 WB 276 WB 277 WB 278 WB 279 WB 280 WB 281 WB 282 WB 283 WB 284 WB 285 WB 286 WB 287 WB 288 WB 289 WB 290 WB 291 WB 292 WB 293 WB 294  Amt. (Mg) 3.86 3.16 3.41 5.92 3.66 4.15 3.97 3.54 3.73 3.59 3.38 6.70 3.66 6.51 3.85 4.80 6.34 3.64 3.73 3.11 3.61 3.25 6.34 3.46 3.82 3.32 4.21 3.52 3.28 3.73 3.60 6.73 3.08 5.09 3.03 3.19  TOC% 0.42 0.60 0.86 0.49 0.37 0.27 0.74 0.57 0.70 0.54 0.41 0.23 0.41 0.27 0.53 0.53 0.16 0.46 0.37 0.70 0.38 0.79 0.35 0.35 0.61 0.14 0.66 0.31 0.90 0.82 0.76 0.28 0.84 0.29 0.81 0.70  δ 13Cadj -24.72 -28.59 -25.31 -27.54 -26.57 -25.94 -26.29 -28.02 -27.34 -28.49 -27.47 -26.59 -27.05 -26.87 -28.16 -28.00 -27.95 -27.35 -28.66 -28.62 -27.01 -28.48 -26.64 -27.43 -28.20 -27.55 -28.56 -28.46 -28.91 -28.50 -28.50 -24.07 -25.22 -25.98 -28.51 -27.83  TN%  δ 15N  Ht (m) 56.15 56.45 56.75 57.05 57.35 57.65 57.95 58.25 58.55 58.85 59.15 59.45 59.75 60.05 60.35 60.65 60.95 61.25 61.55 61.85 62.15 62.45 62.75 63.05 63.35 63.65 63.95 64.25 64.55 64.85 65.15 65.45 65.75 66.05 66.35 66.65 272  Haida Gwaii Whiteaves Bay data Identifier WB 295 WB 02 WB 296 WB 03 WB 04 WB 297 WB 05 WB 06 WB 07 WB 298 WB 08 WB 299 WB 09 WB 10 WB 300 WB 11 WB 301 WB 12 WB 13 WB 302 WB 14 WB 303 WB 15 WB 16 WB 17 WB 18 WB 19 WB 20 WB 21 WB 22 WB 23 WB 24 WB 25 WB 26 WB 27 WB 28  Amt. (Mg) 3.53 4.68 3.15 4.61 4.41 4.31 4.79 4.50 4.27 3.08 4.59 3.50 4.53 4.54 2.48 4.30 2.95 4.38 4.71 2.83 4.54 3.28 4.72 4.59 4.66 4.43 4.36 4.46 4.73 4.32 4.62 4.51 4.61 4.73 4.68 4.50  TOC% 0.50 0.53 0.15 0.80 0.62 0.18 0.70 0.86 0.87 0.14 0.68 0.23 0.71 0.58 0.39 0.67 0.37 0.46 0.34 0.74 0.39 0.29 0.40 0.46 0.42 0.67 0.55 0.81 0.66 0.59 0.43 0.64 0.45 0.81 0.45 0.72  δ 13Cadj -24.96 -28.14 -28.72 -26.61 -27.58 -27.89 -28.46 -27.62 -23.52 -27.51 -28.34 -28.56 -27.24 -28.15 -28.21 -25.40 -27.65 -28.48 -28.73 -28.93 -27.76 -28.19 -27.05 -28.77 -27.76 -28.96 -28.41 -26.37 -27.53 -28.60 -28.86 -28.84 -28.09 -26.48 -27.87 -26.79  TN%  δ 15N  Ht (m) 66.95 67.05 67.25 67.35 67.50 67.55 67.60 67.80 67.84 67.85 68.04 68.15 68.24 68.44 68.45 68.64 68.75 68.84 69.04 69.05 69.24 69.35 69.54 69.74 69.86 70.06 70.33 70.61 70.81 71.01 71.21 71.41 71.61 71.81 72.01 72.21 273  Haida Gwaii Whiteaves Bay data Identifier WB 29 WB 30 WB 31 WB 32 WB 33 WB 34 WB 35 WB 36 WB 37 WB 38 WB 39 WB 40 WB 41 WB 42 WB 43 WB 44 WB 45 WB 46 WB 47 WB 48 WB 49 WB 50 WB 51 WB 52 WB 53 WB 54 WB 55 WB 56 WB 57 WB 304 WB 58 WB 305 WB 59 WB 306 WB 60 WB 307  Amt. (Mg) 4.77 4.65 4.54 4.63 4.72 4.77 4.66 4.69 4.51 4.70 4.39 4.48 4.45 4.04 4.43 4.33 4.56 4.61 4.44 4.80 4.58 4.61 4.74 4.44 4.57 4.61 4.41 4.69 4.79 3.03 4.63 3.88 4.66 3.71 4.46 3.72  TOC% 0.68 0.45 0.39 0.56 0.43 0.47 0.45 0.40 0.54 0.59 0.51 0.70 0.52 0.57 0.52 0.47 0.50 0.69 0.36 0.52 0.61 0.50 0.46 0.65 0.42 0.46 0.61 0.51 0.64 0.53 0.44 0.53 0.47 0.41 0.51 0.51  δ 13Cadj -28.51 -27.32 -26.54 -28.12 -27.07 -27.07 -27.39 -27.13 -28.28 -26.76 -27.68 -27.17 -27.41 -25.38 -29.16 -27.87 -27.51 -27.07 -26.14 -27.75 -28.08 -27.44 -27.39 -27.08 -27.40 -26.97 -27.30 -26.20 -26.31 -26.18 -26.53 -26.16 -26.60 -26.15 -27.09 -26.47  TN%  δ 15N  Ht (m) 72.41 72.61 72.91 73.11 73.31 73.51 73.61 73.71 73.93 74.03 74.16 74.26 74.36 74.46 74.56 74.66 74.76 74.86 74.96 75.06 75.16 75.26 75.36 75.46 75.56 75.66 75.76 75.96 76.16 76.23 76.36 76.53 76.56 76.73 76.76 76.88 274  Haida Gwaii Whiteaves Bay data Identifier WB 61 WB 308 WB 62 WB 309 WB 310 WB 63 WB 311 WB 312 WB 64 WB 313 WB 314 WB 65 WB 315 WB 316 WB 66 WB 317 WB 318 WB 67 WB 319 WB 68 WB 320 WB 69 WB 321 WB 70 WB 322 WB 71 WB 72 WB 323 WB 73 WB 324 WB 74 WB 75 WB 325 WB 76 WB 326 WB 77  Amt. (Mg) 4.32 4.94 3.54 3.05 4.75 3.10 3.74 4.89 3.05 3.54 4.76 3.15 3.63 4.55 3.42 3.07 4.74 3.24 4.41 3.09 4.58 3.07 4.90 3.34 4.34 4.36 3.51 4.45 2.95 4.85 4.70 3.08 4.51 3.11 4.68  TOC% 0.57 0.42 0.22 0.40 0.53 0.36 0.55 0.61 0.52 0.61 0.61 0.47 0.61 0.57 0.55 1.50 0.47 0.45 0.62 0.39 0.65 0.36 0.68 0.43 0.54 0.50 0.48 0.51 0.60 0.75 0.54 0.60 0.73 0.59 0.70 0.56  δ 13Cadj -27.09 -26.95 -26.82 -28.43 -28.38 -27.02 -28.35 -28.14 -27.24 -27.94 -28.17 -27.25 -28.49 -28.45 -26.76 -28.43 -27.97 -27.00 -28.49 -29.02 -28.23 -28.59 -28.05 -28.52 -28.21 -28.48 -28.59 -28.19 -28.51 -28.28 -28.40 -28.52 -28.10 -28.70 -28.33 -28.56  TN%  δ 15N  Ht (m) 76.96 77.03 77.16 77.18 77.28 77.36 77.38 77.48 77.56 77.58 77.68 77.76 77.78 77.88 77.96 77.98 78.08 78.16 78.28 78.36 78.48 78.56 78.68 78.76 78.88 78.96 79.16 79.18 79.36 79.48 79.56 79.76 79.78 79.96 80.08 80.16 275  Haida Gwaii Whiteaves Bay data Identifier WB 78 WB 327 WB 328 WB 329 WB 330 WB 331 WB 332 WB 333 WB 334 WB 335 WB 336 WB 337 WB 338 WB 339 WB 340 WB 341 WB 342 WB 343 WB 344 WB 345 WB 346 WB 347 WB 348 WB 349 WB 350 WB 351 WB 352 WB 353 WB 354 WB 355 WB 356 WB 357 WB 358 WB 359 WB 360 WB 361  Amt. (Mg) 4.35 2.89 3.21 3.57 3.27 3.42 3.22 5.17 3.84 3.62 3.26 3.48 3.07 3.01 3.56 3.49 3.65 3.27 3.51 3.58 3.64 3.49 3.80 3.20 6.23 3.99 5.15 3.21 3.44 3.86 4.21 3.29 3.27 5.06 3.39 3.74  TOC% 0.62 0.80 0.67 0.63 0.82 0.71 0.65 0.63 0.68 0.65 0.79 0.56 0.50 0.33 0.42 0.40 0.34 0.34 0.33 0.41 0.22 0.36 0.50 0.55 0.47 0.44 0.44 0.73 0.39 0.42 0.35 0.34 0.35 0.21 0.30 0.35  δ 13Cadj -28.67 -28.38 -27.60 -28.02 -27.90 -28.31 -28.36 -28.09 -28.36 -28.12 -28.69 -27.85 -25.55 -25.46 -25.32 -25.13 -25.52 -25.87 -25.44 -24.98 -25.91 -25.66 -24.99 -24.71 -24.72 -24.99 -25.16 -25.90 -24.98 -25.66 -25.02 -25.65 -25.80 -25.86 -25.66 -26.11  TN%  δ 15N  Ht (m) 80.36 80.38 80.68 80.98 81.28 81.58 81.88 82.18 82.48 82.78 83.08 83.38 83.78 84.18 84.48 84.78 85.18 85.58 85.88 86.28 86.68 87.08 87.48 87.88 88.28 88.68 89.08 89.48 89.88 90.28 90.68 91.08 91.48 91.88 92.28 92.68 276  Haida Gwaii Whiteaves Bay data Identifier WB 362 WB 363 WB 364 WB 365 WB 366 WB 367 WB 368 WB 369 WB 370 WB 371 WB 372 WB 373 WB 374 WB 375  Amt. (Mg) 3.67 7.18 3.77 3.54 3.67 3.61 3.29 3.37 3.56 3.42 3.12 5.74 5.98 3.73  TOC% 0.25 0.18 0.32 0.31 0.36 0.35 0.37 0.30 0.32 0.34 0.34 0.23 0.25 0.35  δ 13Cadj -26.08 -26.89 -26.04 -25.95 -26.19 -25.15 -25.58 -25.42 -25.15 -25.06 -25.44 -26.03 -26.06 -25.40  TN%  δ 15N  Ht (m) 93.08 93.48 93.88 94.28 94.68 95.08 95.48 95.88 96.28 99.67 100.07 101.97 102.37 102.77  277  Northern Alaska South Barrow #3 Core data core ID SB 001 SB 002 SB 003 SB 004 SB 005 SB 006 SB 007 SB 008 SB 009 SB 010 SB 011 SB 012 SB 013 SB 014 SB 015 SB 016 SB 017 SB 018 SB 019 SB 020 SB 021 SB 022 SB 023 SB 024 SB 025 SB 026 SB 027 SB 028 SB 029 SB 030 SB 031 SB 032 SB 033 SB 034 SB 035 SB 036 SB 037 SB 038  Amt. (Mg) 4.35 4.33 4.25 4.29 4.14 4.24 4.17 4.28 4.19 4.08 4.15 4.21 4.23 4.35 4.32 4.31 4.23 4.01 4.26 4.24 4.01 4.37 4.03 4.04 4.40 4.33 4.11 4.15 4.08 4.34 4.17 4.35 4.14 4.30 4.34 4.01 4.19 4.27  TOC% 1.17 1.02 0.92 1.11 0.95 1.37 0.83 1.21 0.90 1.19 1.12 1.07 0.80 1.03 1.04 0.90 0.80 0.99 0.95 0.87 1.13 0.84 0.86 0.90 1.41 1.22 1.45 2.49 1.77 1.94 0.94 1.07 1.63 1.59 1.06 1.29 0.97 1.93  δ 13Cadj -27.08 -27.35 -27.33 -27.23 -25.92 -27.64 -27.14 -27.68 -27.33 -27.70 -27.60 -27.58 -27.52 -27.62 -27.92 -27.36 -27.50 -27.59 -27.83 -27.38 -27.58 -27.04 -27.11 -27.39 -27.36 -27.49 -27.70 -27.87 -27.67 -28.11 -27.71 -27.72 -28.06 -28.19 -27.79 -27.25 -27.52 -27.75  depth (m) 676.04 675.84 675.64 675.44 675.24 675.04 674.84 674.64 674.44 674.24 674.04 673.84 673.64 673.44 673.24 673.04 672.84 672.64 672.44 672.24 672.04 671.84 671.64 671.44 671.24 671.04 670.84 670.64 670.44 670.24 670.04 669.84 669.64 669.44 669.24 669.04 668.84 668.64 278  Northern Alaska South Barrow #3 Core data core ID SB 039 SB 040 SB 041 SB 042 SB 043 SB 044 SB 045 SB 046 SB 047 SB 048 SB 049 SB 050 SB 051 SB 052 SB 053 SB 054 SB 055 SB 056 SB 057 SB 058 SB 059 SB 060 SB 061 SB 062 SB 063 SB 064 SB 065 SB 066 SB 067 SB 068 SB 069 SB 070 SB 071 SB 072 SB 073 SB 074 SB 075 SB 076  Amt. (Mg) 4.24 4.23 4.16 4.30 4.15 4.38 4.01 4.18 4.17 4.20 4.30 4.13 4.11 4.13 4.40 4.19 4.07 4.08 4.18 4.23 4.27 4.24 4.32 4.23 4.04 4.20 4.21 4.38 4.22 4.00 4.29 4.02 4.05 4.31 4.21 4.07 4.01 4.32  TOC% 1.78 1.27 1.88 1.26 1.01 1.70 2.84 1.46 1.84 1.21 2.27 2.63 0.89 1.01 1.59 1.13 1.47 1.57 1.31 1.57 1.90 1.01 0.98 0.91 0.85 0.90 0.90 0.89 0.81 0.88 0.98 0.94 0.81 1.11 0.82 0.82 1.06 1.46  δ 13Cadj -27.69 -27.99 -28.24 -28.00 -27.97 -28.19 -27.90 -28.22 -28.47 -28.01 -27.86 -28.28 -27.88 -27.62 -27.94 -27.56 -27.99 -27.94 -28.10 -27.93 -28.33 -27.65 -27.50 -27.57 -27.23 -27.36 -26.98 -27.01 -27.26 -26.62 -26.73 -26.51 -26.66 -26.96 -27.08 -26.82 -27.16 -27.41  depth (m) 668.44 668.24 668.04 667.84 667.64 667.44 667.24 667.04 666.84 666.64 666.44 666.24 666.04 665.84 665.64 665.44 665.24 665.04 664.84 664.64 664.44 664.24 664.04 663.84 663.64 663.44 663.24 663.04 662.84 662.64 662.44 662.24 662.04 661.84 661.64 661.44 661.24 661.04 279  Northern Alaska South Barrow #3 Core data core ID SB 077 SB 078 SB 079 SB 080 SB 081 SB 082 SB 083 SB 084 SB 085 SB 086 SB 087 SB 088 SB 089 SB 090 SB 091 SB 092 SB 093 SB 094 SB 095 SB 096 SB 097 SB 098 SB 099 SB 100 SB 101 SB 102 SB 103 SB 104 SB 105 SB 106 SB 107 SB 108 SB 109 SB 110 SB 111 SB 112 SB 113 SB 114  Amt. (Mg) 4.21 4.06 4.21 4.20 4.36 4.31 4.01 4.12 4.20 4.24 4.08 4.18 4.12 4.25 4.18 4.29 4.29 4.23 4.14 4.05 4.23 4.37 4.35 4.06 4.16 4.33 4.33 4.21 4.26 4.23 4.25 4.21 4.17 4.23 4.17 4.16 4.16 4.09  TOC% 1.02 1.30 0.93 1.49 0.79 0.93 0.81 0.74 0.76 0.73 0.65 0.59 0.73 0.79 0.70 2.40 0.70 0.78 0.79 0.57 1.02 1.11 0.96 0.93 0.78 0.77 0.94 0.82 0.98 1.01 1.04 1.01 0.74 0.84 0.93 0.87 0.86 0.69  δ 13Cadj -26.89 -27.22 -26.98 -26.73 -26.91 -27.30 -26.98 -26.82 -26.72 -26.27 -26.38 -26.58 -26.81 -26.62 -26.85 -24.91 -26.45 -26.87 -26.96 -26.66 -26.41 -26.57 -26.44 -25.67 -25.44 -25.80 -25.85 -26.20 -25.96 -26.29 -26.06 -26.00 -26.21 -26.20 -26.10 -26.24 -25.99 -27.21  depth (m) 660.84 660.64 660.44 660.24 660.04 659.84 659.64 659.44 659.24 659.04 658.84 658.64 658.44 658.24 658.04 657.84 657.64 657.44 657.24 657.04 656.84 656.64 656.44 656.24 656.04 655.84 655.64 655.44 655.24 655.04 654.84 654.64 654.44 654.24 654.04 653.84 653.64 653.44 280  Northern Alaska South Barrow #3 Core data core ID SB 115 SB 116 SB 117 SB 118 SB 119  Amt. (Mg) 4.16 4.29 4.25 4.22 4.11  SB 120 SB 121 SB 122 SB 123 SB 124 SB 125 SB 126 SB 127 SB 128 SB 129 SB 130 SB 131 SB 132 SB 133 SB 134 SB 135 SB 136 SB 137 SB 138 SB 139 SB 140 SB 141 SB 142 SB 143 SB 144 SB 145 SB 146 SB 147 SB 148 SB 149 SB 150 SB 151  4.40 4.38 4.03 4.16 4.09 4.01 4.13 4.16 4.29 4.03 4.05 4.15 4.15 4.03 4.35 4.19 4.36 4.23 4.33 4.29 4.02 4.14 4.04 4.27 4.14 4.26 4.15 4.18 4.04 4.36 4.13 4.12  TOC% 0.88 0.86 0.88 0.99 0.90 GAP 1 0.84 0.79 0.79 0.70 0.82 0.77 0.81 0.74 0.83 0.86 0.81 0.87 0.98 0.91 0.91 0.91 0.90 0.99 0.98 0.89 1.47 0.87 0.81 0.81 0.79 0.86 0.79 0.70 0.64 0.72 0.70 0.68  δ 13Cadj -26.12 -25.61 -25.65 -25.12 -25.28  depth (m) 653.24 653.04 652.84 652.64 652.44  -26.48 -26.78 -26.59 -26.51 -27.01 -26.59 -26.83 -26.96 -26.55 -26.35 -26.11 -26.54 -25.77 -25.97 -25.54 -25.62 -25.82 -25.87 -26.10 -26.42 -26.97 -26.19 -26.69 -26.72 -26.51 -26.69 -26.74 -26.75 -26.63 -26.92 -27.14 -26.30  645.87 645.72 645.57 645.42 645.27 645.12 644.97 644.82 644.67 644.52 644.37 644.22 644.07 643.92 643.77 643.62 643.47 643.32 643.17 643.02 642.87 642.72 642.57 642.42 642.27 642.12 641.97 641.82 641.67 641.52 641.37 641.22 281  Northern Alaska South Barrow #3 Core data core ID SB 152 SB 153 SB 154 SB 155 SB 156 SB 157 SB 158 SB 159 SB 160 SB 161 SB 162 SB 163 SB 164 SB 165 SB 166 SB 167 SB 168 SB 169 SB 170 SB 171 SB 172 SB 173 SB 174 SB 175 SB 176 SB 177 SB 178 SB 179 SB 180 SB 181 SB 182 SB 183 SB 184 SB 185 SB 186 SB 187 SB 188 SB 189  Amt. (Mg) 4.17 4.34 4.23 4.12 4.12 4.15 4.30 4.16 4.07 4.33 4.19 4.34 4.20 4.23 4.40 4.06 4.17 4.11 4.22 4.20 4.21 4.31 4.33 4.34 4.31 4.33 4.31 4.01 4.08 4.32 4.11 4.05 4.23 4.14 4.27 4.05 4.19 4.15  TOC% 0.66 0.69 0.74 0.77 0.88 0.76 0.78 0.72 0.83 0.70 0.75 0.74 0.88 0.87 0.66 0.65 0.67 0.74 0.65 0.60 0.70 0.69 0.78 0.75 0.70 0.86 0.75 0.72 0.61 0.65 0.69 0.73 0.66 0.65 0.56 0.61 0.59 0.68  δ 13Cadj -27.24 -27.16 -27.45 -27.28 -27.26 -27.23 -27.32 -27.49 -25.85 -27.27 -27.54 -27.56 -26.80 -27.38 -27.13 -26.98 -27.04 -27.68 -27.50 -27.19 -27.89 -27.69 -27.71 -27.95 -27.72 -27.94 -27.83 -27.93 -27.37 -27.37 -27.39 -27.46 -27.41 -27.56 -27.51 -27.67 -27.35 -27.94  depth (m) 641.07 640.92 640.77 640.62 640.47 640.32 640.17 640.02 639.87 639.72 639.57 639.42 639.27 639.12 638.97 638.82 636.72 636.57 636.42 636.27 636.12 635.97 635.82 635.67 635.52 635.37 635.22 635.07 634.92 634.77 634.62 634.47 634.32 634.17 634.02 633.87 633.72 633.57 282  Northern Alaska South Barrow #3 Core data core ID SB 190 SB 191 SB 192 SB 193 SB 194 SB 195 SB 196 SB 197 SB 198 SB 199 SB 200 SB 201 SB 202 SB 203 SB 204 SB 205 SB 206 SB 207 SB 208 SB 209 SB 210 SB 211 SB 212 SB 213 SB 214 SB 215 SB 216 SB 217 SB 218 SB 219 SB 220 SB 221 SB 222 SB 223 SB 224 SB 225 SB 226 SB 227  Amt. (Mg) 4.06 4.29 4.08 4.26 4.36 4.28 4.18 4.12 4.17 4.27 4.22 4.16 4.16 4.18 4.33 4.19 4.17 4.19 4.22 4.30 4.23 4.24 4.03 4.34 4.34 4.28 4.22 4.08 4.01 4.27 4.18 4.33 4.32 4.21 4.22 4.31 4.19 4.20  TOC% 0.74 0.55 0.62 0.59 0.69 0.66 0.64 0.65 0.55 0.55 0.58 0.59 0.57 0.52 0.52 0.48 0.55 0.55 0.63 0.58 0.55 0.55 0.51 0.68 1.26 0.48 0.61 0.87 0.97 0.61 0.61 0.59 0.54 0.52 0.72 0.71 0.71 0.72  δ 13Cadj -27.99 -27.82 -27.98 -27.69 -27.92 -27.80 -27.91 -27.84 -27.52 -27.73 -27.67 -27.75 -27.79 -27.55 -27.83 -27.21 -27.55 -27.45 -27.59 -27.58 -27.29 -27.16 -27.56 -28.53 -27.93 -27.60 -27.95 -28.87 -28.84 -28.13 -27.64 -28.10 -27.77 -27.83 -27.92 -28.18 -28.17 -27.82  depth (m) 633.42 633.27 633.12 632.97 632.82 632.67 632.52 632.37 632.22 632.07 631.92 631.77 631.67 631.57 631.47 631.37 631.27 631.17 631.07 630.97 630.87 630.77 630.67 630.57 630.47 630.37 630.27 630.17 630.07 629.97 629.87 629.77 629.67 629.57 629.47 629.37 627.27 627.17 283  Northern Alaska South Barrow #3 Core data core ID SB 228 SB 229 SB 230 SB 231 SB 232 SB 233 SB 234 SB 235 SB 236 SB 237 SB 238 SB 239 SB 240 SB 241 SB 242 SB 243 SB 244 SB 245 SB 246 SB 247 SB 248 SB 249 SB 250 SB 251 SB 252 SB 253 SB 254 SB 255 SB 256 SB 257 SB 258 SB 259 SB 260 SB 261 SB 262 SB 263 SB 264 SB 265  Amt. (Mg) 4.26 4.35 4.07 4.15 4.32 4.38 4.06 4.02 4.16 4.11 4.35 4.36 4.27 4.07 4.24 4.04 4.20 4.17 4.16 4.13 4.13 4.28 4.35 4.19 4.18 4.30 4.25 4.35 4.17 4.32 4.15 4.36 4.31 4.30 4.14 4.07 4.05 4.35  TOC% 0.77 1.67 0.84 0.80 1.27 0.94 0.57 0.88 0.69 0.98 0.94 0.87 0.87 0.95 0.94 0.99 0.81 0.92 1.04 0.97 1.04 1.12 1.19 1.16 1.23 1.26 1.21 1.19 1.15 1.08 1.23 1.10 1.07 1.02 1.03 1.00 0.90 0.98  δ 13Cadj -27.87 -29.51 -28.47 -28.60 -28.12 -28.16 -27.92 -27.86 -28.05 -27.50 -27.20 -27.28 -26.79 -26.84 -26.99 -27.09 -26.74 -27.03 -27.11 -27.34 -27.35 -27.72 -27.46 -27.35 -27.72 -27.68 -27.54 -27.60 -27.63 -27.53 -27.77 -27.38 -27.29 -27.34 -27.39 -27.37 -27.28 -27.42  depth (m) 627.07 626.97 626.87 626.77 626.67 626.57 626.47 626.37 626.27 624.23 624.13 624.03 623.93 623.83 623.73 623.63 623.53 623.43 623.33 623.23 623.13 623.03 622.93 622.83 622.73 622.63 622.53 622.43 622.33 622.23 622.13 622.03 621.93 621.83 621.73 621.63 621.53 621.43 284  Northern Alaska South Barrow #3 Core data core ID SB 266 SB 267 SB 268 SB 269 SB 270 SB 271 SB 272 SB 273 SB 274 SB 275 SB 276 SB 277 SB 278 SB 279 SB 280 SB 281 SB 282 SB 283 SB 284 SB 285 SB 286 SB 287 SB 288 SB 289 SB 290 SB 291 SB 292 SB 293 SB 294 SB 295 SB 296 SB 297 SB 298 SB 299 SB 300 SB 301 SB 302 SB 303  Amt. (Mg) 4.26 4.23 4.20 4.17 4.11 4.12 4.12 4.29 4.20 4.31 4.16 4.22 4.31 4.12 4.29 4.29 4.26 4.10 4.10 4.01 4.13 4.26 4.12 4.07 4.29 4.12 4.03 4.06 4.22 4.29 4.15 4.03 4.17 4.23 4.26 4.05 4.29 4.13  TOC% 0.97 0.93 0.83 1.14 0.93 1.40 1.40 1.07 0.98 1.03 1.07 1.30 1.12 1.06 1.09 1.14 1.05 1.11 1.12 1.04 1.10 1.10 0.98 1.83 1.05 1.30 1.07 1.14 1.33 1.27 1.24 1.23 1.11 1.15 1.05 1.33 1.10 1.14  δ 13Cadj -27.42 -27.35 -27.28 -27.36 -27.60 -26.95 -27.59 -27.60 -27.57 -27.50 -27.44 -27.33 -27.16 -27.20 -27.36 -27.48 -27.45 -27.59 -27.37 -27.34 -27.34 -27.20 -27.68 -27.22 -27.96 -28.26 -28.24 -28.39 -28.49 -28.28 -28.25 -28.21 -28.02 -27.97 -27.80 -27.79 -27.55 -27.52  depth (m) 621.33 621.23 621.13 621.03 620.93 620.83 620.73 620.63 620.53 620.43 620.33 620.23 618.13 618.03 617.93 617.83 617.73 617.63 617.53 617.43 617.33 617.23 617.13 617.03 616.93 616.83 616.73 616.63 616.53 616.43 616.33 616.23 616.13 616.03 615.93 615.83 615.73 615.63 285  Northern Alaska South Barrow #3 Core data core ID SB 304 SB 305 SB 306 SB 307 SB 308 SB 309 SB 310 SB 311 SB 312 SB 313 SB 314 SB 315 SB 316 SB 317 SB 318 SB 319 SB 320 SB 321 SB 322 SB 323 SB 324 SB 325 SB 326 SB 327 SB 328 SB 329 SB 330 SB 331 SB 332 SB 333 SB 334 SB 335 SB 336 SB 337 SB 338 SB 339 SB 340 SB 341  Amt. (Mg) 4.17 4.12 4.06 4.16 4.34 4.01 4.26 4.29 4.16 4.31 4.20 4.35 4.33 4.33 4.63 4.36 4.24 4.78 4.87 4.68 4.50 4.15 4.21 4.19 4.24 4.26 4.67 4.02 4.25 4.76 4.21 4.85 4.44 4.85 4.81 4.50 4.69 4.17  TOC% 1.03 0.95 1.26 1.47 1.35 1.44 1.45 1.28 1.32 1.36 1.34 1.10 1.02 0.77 0.88 0.96 1.00 0.85 0.90 1.15 1.10 1.07 1.04 2.27 1.18 1.04 1.06 1.07 0.91 0.98 0.97 0.95 1.01 0.86 0.86 0.92 0.92 0.85  δ 13Cadj -27.56 -27.19 -27.71 -28.10 -28.02 -28.00 -28.06 -27.93 -27.92 -27.90 -27.90 -27.93 -27.95 -28.13 -28.12 -28.32 -28.26 -28.29 -28.38 -28.43 -28.27 -27.99 -28.05 -28.24 -27.86 -27.53 -27.80 -27.74 -27.67 -27.49 -27.13 -27.43 -27.32 -27.36 -26.94 -27.48 -27.55 -27.19  depth (m) 615.53 615.43 615.33 615.23 615.13 615.03 614.93 614.83 614.73 614.63 614.53 614.43 614.33 614.23 614.13 614.03 613.93 613.83 613.73 613.63 613.53 613.43 613.33 613.23 613.13 613.03 612.93 612.83 612.73 612.63 612.53 612.43 612.33 612.23 612.13 612.03 611.93 611.83 286  Northern Alaska South Barrow #3 Core data core ID SB 342 SB 343 SB 344 SB 345 SB 346 SB 347 SB 348 SB 349 SB 350 SB 351 SB 352 SB 353 SB 354 SB 355 SB 356 SB 357 SB 358 SB 359 SB 360 SB 361 SB 362 SB 363 SB 364 SB 365 SB 366 SB 367 SB 368 SB 369 SB 370 SB 371 SB 372 SB 373 SB 374 SB 375 SB 376 SB 377 SB 378 SB 379  Amt. (Mg) 4.91 4.36 4.17 4.81 4.26 4.83 4.85 4.29 4.05 4.60 4.77 4.78 4.60 4.27 4.32 4.24 4.61 4.58 4.27 4.28 4.19 4.76 4.46 4.12 4.91 4.77 4.51 4.19 4.33 4.23 4.39 4.92 4.71 4.44 4.29 4.89 4.35 4.18  TOC% 0.83 0.84 0.89 0.79 0.77 0.79 0.78 0.75 0.76 0.72 0.64 0.81 0.67 1.10 0.96 0.92 0.96 0.92 0.96 0.96 0.82 0.89 0.95 0.86 0.96 0.83 0.84 0.86 0.81 0.82 0.82 0.85 0.81 0.84 0.74 0.80 0.97 0.93  δ 13Cadj -27.12 -27.01 -27.34 -26.64 -26.76 -27.23 -26.87 -26.90 -27.06 -26.44 -27.03 -27.28 -27.37 -27.59 -27.41 -27.48 -27.39 -27.18 -27.51 -27.36 -27.47 -27.97 -27.59 -27.24 -27.46 -27.53 -27.57 -27.19 -27.42 -27.45 -27.45 -27.38 -27.26 -27.14 -27.37 -27.23 -27.25 -27.23  depth (m) 611.73 611.63 611.53 611.43 611.33 611.23 611.13 608.68 608.58 608.48 608.38 608.28 608.18 608.08 607.98 607.88 607.78 607.68 607.58 607.48 607.38 607.28 607.18 607.08 606.98 606.88 606.78 606.68 606.58 606.48 606.38 606.28 606.18 606.08 605.98 605.88 605.78 605.68 287  Northern Alaska South Barrow #3 Core data core ID SB 380 SB 381 SB 382 SB 383 SB 384 SB 385 SB 386 SB 387 SB 388 SB 389 SB 390 SB 391 SB 392 SB 393 SB 394 SB 395 SB 396 SB 397  Amt. (Mg) 4.66 4.37 4.90 4.64 4.52 4.59 4.48 4.48 4.50 4.20 4.08 4.56 4.43 4.69 4.17 4.04 4.76 4.62  SB 398 SB 399 SB 400 SB 401 SB 402 SB 403 SB 404 SB 405 SB 406 SB 407 SB 408 SB 409 SB 410 SB 411 SB 412 SB 413 SB 414 SB 415 SB 416  4.38 4.76 4.72 4.79 4.06 4.66 4.07 4.17 4.28 4.26 4.27 4.18 4.31 4.19 4.16 4.31 4.10 4.25 4.36  TOC% 0.85 0.93 0.87 0.93 1.01 0.86 0.81 0.87 0.95 0.92 0.87 0.68 0.87 0.78 0.91 1.02 0.95 0.97 GAP 2 0.64 0.63 0.61 0.72 0.67 0.56 0.71 0.75 0.67 0.73 0.72 0.65 0.80 0.61 0.72 0.79 0.93 0.99 0.69  δ 13Cadj -27.07 -27.34 -27.29 -27.36 -27.20 -26.90 -27.10 -27.08 -27.30 -27.28 -27.17 -27.11 -27.39 -27.46 -27.19 -27.20 -27.55 -27.44  depth (m) 605.58 605.48 605.38 605.28 605.18 605.08 604.98 604.88 604.78 604.68 604.58 604.48 604.38 604.28 604.18 604.08 603.98 603.88  -26.63 -26.98 -27.29 -27.63 -26.89 -26.72 -26.27 -26.84 -26.93 -26.93 -27.06 -26.86 -26.96 -26.09 -26.65 -27.06 -27.53 -28.04 -26.39  548.03 547.98 547.93 547.88 547.83 547.78 547.73 547.68 547.63 547.58 547.53 547.48 547.43 547.38 547.33 547.28 547.23 547.18 547.13 288  Northern Alaska South Barrow #3 Core data core ID SB 417 SB 418 SB 419 SB 420 SB 421 SB 422 SB 423 SB 424 SB 425 SB 426 SB 427 SB 428 SB 429 SB 430 SB 431 SB 432 SB 433 SB 434 SB 435 SB 436 SB 437 SB 438 SB 439 SB 440 SB 441 SB 442 SB 443 SB 444 SB 445 SB 446 SB 447 SB 448 SB 449 SB 450 SB 451 SB 452 SB 453 SB 454  Amt. (Mg) 4.03 4.31 4.34 4.40 4.32 4.29 4.16 4.35 4.23 4.33 4.09 4.36 4.23 4.37 4.24 4.15 4.18 4.38 4.10 4.17 4.27 4.38 4.36 4.38 4.31 4.04 4.25 4.35 4.35 4.36 4.22 4.07 4.17 4.18 4.06 4.34 4.03 4.11  TOC% 0.80 0.72 0.81 0.77 0.73 0.73 0.69 0.68 0.80 0.70 0.80 0.63 0.69 0.75 0.51 0.74 0.68 0.64 0.66 0.70 0.79 0.93 0.61 0.68 0.63 0.73 0.66 0.74 0.59 0.85 0.68 0.70 0.71 0.68 0.74 0.74 0.75 0.82  δ 13Cadj -27.04 -27.02 -27.66 -26.83 -27.64 -27.00 -26.51 -26.65 -27.07 -26.60 -26.82 -26.75 -26.44 -27.05 -26.64 -27.00 -26.29 -27.04 -26.77 -26.98 -27.45 -27.36 -26.39 -26.87 -26.86 -27.26 -27.08 -27.03 -27.20 -27.11 -26.61 -27.04 -26.86 -27.59 -27.41 -27.11 -27.30 -27.75  depth (m) 547.08 547.03 546.98 546.93 546.88 546.83 546.78 546.73 546.68 546.63 546.58 546.53 546.48 546.43 546.38 546.33 546.28 546.23 546.18 546.13 546.08 546.03 545.98 545.93 545.88 545.83 545.78 545.73 545.68 545.63 545.58 545.53 545.48 545.43 545.38 545.33 545.28 545.23 289  Northern Alaska South Barrow #3 Core data core ID SB 455 SB 456 SB 457 SB 458 SB 459 SB 460 SB 461 SB 462 SB 463 SB 464 SB 465 SB 466 SB 467 SB 468 SB 469 SB 470 SB 471 SB 472 SB 473 SB 474 SB 475 SB 476 SB 477 SB 478 SB 479 SB 480 SB 481 SB 482 SB 483 SB 484 SB 485 SB 486 SB 487 SB 488 SB 489 SB 490 SB 491 SB 492  Amt. (Mg) 4.23 4.15 4.10 4.25 4.18 4.25 4.08 4.37 4.25 4.34 4.31 4.14 4.29 4.25 4.30 4.01 4.19 4.07 4.06 4.32 4.34 4.11 4.36 4.07 4.28 4.28 4.30 4.38 4.03 4.18 4.26 4.12 4.23 4.37 4.04 4.39 4.18 4.06  TOC% 0.75 0.77 0.68 0.75 0.43 0.67 0.80 0.76 0.76 0.89 0.67 0.74 0.75 0.72 0.77 0.62 0.72 0.73 0.68 0.73 0.75 0.85 0.85 0.74 0.87 0.80 0.83 0.64 0.77 0.76 0.71 0.75 0.70 0.72 0.70 0.65 0.76 0.69  δ 13Cadj -27.50 -27.20 -27.64 -27.14 -28.29 -27.90 -27.91 -27.76 -27.75 -27.80 -27.99 -27.19 -27.22 -27.76 -26.86 -27.22 -27.33 -27.11 -26.46 -26.69 -26.92 -26.70 -27.12 -27.05 -27.46 -27.46 -27.20 -27.69 -27.35 -27.47 -27.17 -27.58 -27.93 -27.48 -27.29 -27.57 -27.52 -28.28  depth (m) 545.18 545.13 545.08 545.03 544.98 544.93 544.88 544.83 544.78 544.73 544.68 544.63 544.58 544.53 544.48 544.43 544.38 544.33 544.28 544.23 544.18 544.13 544.08 544.03 543.98 543.93 543.88 543.83 543.78 543.73 543.68 543.63 543.58 543.53 543.48 541.93 541.88 541.83 290  Northern Alaska South Barrow #3 Core data core ID SB 493 SB 494 SB 495 SB 496 SB 497 SB 498 SB 499 SB 500 SB 501 SB 502 SB 503 SB 504 SB 505 SB 506 SB 507 SB 508 SB 509 SB 510 SB 511 SB 512 SB 513 SB 514 SB 515 SB 516 SB 517 SB 518 SB 519 SB 520 SB 521 SB 522 SB 523 SB 524 SB 525 SB 526 SB 527 SB 528 SB 529 SB 530  Amt. (Mg) 4.35 4.29 4.09 4.40 4.28 4.30 4.25 4.23 4.32 4.19 4.23 4.10 4.22 4.32 4.04 4.07 4.28 4.23 4.35 4.23 4.36 4.05 4.30 4.07 4.05 4.28 4.09 4.25 4.05 4.35 4.28 4.36 4.21 4.17 4.27 4.25 4.40 4.38  TOC% 0.87 0.80 0.84 0.68 0.59 0.68 0.82 0.71 0.75 0.68 0.85 0.68 0.65 1.11 0.72 0.69 0.76 0.81 0.71 0.67 0.70 0.65 0.61 0.66 0.69 0.68 0.82 0.74 0.85 0.83 0.99 0.88 1.00 1.05 0.92 0.85 0.90 0.75  δ 13Cadj -27.71 -27.75 -28.10 -27.53 -28.11 -28.13 -28.08 -27.82 -27.87 -27.64 -28.05 -27.93 -27.62 -28.03 -28.04 -27.49 -27.89 -27.81 -27.14 -27.37 -27.62 -27.53 -27.32 -27.24 -27.50 -27.22 -27.91 -27.47 -27.76 -27.73 -27.98 -27.46 -27.58 -28.00 -27.95 -27.87 -27.31 -27.61  depth (m) 541.78 541.73 541.68 541.63 541.58 541.53 541.48 541.43 541.38 541.33 541.28 541.23 541.18 541.13 541.08 541.03 540.98 540.93 540.88 540.83 540.78 540.73 540.68 540.63 540.58 540.53 540.48 540.43 540.38 540.33 540.28 540.23 540.18 540.13 540.08 540.03 539.98 539.93 291  Northern Alaska South Barrow #3 Core data core ID SB 531 SB 532 SB 533 SB 534 SB 535 SB 536 SB 537 SB 538 SB 539 SB 540 SB 541 SB 542 SB 543 SB 544 SB 545 SB 546 SB 547 SB 548 SB 549 SB 550 SB 551 SB 552 SB 553 SB 554 SB 555 SB 556 SB 557 SB 558 SB 559 SB 560 SB 561 SB 562 SB 563 SB 564 SB 565 SB 566 SB 567 SB 568  Amt. (Mg) 4.38 4.26 4.15 4.20 4.34 4.31 4.22 4.35 4.14 4.04 4.26 4.15 4.23 4.17 4.21 4.11 4.05 4.37 4.08 4.22 4.21 4.09 4.28 4.37 4.30 4.36 4.25 4.26 4.22 4.06 4.27 4.31 4.27 4.12 4.17 4.20 4.27 4.31  TOC% 0.94 0.87 0.83 0.74 1.01 0.91 0.79 0.75 0.94 0.85 0.86 0.64 1.02 0.80 0.80 0.68 0.93 0.95 0.68 0.75 0.89 0.67 0.78 0.76 0.79 1.01 1.07 1.12 1.09 0.88 0.95 0.94 1.04 0.94 0.99 0.85 0.98 1.25  δ 13Cadj -27.73 -27.73 -27.65 -27.83 -27.76 -27.61 -27.84 -27.08 -27.48 -27.76 -28.02 -27.76 -27.45 -27.58 -27.22 -27.35 -27.90 -27.68 -27.62 -27.40 -27.31 -27.11 -27.81 -27.64 -27.42 -27.90 -27.77 -27.94 -27.91 -27.77 -27.77 -27.38 -27.54 -27.80 -27.52 -28.05 -27.89 -27.94  depth (m) 539.88 539.83 539.78 539.73 539.68 539.63 539.58 539.53 539.48 539.43 539.38 539.33 539.28 539.23 539.18 539.13 539.08 539.03 538.98 538.93 538.88 538.83 538.78 538.73 538.68 538.63 538.58 538.53 538.48 538.43 538.38 538.33 538.28 538.23 538.18 538.13 535.83 535.78 292  Northern Alaska South Barrow #3 Core data core ID SB 569 SB 570 SB 571 SB 572 SB 573 SB 574 SB 575 SB 576 SB 577 SB 578 SB 579 SB 580 SB 581 SB 582 SB 583 SB 584 SB 585 SB 586 SB 587 SB 588 SB 589 SB 590 SB 591 SB 592 SB 593 SB 594 SB 595 SB 596 SB 597 SB 598 SB 599 SB 600 SB 601 SB 602 SB 603  Amt. (Mg) 4.40 4.30 4.17 4.16 4.20 4.04 4.24 4.08 4.22 4.11 4.19 4.25 4.31 4.18 4.28 4.39 4.32 4.11 4.33 4.13 4.38 4.20 4.34 4.04 4.36 4.25 4.26 4.11 4.28 4.35 4.01 4.25 4.15 4.12 4.31  TOC% 0.71 0.90 1.03 1.02 1.06 1.13 0.79 0.38 1.01 0.90 0.93 0.90 0.77 0.82 0.76 0.81 0.77 0.76 0.79 0.99 0.84 0.82 0.77 0.82 0.83 0.88 0.89 0.97 0.76 0.67 0.82 0.87 0.81 0.69 0.77  δ 13Cadj -28.14 -27.93 -27.79 -27.87 -27.82 -27.87 -27.70 -28.75 -27.67 -27.79 -27.45 -27.60 -28.04 -27.74 -27.59 -27.73 -27.35 -27.48 -27.96 -27.61 -27.46 -28.02 -28.18 -27.80 -27.59 -28.06 -27.75 -27.91 -27.97 -27.46 -27.52 -27.94 -27.91 -27.75 -27.44  depth (m) 535.73 535.68 535.63 535.58 535.53 535.48 535.43 535.38 535.33 535.28 535.23 535.18 535.13 535.08 535.03 534.98 534.93 534.88 534.83 534.78 534.73 534.68 534.63 534.58 534.53 534.48 534.43 534.38 534.33 534.28 534.23 534.18 534.13 534.08 534.03  293  

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