Characterizing Oil, Gas Hydrate, and Sedimentary Systems Via

CHARACTERIZING OIL, GAS HYDRATE, AND SEDIMENTARY SYSTEMS VIA

GEOCHEMISTRY, COMPUTATIONAL MODELING, AND GLOBAL SYNTHESIS

A DISSERTATION

SUBMITTED TO THE DEPARTMENT OF

GEOLOGICAL AND ENVIRONMENTAL SCIENCES

AND THE COMMITTEE ON GRADUATE STUDIES

OF STANFORD UNIVERSITY

IN PARTIAL FULFILLMENT OF THE REQUIREMENTS

FOR THE DEGREE OF

DOCTOR OF PHILOSOPHY

ZACHARY FLORENTINO MURGUIA BURTON

DECEMBER 2020

This work is licensed under a Creative Commons Attribution- Noncommercial 3.0 United States License. http://creativecommons.org/licenses/by-nc/3.0/us/

This dissertation is online at: http://purl.stanford.edu/yt550pv4632

ii I certify that I have read this dissertation and that, in my opinion, it is fully adequate in scope and quality as a dissertation for the degree of Doctor of Philosophy.

Stephan Graham, Primary Adviser

I certify that I have read this dissertation and that, in my opinion, it is fully adequate in scope and quality as a dissertation for the degree of Doctor of Philosophy.

J. Moldowan

I certify that I have read this dissertation and that, in my opinion, it is fully adequate in scope and quality as a dissertation for the degree of Doctor of Philosophy.

Tapan Mukerji

Approved for the Stanford University Committee on Graduate Studies. Stacey F. Bent, Vice Provost for Graduate Education

This signature page was generated electronically upon submission of this dissertation in electronic format. An original signed hard copy of the signature page is on file in University Archives.

iii ABSTRACT

At the highest level, this dissertation investigates marine systems from a geochemical and sedimentological perspective through the undertaking of six distinct and only somewhat tenuously interrelated contributions. Nonetheless, the work detailed herein might be thought of as fitting, very broadly speaking, under two central themes, being (i) investigations into the formation, evolution, and destruction of hydrocarbon systems and

(ii) advancement and refinement of the understanding of global controls on deep-sea clastic depositional systems. These investigations have important bearing on the interplay of climate, tectonics, and paleoceanography and attendant modulation of deep-marine depositional systems and the oceanic carbon cycle at local, regional, and global scales, and throughout Earth history.

Chapter 1 of this dissertation presents an organic geochemical assessment of petroleum quality using petroleum seep samples from frontier exploration regions along the east coast of New Zealand’s North Island and northern South Island. We analyze biomarker compounds to assess the degree of biodegradation to which these samples have been subjected, and use biomarker and diamondoid analysis to estimate the thermal maturity of these oils and to identify petroleum mixing that may have occurred during migration and accumulation of these oils. Assessment of biodegradation based upon distributions of n- alkane, isoprenoid, sterane, and terpane compounds indicates negligible biodegradation of oils from the northern portion of our study area and low to moderate biodegradation in southern oils, with measurement of 25-norhopane further suggesting that biodegradation in southern oils may be solely aerobic. Taken together, these findings suggest fresh, active

iv seepage (and thus, petroleum systems) in the north, and suggest that the quality of potential subsurface accumulations throughout the studied regions is potentially unaffected by biodegradation. Assessment of thermal maturity, oil-to-gas conversion, and hydrocarbon mixing using sterane, terpane, and diamondoid compounds reveals three distinct petroleum mixtures among the seeps, including components generated throughout the oil window as well as intensely cracked condensate-wet gas components. Identification of black oil components might indicate the presence of actively producing source rock in all regions represented by these seeps, while the intensely cracked components indicate petroleum mixing via thermogenic gas infiltration, and thus suggest an impact on oil quality.

Chapter 1 has been published in Energy & Fuels (2018, vol. 32, p. 1287–1296) with coauthors Mike Moldowan, Richard Sykes, and Steve Graham, all of whom provided input and discussion important to the scope, design, and implications of the study. Mike

Moldowan provided me the training and laboratory space necessary to the execution of the experimental portion of this work, and was particularly helpful in guiding and informing my interpretations of the results described herein. Richard Sykes provided me with the samples analyzed herein, courtesy of longtime Stanford geosciences collaborator GNS

Science in New Zealand, and also provided key contextual information and public data on petroleum geochemistry of New Zealand oil and gas samples. Steve Graham was the one who, alongside Jared Gooley, first suggested I consider undertaking investigation of the frontier basins off New Zealand’s east coast of New Zealand, supported my early forays into petroleum geochemistry, and provided financial support for this study.

Chapter 2 presents a geochemical assessment of petroleum source rock depositional environments and age based upon expanded analysis of the set of New Zealand oil seep

v samples investigated in Chapter 1. We use biomarker distributions, stable carbon isotope distributions, and sulfur concentrations of these oil seep samples to interpret source rock characteristics including type of organic matter input, redox conditions, sedimentary facies, and age. Results show that samples generally cluster into two groups on the basis of these characteristics. These groups correlate with geographic location of seeps—namely, a distinct northern group and southern group emerge. While source rocks associated with all seep samples are interpreted to be marine, results suggest northern samples had more terrigenous organic matter input to their source rock(s), while southern samples had more marine input. Results suggest northern sample source rock(s) had more oxic depositional environments, whereas southern sample source rock(s) had more reducing environments.

A shale source rock sedimentary facies was indicated for all samples. These observations suggest that southern samples may be derived from slightly higher quality source rocks

(i.e., higher hydrogen index, deposited in more reducing conditions), although source rocks in both regions are oil prone. Biomarker age parameters suggest that the northern oil samples are from a younger (Cenozoic) source rock, whereas the southern oil samples are from an older (Cretaceous) source rock. Taken together, source rock characteristics

(depositional environment and age) examined herein indicate the presence of at least two different source rocks—and, advance our understanding of paleodepositional environments and prospective petroleum systems in this underexplored frontier setting.

Chapter 2 has been published in International Journal of Earth Sciences (2019, vol. 108, p. 1079–1091) with coauthors Mike Moldowan, Les Magoon, Richard Sykes, and Steve

Graham. Coauthor contributions to this study are as detailed for Chapter 1, above, with additional key contributions by Les Magoon, who provided useful discussion, framing, and

vi input—particularly with regard to assessment of oil seeps and correlative source rocks from a holistic petroleum systems perspective.

Chapter 3 presents findings from an integrated computational modeling (i.e., basin and petroleum system modeling) study of the impact of tectonic uplift on the stability of submarine gas hydrate systems. We use two-dimensional modeling of a structurally restored transect from the Hikurangi margin of New Zealand (a region with confirmed presence of subsea gas hydrates) to investigate the potential for tectonic shortening and uplift to drive changes in the thickness and extent of the gas hydrate stability zone. We predict evolution of the gas hydrate stability zone from the late Oligocene to present, and demonstrate substantial (~70%) decreases in the extent of the gas hydrate stability zone over this period of tectonic uplift and attendant decreases in water depth along the modeled transect. These results provide modeling-based validation that tectonic uplift can destabilize subsea gas hydrate, with important implications for ocean chemistry and the carbon cycle, particularly during periods of increased plate convergence.

Chapter 3 has been published in Geophysical Research Letters (2020, vol. 47) with coauthors Karsten Kroeger, Allegra Hosford Scheirer, Yongkoo Seol, Blair Burgreen-

Chan, and Steve Graham, all of whom contributed to discussion of and publication of this work. Karsten Kroeger provided key discussion of gas hydrate systems and input to the overall framing of this study, with particularly helpful input regarding the modeling of gas hydrates, gas hydrate and sedimentary depositional systems of New Zealand, and interpretation of the gas hydrate stability zone from seismic datasets of offshore New

Zealand. Allegra Hosford Scheirer suggested the idea of using Blair Burgreen-Chan’s

Hikurangi margin basin model to explore model capabilities with regard to characterization

vii of gas hydrate systems, and alongside other instructors in our group, was instrumental to my training in the use of basin and petroleum system modeling software. Yongkoo Seol provided insight into gas hydrate system dynamics (especially gas hydrate stability) and has provided financial support of my gas hydrate research efforts since late 2016. Blair

Burgreen-Chan achieved the tremendous task of constructing the two-dimensional model

(and performing the painstakingly detailed structural restorations incorporated into this model) that was adapted in this study for investigation of gas hydrate systems. Steve

Graham provided unwavering encouragement of my pursuit of this novel approach to investigating gas hydrate stability from the time I first embarked on this work in early 2018, and also provided financial support.

Chapter 4 presents a global compilation of sediment organic properties at marine biogenic gas hydrate localities to shed light on the minimum organic carbon and hydrogen contents required for the formation of gas hydrate—given that organic matter in shallow sediments is requisite as the feedstock for microbial metabolism and resultant generation of methane.

We undertake a comprehensive synthesis of published total organic carbon (TOC) and hydrogen index (HI) values from analysis by previous workers of shallow (up to ~2 km subseafloor) sediment at marine gas hydrate localities. We compile data from 43 known gas hydrate sites along five continents, which include average site-by-site TOC values ranging from 0.3–3+ wt. % (compatible with previous studies suggesting a minimum TOC of 0.3–0.5 wt. % needed for gas hydrate formation) and an average TOC for all sites of 1.2 wt. % (n = 2000+ total analyses). HI values range from 50–400+ mg hydrocarbon/g TOC, suggesting an HI of 50 may result in gas hydrate formation. Average HI for all sites is 170

(n = 500+ total analyses). This compilation provides both understanding of average organic

viii properties of gas hydrate-associated sediment as well as possible empirical evidence of the minimum TOC and HI necessary for the formation of gas hydrate.

Chapter 4 is in preparation for submission to Marine and Petroleum Geology with coauthors Allegra Hosford Scheirer and Ken Peters, both of whom have provided helpful revisions of the manuscript. Allegra Hosford Scheirer conceived of the idea of compiling a catalogue of gas hydrate-associated sediment organic properties to be used in the basin and petroleum system modeling of gas hydrate systems. Ken Peters has provided critical input on biogenic gas systems and sedimentary organic geochemistry.

Chapter 5 presents a global compilation of deep-marine sand-rich systems deposited during the exceptionally warm (i.e., hothouse and greenhouse) climates of the Paleocene through middle Eocene to investigate the nature of deep-sea siliciclastic deposition during periods of high global sea level—which should, due to resultant maximization of shelfal accommodation space, be times of minimal sand-rich deposition in deep-water passive margin settings, according to traditional notions of deep-water deposition. We focus in particular on deep-sea sand-rich systems of early Eocene age, as the early Eocene

“hothouse” has been identified as the warmest extended climate interval of the Cenozoic, and, experienced the highest global mean sea level of the Cenozoic. We document 59 examples of early Eocene turbidite systems from all continental margins except Antarctica.

Sand-rich systems were widespread on active margins (42 examples), but were also present on passive margins (17 examples). Along passive margins, 13 of 17 (~76%) turbidite systems are associated with significant Eocene-age fluvial systems. We find that sediment from active margins and sediment delivered to passive margins by integrated drainages and elevated hothouse denudation cause sediment supply to exceed shelf accommodation

ix despite exceptionally high global sea level. We therefore suggest that—in addition to the paradigm of eustasy—tectonism and climate-driven change (e.g., drainage integration, global rainfall, denudation) may significantly influence global distribution of coarse- grained deep-marine depositional systems, and may overwhelm eustatic control entirely, especially during periods of high sea level.

Chapter 5 is in revision for resubmission to Geology with coauthors Tim McHargue, Tyler

Kukla, Chris Kremer, Roger Bloch, Jared Gooley, Chayawan Jaikla, Jake Harrington, and

Steve Graham, all of whom took part in a seminar arranged by Tim McHargue to investigate this topic in April through June of 2017. Continued discussion and work with

Tim McHargue and Steve Graham has been central to the completion of this work. Tyler

Kukla provided especially critical input on climate history and climate dynamics and developed the conceptual model that we present herein to describe the linkage between global climate and deep-marine sand-rich system occurrence. Chris Kremer provided suggestions especially helpful to the writing of this manuscript.

Chapter 6 presents a global compilation of unconformities that occur at the Eocene-

Oligocene boundary within deep-marine clastic successions (mostly within sedimentary basins). The Eocene-Oligocene boundary marks a major transition from the warm, largely ice-free conditions (i.e., hothouse and greenhouse worlds) of the Eocene to the cool, glaciated conditions (i.e., icehouse world) that have typified the Oligocene to recent times.

Attendant to this profound climatic transition was the largest fall in global mean sea level in the entirety of the Cenozoic (excepting perhaps during glacial-interglacial cycles of the

Quaternary), as well as major changes to ocean circulation (including significant intensification of deep ocean circulation). Extensive deep-marine hiatuses of this age have

x been documented within pelagic successions investigated during activities of the Deep Sea

Drilling Project, but, documentation of interregional Eocene-Oligocene hiatuses within clastic successions is lacking at a global scale. We undertake this effort by synthesizing results from published work. We document 93 distinct locations—along the margins of all seven continents—where unconformities of Eocene-Oligocene age have been recorded in the literature. These include 28 locations along the African continental margin, 1 along the

Antarctic margin, 12 along Asian margins, 12 in the Oceania region, 12 along European margins, 19 along the North American margin, and 9 along the South American margin.

Of the 93 total locations, 37 are in passive margin settings and 56 are in active margin settings or passive margin settings with local tectonism. Taken together, we suggest that these unconformities are clear manifestation of major oceanographic change at the Eocene-

Oligocene boundary. We suggest that some combination of the activity of intensified deep- marine currents and of the influence of pronounced sea-level fall resulted in globally widespread erosion of deep-marine clastic sediments during this time, while noting that a possible sea-level control must go beyond traditional notions of emergence-driven unconformity development (given that the sections investigated occur in deep-marine sediments, i.e., would have been far beneath the level of any subaerial exposure driven by the estimated ~55-meter drop in sea level at the Eocene-Oligocene Transition). The relative significance and plausibility of these ocean circulation and falling sea level scenarios, as well as of other potential mechanisms such as nondeposition, call for further investigation.

Chapter 6 is in preparation as an invited submission for Earth-Science Reviews with coauthors Tim McHargue and Steve Graham, both of whom have been instrumental in the discussion, design, and interpretation of this work.

ACKNOWLEDGEMENTS

As a crisp autumn sunrise again hints at the rapidly-expiring eleventh hour of my time at

Stanford, it is wholly impossible to even begin to reflect on what this past half-decade (and three months) has represented to me without tears immediately rendering futile any effort at a measured, logical collection of nods to those who have meant the world to me these past few years. And so, I apologize in advance for the inevitably incomplete attempt that here follows.

First, a deep thanks to the Department of Geological Sciences and Stanford School of

Earth, Energy, and Environmental Sciences staff who keep the cogs of our tight-knit scientific and social community well-oiled and ever-whirling, and without whom none of my own research efforts, conference travel, or holiday party planning would ever have come to fruition. A special and non-exhaustive shout-out to Alyssa Ferree, Yvonne Lopez,

Stephanie James, Julie Hitchcock, Claudia Baroni, Lauren Nelson, Debbie Spaizman,

Lauren Mendoza-Tabinas, Susan French, Sylvia Wasmer, Kelly Wells, and Javier Illueca.

Thank you in particular for your kindness, compassion, and guidance to all of us bewildered and disorganized graduate students near-drowning in mountains of crumpled and soiled receipts—I hope, at least, I earned a bit of good karma by never coming to any of you with a Patagonia reimbursement request.

I would also like to express my sincere gratitude to Stanford’s Basin and Petroleum System

Modeling Industrial Affiliates Program and to the U.S. Department of Energy National

Energy Technology Laboratory for providing the funding that made this research—and my time at Stanford—possible. Thanks to the McGee/Levorsen Research Grant Program, xii

AAPG Foundation Grants-in-Aid Program, Gordon Research Conference Carl Storm

Fellowship Program, and Shell Foundation Grant Program for additional funding specifically supporting the research documented herein, as well as to the Geological

Society of America, Clay Minerals Society, SETI Institute, National Science Foundation, and NASA Astrobiology Institute for supporting additional geochemistry and geology research I have conducted while at Stanford.

Of course, at the very core of my ability to have undertaken the research described below has been collaboration and discussion with many brilliant scientists here at Stanford. I’d like to especially acknowledge my doctoral committee (Steve Graham, Allegra Hosford

Scheirer, Mike Moldowan, Tim McHargue, and Tapan Mukerji—with special thanks to

Tony Kovscek for chairing my defense), Les Magoon, Ken Peters, Alan Burnham, Kris

Meisling, Noelle Schoellkopf, Don Lowe, Matt Malkowski, Simon Klemperer, Nader

Dutta, George Hilley, and Jeremy Dahl as astounding faculty I’ve had the privilege to work with and learn from, and, to especially acknowledge “official” collaboration with fellow students Jared Gooley, Chris Kremer, Tyler Kukla, Earth Jaikla, and Jake Harrington on the early Eocene, Laura Dafov on gas hydrate systems, Jared and Zach Sickmann for my earliest experiences with New Zealand’s geology, and Inessa Yurchenko and Will

Thompson for camaraderie through biomarker lab work juggling many, many samples— while at the moment refraining from delving into the tremendous influence so many other fellow students and postdocs at Stanford have had specifically on my research and on my growth as a scientist.

Also central to my completion of the research detailed herein have been the contributions and support of scientific collaborators outside of Stanford. My New Zealand work

xiii benefited from close collaboration with the terrific folks at GNS Science in Wellington

(special thanks to Greg Browne, Karsten Kroeger, Richard Sykes, Kyle Bland, Gareth

Crutchley, Rob Funnell, and Peter King), from extended discussions with folks at

Chevron’s Houston and Perth offices (thanks to Gary Muscio and Nathan Palmer, in particular), and from the input of BPSM alumna Blair Burgreen-Chan of ConocoPhillips in Houston. The staff at Biomarker Technologies Incorporated in Rohnert Park provided me with essential training and support toward the completion of my New Zealand oil seep work, while those at Infometrix in Seattle and GeoMark Research in Houston provided additional support to my petroleum geochemistry-related endeavors. I would not have been afforded the opportunity to work on gas hydrate systems without the support of collaborator

Yongkoo Seol of DOE NETL’s Morgantown location. Much appreciation is owed to folks at Schlumberger’s Aachen office for PetroMod support, to Roger Bloch for work together on the early Eocene, and to Tim Collett and Seth Haines of the USGS’s Denver office, Ray

Boswell of DOE NETL, James Cassanelli of Occidental Petroleum’s Houston office, and the Japan Oil, Gas and Metals National Corporation for continued discussion of gas hydrates. Lastly, thanks are again due to the BPSM and SPODDS groups at large (affiliates, faculty, staff, postdocs, students, collaborators, and more) for their role in my development and completion of this research.

In the realm of professional development in the geology sector but outside of the pages of this dissertation, I learned immense amounts—under the indefatigable guidance of mentor

(and now, dear friend) Lara Heister—about source rock geochemistry and making it to the office by 9:00 a.m. (well, mostly) during a summer internship at Anadarko Petroleum

Corporation in Houston and benefited from a fall internship at California Resources

xiv

Corporation in Bakersfield (particularly by way of Hendrick’s plus thinly-sliced cucumber administered under the careful supervision of dear friend Lisa Alpert). And, in the realm of scientific inquiry, I have treasured few experiences more than my work as a volunteer research assistant for Janice Bishop, with whom I have been investigating the geochemistry and mineralogy of Mars analogue sediments since 2017 at the SETI Institute and through the NASA Astrobiology Institute, and who has been (and very much continues to be) one of my most influential scientific mentors. Outside the sciences sensu stricto—but still largely within the realm of my development as a research scientist these past few years—I owe my deep gratitude to Sonali Rammohan for mentoring me through the course of a summer research internship with the Graduate School of Business’s Institute for Innovation in Developing Economies (Stanford Seed), all while managing our massive time difference between her Stanford office and my coworking space in Chennai, India (thanks as well to

Chitra Ravi of Chrysalis for working closely with me that summer, and for our continued dialogue). Alongside Janice and my doctoral committee, there are few other mentors who have been as influential to my own development while at Stanford as Dian Grueneich. Dian took a chance in taking a geologist under her wing—she was my mentor from day one in launching and leading Stanford’s Energy Policy Community, brought me on for nine months of policy research at the Precourt Institute for Energy, and continues to advise and encourage me along my professional journey

As far as the foundations to my becoming a Stanford geology graduate student (and almost- graduate) and scientist, I owe my deepest thanks to the amazing, amazing professors from my days in Bowdoin College’s Earth and Oceanographic Science department—beginning with Peter Lea’s and Phil Camill’s intro courses, to the constant support, extremely

xv challenging coursework, and encouragement from sage “department mothers” (i.e., geologist Rachel Beane and oceanographer Collin Roesler), and of course, my sensational undergraduate thesis advisor Emily Peterman, who set alight my love for geologic research.

Thanks as well to long, sweaty summer days at the Indiana University Geologic Field

Station in Montana and during the GSA/ExxonMobil field seminar in Bighorn Basin, and to late nights at UCSB’s LA-ICP-MS facility for helping me cut my teeth in 2014 and 2015.

Finally, in reflecting on my research at Stanford and on my evolution as a scientist, my doctoral committee deserves the largest thanks and deepest gratitude I can possibly muster…

To Tapan Mukerji: Though we did not work directly on my dissertation research covered in the chapters below, your influence on my scientific development as part of the BPSM group is something I value deeply. Your tremendous facility to shed light on the physical underpinnings of the natural systems subject to our group’s lines of inquiry (perhaps best emblazoned upon my memory as a recollection of hours spent auditing your numerical methods course, wherein you somehow managed to get even me to manipulate COMSOL

Multiphysics with some fluency) is an attribute that—alongside your observations and recommendations regarding my scientific journey over these past few years—I can only describe as sagacious. I hope to take with me your ability to objectively and critically consider natural systems with ever-fresh eyes and openness to new perspectives and possibilities. Thank you, Tapan.

To Tim McHargue: I could not have predicted that what began during one of your illustrious “Timinars” in early 2017 would—after going near-dormant for over a year and a half—become an all-consuming realization of what I naively proposed to Steve during a xvi

March 2015 visit in what now feels like a different lifetime: the elusive and generally ill- advised (though terrifically executed by academic predecessor and collaborator Blair)

BPSM × SPODDS dissertation. But, I won’t pretend for a geologically instantaneous moment that there isn’t a great deal of unfinished work to be done: 1) The early Eocene will see the light of day—and, soon! 2) The Eocene-Oligocene boundary shall be missing in the sedimentary record of deep-marine clastic basins no longer—that is to say, it shall at least no longer be missing in the record of the missing sedimentary record! 3) Contourites!

4) Eel Canyon?! Okay, I promise completion of at least two of these items by the close of

2021—or else, I am henceforth sworn to make an earnest attempt (beginning on January

1st, 2022) to supply you with at least one (unopened) Pepsi any and every time I set foot within a one-hour radius of the Bay Area, broadly defined. Thanks, Tim.

To Mike Moldowan: I feel compelled to here admit to you for the first time that, throughout my sentient life in the years leading up to a momentous August–September 2016 stint spent in Rohnert Park, there is no discipline I feared and frantically avoided more than organic chemistry. But your carrot-rather-than-stick approach of lab time intercalated with lunchtime walks to the casino for the BBQ salads that Will and I still speak of in reverent

(i.e., fanatical) tones truly worked wonders—and with trusty biomarker “bible” (Peters and

Moldowan, 1993; Peters, Walters, and Moldowan, 2005) in hand and your ever-patient and inspiring mentorship, I gained a “license to distill [liquid mixtures via rotary evaporation and valuable information on biodegradation, thermal maturity, and paleodepositional environments via interpretation of molecular fossils].” Thanks, Mike. P.S.: I am highly amenable to being lured back to Sonoma County with even a faint promise of said BBQ salad—and, maybe even some lab work here and there—but, I’ll also be hoping for the

xvii appearance of a certain bandana and leather jacket I seem to remember from a Salt Lake

City dancefloor at the 2018 AAPG gathering.

To Allegra Hosford Scheirer: Your tireless championing of our BPSM group—and of all of us students—warrants hundreds of words alone. Similarly, I cannot thank you enough for supporting, alongside Steve, my ofttimes scattered and (probably ofttimes frustratingly) diverse exploration of countless research rabbit holes! But here I especially want to thank you so very deeply for your support as a dear friend—as an ever-energetic and deeply compassionate presence to me during both my professional and personal journey (and—as you know better than most—many profound challenges) throughout my entire time at

Stanford. There are too many heartwarming experiences to relate, as well as some hilarious ones—for instance, I’ll never know how you tolerated the whole “(I) We’re at a highly important BPSM meeting in southern California; (II) Zack is the only registered driver for that rental van; (III) Zack dips his hands in that active oil seep and is immediately inundated in sticky (and I mean sticky) tar; (IV) Zack drives with grocery bags over his hands and affiliates in tow to a Home Depot to get industrial-grade tar ‘n’ grease remover, delaying the expedition by some 75 minutes; (V) No one is pleased except Allegra, who finds the situation rather humorous (and I’m quite sure has [confidential!] photographic evidence of the whole debacle).” Thank you, Allegra.

To Steve “Papa Steve” Graham: At risk of (and in hopes of) sounding cheesy to mask my indescribable gratitude and admiration—thank you for bringing me to Stanford. Thank you for handing me the true Golden Ticket (no loopholes, unlike poor Charlie’s PhD aspirations!), and for providing me with the means, confidence, and complete freedom to explore and develop as a scientist and as a person. To directly quote “Papa Junior” (i.e.,

xviii dear friend Matt Malkowski), who captured much of my sentiment so perfectly: “… I am tremendously grateful for his patience and tolerance with me during this process. Steve has always supported me and encouraged me to become my own scientist, pursue the questions that interest me most, and to do the type of geology that I want to do. It will never cease to amaze me how a man can wear so many hats and wear them all so well. Despite the many hats and the many advisees, Steve always made himself available for me when I asked. He has shown me how to put his students’ careers and success ahead of his own and has been a great role model…” Steve, thank you for not only your absolute tolerance, but also your encouragement as I continue to try on hat upon hat (and usually in an effort to—like you—wear many all at once, but never nearly as well). I am also ever in awe at your uncanny ability to discern the precise moment when a hat is just ever so large that it slips down—and down a bit further—as it threatens to cover some advisee’s eyes as they inadvertently wander precipitously close to some craggy cliffside, before gently turning their shoulders some 90–180° and nudging them ahead, allowing them to continue merrily on their way—but this time into a grassy field abutted by stunning outcrop of undeformed deep-water siliciclastic deposits (metaphorically speaking, of course—part of your unending tolerance and openness was also tolerating the fact that my work herein did not involve analysis of a rock sample or outcrop in any sort of traditional sedimentological or stratigraphic fashion). Thank you for accepting and raising me as part of the family. I am forever indebted for your compassionate, sage mentorship—and will ever look up to you as one of the greatest role models I’ll ever be lucky and privileged enough to know.

Now, while various NDAs (whether explicitly communicated or not) restrict my ability to fully enumerate and/or elaborate upon the vast number of reasons I have to be so incredibly

xix thankful to a good number of incredible friends, I’ll attempt to scratch the surface, and will at times protect myself with occasional, non-exhaustive assignment of some nonbinding

(i.e., non-accusatory) nicknames, either known or fabricated for these purposes.

Under the broad umbrella of GS, BPSM, SPODDS, SedGroup, and the like—“definitely the best 5.25 years of my life so far” would be nothing more than “5.25 pretty busy years of my life so far” without (in no particular order at all) Matt “Dad”/“Malko” Malkowski,

Chris “Ey Meynge”/“Free Food Fall Guy”/“Boxers and Button-Ups” Kremer, Inessa

“Burton Mother” Yurchenko, Zach [sic] Sickmann, Cody “DJ Slurry Drop”/“Elon’s

Biggest Stan” Trigg, Rich “Is the Oven on?”/“He Who Beareth the Domino’s” Stockey,

Colin “cjw2934”/“Lit” White, Xiaowei “Frank”/“Francis” Li, Nilay “Nilaybon”/

“Nilayyyy” Gungor, Anatoly “Twitter King” Aseev, Aaron “Ca[c]t[us] Lady”/“Yellow

Muppet”/“Marvelous Mullet” Steelquist, Best “Van Life”/“What Happened to Your

Ankle?” Chaipornkaew, Will “Why Does He Insist on Walking?”/“Fiddy” Thompson,

Devon “Dancefloor Dominator at—e.g.—The Opal” Orme, Eli “Moy” Hernandez, Tanvi

“Tech Mastermind Moonlighting as Geologist” Chheda, Tom “Loris”/“Jeff, Where Ya

Goin?” Boag, “the Nieminskis” (Jared, Nora, Reyna), Laura “YouTube Co-Star”/“The

Only One Whose Costume is Always Better than Mine” Dafov, Wisam “The Nicest Wizard

I’ve Ever Met” AlKawai, Marisa “Wuthering Heights” Mayer, Drew Laskowski, Mary

“The Real Santa” Reagan, Mustafa Al Ibrahim, Steve “Pioneer Open Mic Wunderkind”

Dobbs, Natalie Sievers, Tyler “CuFeS2” Hall, Earth “The Only Nice One in this

Town”/“Pigeon Point Prodigy” Jaikla, Stephen “Bear Slayer”/“Dill Pickle Chips” Pearcey,

Dan “Paleoclimate” Ibarra, Yao Tong, Tyler “Tyler? He’s in the Maze.” Kukla, Jackson

“That Denim Poker Shirt” Borchardt, Nadja “Hallo Schatz”/“Kebet Stone Conqueror”

Drabon, Anshuman Pradhan, Sam “ShaleLord” Ritzer, Molly Witter, Bradley “Queso

King” Tolar, Jens “Schmetterling” Lund Snee, Will “Grill Guru” Gearty, Pulkit “Point

Count” Singh, Pablo “Squirrel Chef” García del Real, Marcelo Silka, Ginny “Climbing

Buddy”/“Holiday Party Mastermind” Isava, Ziva “Zeevs” Shulaker, Jake “The Ohio

State”/“Noble Gas Geochem” Harrington, Noah Athens, and Marianne

“Cheese”/“Hi.”/“Collaborator” Coholich. Nor would it be so without those who came before me, in a distant and sometimes fuzzy past, including Larisa Masalimova, Sam

Johnstone, Dana Thomas, Joey Nelson, Katlyn Turner, Tess Menotti, Theresa Schwartz,

Matt Thomas, Tom Benson, Jesse Bateman, Nikolaus Deems, Blair Burgreen-Chan, Holly

Young, Keisha Durant, Danica Dralus, Tim O'Brien, Lauren Shumaker, Lauren Schultz,

Meredith Townsend, Brad Ritts, Andrew Hanson, and Jake Covault. Nor, without Stanford

Earth and GS comrades and affiliates thereof, broadly defined—Courtney Payne, Alandra

Lopez, Kevin McCormack, Kirellos Sefein, Dylan Rittman, Abi Vega, Mei Mei, Danielle

Touma, Cecilia Endriga, Sami Chen, Curtis Baden, Arden Wells, Elizabeth Miller, Mathieu

Lapôtre, Alex Kendrick, Cooper Elsworth, Omar Alamoudi, Tristan Ballard, Colette Kelly,

Laura Schaefer, Megan (Dustin) Mahajan, Shersingh Joseph Tumber-Dávila, Hannah Joy-

Warren, Vishal Das, Ian Gottschalk, MC Anderson, Brent Lunghino, Noah Dewar, Taylor

Martin, Josue Fonseca, Kat Gonzales, Krongrath Suwannasri, Jackson MacFarlane,

Alexander Bakay, Elizabeth Johnston, Nicolette Meyer, Bennett Kapili, Erik Sperling, Jon

Payne, Carolyn Lampe, Selena Perrin, EKela Autry, Humberto Arévalo, Liam Bhajan,

Dante Orta, Robert Collar, Greg Ledingham, Malcolm Hodgskiss, Audrey Yau, Elenita

Makani Nicholas, Wen Song, Ethan Williams, Erin Barry, Fantine Huot, and Emily Stoll.

Nor, without those in the general Stanford vicinity who have (in addition to many of those

xxi listed above) been utter anchors to my happiness, well-being, and survival (literally) at many or all points during these past years—Elisa Hofmeister, Stuart

“Stubacca”/“Stululemon”/“Coach”/ “We Shall Subconsciously Replace ‘You’ with ‘Stu’ in Basically Every Song Imaginable, Forevermore”/“Even-More-Hobbies-than-

Nicknames” Farris, Mitch “Meetchellll” Watt, Anna-Katharina von Krauland, Arnav

Mariwala (and the whole Mariwala family!), Grégoire “G”/“Hell on Wheels” Faucher, Ali

“Bioturbation Boss” Cribb, Isabella “Definitely in Charge” Kahhalé (and Mimi and

Momo), Lara “Got Source Rock?”/“Runs Robard’s”/“Take the Bag of Money and Run”

Heister, Emily “Simply Amazing”/“Soil Sorceress” Lacroix and Joe “EDC”/“Fellow

Tavour Addict”/“Bromance” Soultanis, Tamara “Tomorrow”/“Thanks for the Milk”/“Jean

Jacket” Kahhalé, Annie “BOA’d” Mosher, Himanshu Gupta, Igor “Gary”/“Bromance”

Yurchenko and the Yurchenko and Sickmann families at large, Frida “Throw Me Your

Pizza Crust”/“LBD”/“Known (and Feared) by Delivery Persons Across the 94306 Zip

Code” Margarita, “Dr.” James “Canyon Creek”/“LIIT” Cassanelli, Navya Konda, Nathan

“Monkey!”/“Dese Boys Lookin for Pappy!”/“How Did We Possibly Survive All those

Road Trips?” Spielberg, Katie “Lyse that Bacterium!” Bodner, and the Spielberg and

Bodner families, Kait “Trash Queen”/“The Glue That’ll Keep Us All Bound into our mid-

Eighties” Cyr, Zach “The Barefoot Spirit” Stuart, Samir “The Best ‘Za Chef in All the

Land”/“Knows More About Mineral Resources than Most Geologists” Patel, Payton

Hagyard, Lisa “Love” Alpert, Haley “Undefeated Pong Champion” Feck, The Manic

Monologues family and mental health community (a thousand-, thousandfold), the terrific folks of the Stanford Energy Club and planetary science (plus poetry?) communities,

Hursula, Clifford, Los Carnalitos y El Grullense, cabbage salad, SedFarm, the garden, the

xxii chickens (especially Egg, Zack/“the annoying one,” and the Silkies) and squirrels, Rains,

Brown Bag frozen burritos and pizza, unhealthy energy drinks/reused coffee beans, Google

Scholar, Microsoft Word, Ari and House, El Patio, and LinkedIn. Also those Bowdoin cronies, including, in no particular order—Courtney Payne (here’s to always being your #1 bridesmaid), The DFP, Nick “Grey Goose” Cast, RJ “Nasty” Dellecese, Patrick “Bae”

Millet, Justin “President” Pearson, Sean “Baby” Moran, Emery “Brother” Ahoua, Grégoire

“G” Faucher, Bill De La Rosa, and Andy Zheng (you’ll always be my first roommate)— with thanks as well to Steve Cerf and Birgit Tautz for your love and support.

Finally, my family—Mom, Dad, Owen, Jolene, Lando, Alex, Franck, Simone, Grandma

Rita, Grandpa Sef, Grandma Betsy, Grandma Liz, Grandpa Sam, Grandpa John, my dear cousins and aunts and uncles, and the many, many in our respective tribes—I love you, and thank you—now and forever. I wouldn’t be here without you.

As with my undergraduate thesis, I dedicate this dissertation to my brother, Onion.

xxiii

TABLE OF CONTENTS

ABSTRACT ...... iv ACKNOWLEDGEMENTS ...... xii TABLE OF CONTENTS ...... xxiv LIST OF APPENDICES ...... xxviii LIST OF TABLES ...... xxix LIST OF FIGURES ...... xxxi CHAPTER 1: ASSESSING PETROLEUM BIODEGRADATION, SOURCE ROCK THERMAL MATURITY, AND HYDROCARBON MIXING USING SEEP OIL GEOCHEMISTRY, EAST COAST OF NEW ZEALAND ...... 1 ABSTRACT ...... 2 1.1. INTRODUCTION ...... 4 1.2. SAMPLES AND METHODS ...... 6 1.2.1. SAMPLE SELECTION ...... 6 1.2.2. METHODS ...... 6 1.3. RESULTS AND DISCUSSION ...... 9 1.3.1. BIODEGRADATION ...... 9 1.3.2. THERMAL MATURITY ...... 12 1.3.3. IMPLICATIONS FOR PETROLEUM MIXING AND CHARGE HISTORY ...... 18 1.4. CONCLUSIONS ...... 20 ACKNOWLEDGEMENTS ...... 23 REFERENCES ...... 24 CHAPTER 2: INTERPRETING SOURCE ROCK DEPOSITIONAL ENVIRONMENT FROM SEEP OIL ISOTOPE AND ORGANIC GEOCHEMISTRY, EAST COAST OF NEW ZEALAND ...... 45 ABSTRACT ...... 46 2.1. INTRODUCTION ...... 48 2.2. BACKGROUND ...... 50 2.3. MATERIALS AND METHODS ...... 54

xxiv

2.3.1. MATERIALS ...... 54 2.3.2. METHODS ...... 54 2.4. RESULTS AND DISCUSSION ...... 57 2.4.1. SOURCE ROCK ORGANIC MATTER INPUT ...... 57 2.4.2. SOURCE ROCK DEPOSITIONAL ENVIRONMENT OXICITY ...... 60 2.4.3. SOURCE ROCK SEDIMENTARY FACIES ...... 61 2.4.4. SOURCE ROCK AGE ...... 63 2.4.5. IMPLICATIONS ...... 65 2.5. CONCLUSIONS ...... 67 ACKNOWLEDGEMENTS ...... 69 REFERENCES ...... 70 CHAPTER 3: TECTONIC UPLIFT DESTABILIZES SUBSEA GAS HYDRATE: A MODEL EXAMPLE FROM HIKURANGI MARGIN, NEW ZEALAND ...... 92 ABSTRACT ...... 93 3.1. INTRODUCTION ...... 94 3.2. METHODS...... 97 3.3. RESULTS...... 100 3.4. DISCUSSION ...... 102 3.5. CONCLUSIONS ...... 104 ACKNOWLEDGEMENTS ...... 105 REFERENCES ...... 106 CHAPTER 4: GLOBAL INVENTORY OF BIOGENIC GAS HYDRATE-ASSOCIATED SEDIMENT ORGANIC PROPERTIES ...... 125 ABSTRACT ...... 126 4.1. INTRODUCTION ...... 128 4.2. BACKGROUND ...... 130 4.2.1. BIOGENIC GAS GENERATION ...... 130 4.2.2. GAS HYDRATE FORMATION ...... 133 4.2.3. SEDIMENT ORGANIC PROPERTIES, BIOGENIC GAS, AND GAS HYDRATES ...... 135 4.3. METHODS...... 138 4.4. GLOBAL INVENTORY OF SEDIMENT ORGANIC PROPERTIES ...... 142 4.4.1. NORTH AMERICA ...... 142

xxv

4.4.2. SOUTH AMERICA ...... 146 4.4.3. EUROPE ...... 148 4.4.4. AFRICA...... 149 4.4.5. INDIA ...... 149 4.4.6. ASIA ...... 151 4.4.7. OCEANIA ...... 153 4.4.8. ANTARCTICA ...... 154 4.4.9. PERMAFROST-ASSOCIATED GAS HYDRATE ...... 154 4.4.10. SUMMARY OF GLOBAL DRILLING DATA FROM GAS HYDRATE LOCALITIES ...... 157 4.5. CONCLUSIONS ...... 160 ACKNOWLEDGEMENTS ...... 161 REFERENCES ...... 162 CHAPTER 5: GLOBALLY DISTRIBUTED PALEOGENE DEEP-WATER SAND- RICH SYSTEMS DESPITE HIGH SEA LEVEL DURING WARM CLIMATES ...... 219 ABSTRACT ...... 220 5.1. INTRODUCTION ...... 222 5.1.1. THE PARADIGM OF DEEP-SEA SILICICLASTIC DEPOSITION ...... 222 5.1.2. THE EARLY EOCENE HOTHOUSE ...... 223 5.2. METHODS...... 225 5.3. ABUNDANT PALEOGENE HOTHOUSE TURBIDITE SYSTEMS ...... 226 5.4. DISCUSSION ...... 228 5.4.1. THE PARADIGM OF DEEP-SEA SILICICLASTIC DEPOSITION ...... 228 5.4.2. MECHANISMS OF SANDY TURBIDITE DEPOSITION DURING EUSTATIC HIGHS ...... 228 5.4.3. CONCEPTUAL MODEL LINKING GLOBAL CLIMATE AND SAND- RICH TURBIDITE SYSTEMS...... 229 5.4.4. SUGGESTIONS FOR FUTURE RESEARCH ...... 230 5.5. CONCLUSIONS ...... 232 ACKNOWLEDGEMENTS ...... 233 REFERENCES ...... 234 CHAPTER 6: GLOBAL EOCENE-OLIGOCENE UNCONFORMITY DUE TO CLIMATIC COOLING-DRIVEN SUBMARINE EROSION IN SEDIMENTARY BASINS ...... 243

xxvi

ABSTRACT ...... 244 6.1. INTRODUCTION ...... 246 6.1.1. THE EOCENE-OLIGOCENE TRANSITION ...... 246 6.1.2. DEEP-MARINE HIATUSES ...... 248 6.2. METHODS...... 254 6.3. RESULTS...... 256 6.3.1. PASSIVE MARGIN UNCONFORMITIES ...... 256 6.3.2. ACTIVE MARGIN UNCONFORMITIES ...... 257 6.4. DISCUSSION ...... 258 6.4.1. REGIONS LACKING UNCONFORMITIES...... 258 6.4.2. COMPARISON WITH SEISMIC SEQUENCE STRATIGRAPHIC WORK ...... 259 6.5. CONCLUSIONS ...... 262 ACKNOWLEDGEMENTS ...... 263 REFERENCES ...... 264

xxvii

LIST OF APPENDICES

APPENDIX A: SUPPLEMENTARY MATERIAL FOR CHAPTERS 1 AND 2

Appendix A-1. Analytical data: Tables…………………………………………………279

Appendix A-2. Analytical data: Chromatograms.……………………………………...290

APPENDIX B: SUPPLEMENTARY MATERIAL FOR CHAPTER 2

Appendix B-1: Correction to International Journal of Earth Sciences publication…...300

APPENDIX C: SUPPLEMENTARY MATERIAL FOR CHAPTER 5

Appendix C-1. Detailed methodology………………………………………………….302

Appendix C-2. Detailed description of turbidite occurrences…………………………..305

Appendix C-3. References cited………………………………………………………..382

APPENDIX D: SUPPLEMENTARY MATERIAL FOR CHAPTER 6

Appendix D-1. Detailed description of unconformities………………………………...438

Appendix D-2. References cited………………………………………………………..479

xxviii

LIST OF TABLES

CHAPTER 1

Table 1.1. Diasterane and sterane peak assignment for Figure 1.3……………………….33

Table 1.2. Sterane and diasterane %C27−29 for all samples……………………………….34

Table 1.3. C31−35 17α-hopane 22S/(22S + 22R) isomerization ratios…………………….35

Table 1.4. Extent of cracking for the mixed-source samples……………………………..36

CHAPTER 2

Table 2.1. Stable carbon isotope and canonical variable values and interpretations……...78

Table 2.2. C30 sterane indices from GC-MS-MS…………………………………………79

Table 2.3. Regular steranes, diasteranes, and diasterane ratio from GC-MS-MS………...80

CHAPTER 3

Table 3.1. Predicted gas hydrate stability zone depth and thickness through time...……118

CHAPTER 4

Table 4.1. North American gas hydrate sites……………………………………………190

Table 4.2. South American gas hydrate sites…………………………………………..197 xxix

Table 4.3. European gas hydrate sites………………………….………………………200

Table 4.4. Indian gas hydrate sites…………………….……………………………….201

Table 4.5. Asian gas hydrate sites………………….…………………………………..202

Table 4.6. Australian gas hydrate sites…………………………………………………205

xxx

LIST OF FIGURES

CHAPTER 1

Figure 1.1. Oil seep locations and general setting in New Zealand.……………………...37

Figure 1.2. Whole-oil (GC-FID) gas chromatograms.…………………………………...38

Figure 1.3. m/z 217 fragmentograms illustrating sample sterane distributions.………….39

Figure 1.4. C27, C28, and C29 regular steranes vs. βα-diasteranes…..……………………..40

Figure 1.5. m/z 191 fragmentograms illustrating sample hopane distributions…………..41

Figure 1.6. Investigation of C29 25-Nor-17α(H)-Hopane………………………………...42

Figure 1.7. Various biomarker maturity parameters calculated from GC-MS-MS………43

Figure 1.8. Methyldiamantane vs. stigmastane concentrations…………………………..44

CHAPTER 2

Figure 2.1. Oil seep locations and general setting in New Zealand.……………………...81

Figure 2.2. Chronostratigraphic chart and regional tectonic history……………………..82

13 13 Figure 2.3. δ Csaturates versus δ Caromatics.………………………………………………..83

Figure 2.4. Ternary diagram of C27, C28, and C29 sterane abundances……………………84

Figure 2.5. Plot of wt. % sulfur versus oleanane index…………………………………...85

Figure 2.6. Tetracyclic polyprenoid ratios vs. %C30 diasteranes…………………………86 xxxi

Figure 2.7. Wt. % sulfur versus C35 homohopane index………………………………….87

Figure 2.8. Gammacerane index versus C35/C34 homohopanes………………………….88

Figure 2.9. Various biomarker parameters indicative of sedimentary facies……………..89

13 13 Figure 2.10. δ Csaturates versus δ Caromatics for various New Zealand oils………………..90

Figure 2.11. % C26 24/(24+27) Nordiacholestanes versus oleanane index……………….91

CHAPTER 3

Figure 3.1. Location of modeled transect along New Zealand’s Hikurangi margin...…..119

Figure 3.2. Seismic line and corresponding 2D model of gas hydrate stability zone……120

Figure 3.3. Modeled time steps (24–0 Ma) with predicted hydrate stability zone………121

Figure 3.4. Modeled water depths and corresponding stability zone thicknesses……….123

Figure 3.5. Seismic- vs. modeling-based prediction of hydrate stability zone extent…...124

CHAPTER 4

Figure 4.1. Global gas hydrate sites from which TOC and HI data were compiled……..206

Figure 4.2. North American sites from which TOC and HI data were compiled………..207

Figure 4.3. South American sites from which TOC and HI data were compiled………..208

Figure 4.4. European sites from which TOC data were compiled………………………209

xxxii

Figure 4.5. Indian sites from which TOC data were compiled………………………….210

Figure 4.6. Asian sites from which TOC and HI data were compiled…………………...211

Figure 4.7. New Zealand site from which TOC data were compiled……………………212

Figure 4.8. Permafrost-associated gas hydrate sites…………………………………….213

Figure 4.9. Distribution of TOC values of gas hydrate-associated sediment……………214

Figure 4.10. Distribution of HI values of gas hydrate-associated sediment…………….215

Figure 4.11. Box plot of TOC values from gas hydrate-associated sediment…………...216

Figure 4.12. Box plot of HI values from gas hydrate-associated sediment……………...217

Figure 4.13. Box plots of TOC and HI of sediment below, within, above the GHSZ…..218

CHAPTER 5

Figure 5.1. Sea-level curves and counts of early Paleogene sandy systems…………….239

Figure 5.2. Paleocene, early Eocene, and middle Eocene locations examined………….240

Figure 5.3. Conceptual model linking global climate and sandy deep-sea systems……..241

Figure 5.4. Schematic of sandy deep-sea deposition under falling vs. high sea level…...242

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CHAPTER 6

Figure 6.1. Cenozoic sea-level curves………………………………………………….273

Figure 6.2. Eocene-Oligocene unconformities documented herein.……………………274

Figure 6.3. Locations Vail et al. (1977) used to reconstruct Phanerozoic sea-level……..275

xxxiv

CHAPTER 1

ASSESSING PETROLEUM BIODEGRADATION, SOURCE ROCK THERMAL

MATURITY, AND HYDROCARBON MIXING USING SEEP OIL

GEOCHEMISTRY, EAST COAST OF NEW ZEALAND

ASSESSING PETROLEUM BIODEGRADATION, SOURCE ROCK THERMAL

MATURITY, AND HYDROCARBON MIXING USING SEEP OIL

GEOCHEMISTRY, EAST COAST OF NEW ZEALAND

Zachary F. M. Burton1, J. Michael Moldowan1,2, Richard Sykes3, and Stephan A.

Graham1

1Department of Geological Sciences, Stanford University, Stanford, CA, USA

2Biomarker Technologies Incorporated, Rohnert Park, CA, USA

3GNS Science, Lower Hutt, New Zealand

ABSTRACT

Determining oil quality is essential to identifying valuable resource accumulations.

However, in new areas of exploration, little information is available on the processes affecting resource quality. Geochemical analyses of oil seeps from frontier regions of New

Zealand’s east coast illustrate an application of underutilized resource quality assessment techniques. Distributions of n-alkanes and isoprenoids reveal biodegradation, and thus potentially lower oil quality in the “southern” versus the “northern” oil seeps. However, sterane and terpane compounds are unaltered, indicating overall biodegradation of these oils is low to moderate. Additionally, lack of 25-norhopane indicates degradation of southern oils may be solely aerobic. Therefore, any subsurface accumulations are potentially unaffected. Investigation of sterane and hopane isomerization ratios and

2 additional sterane and terpane maturity parameters is paired with diamondoid analyses of oil-to-gas conversion and petroleum mixing. Three distinct petroleum mixtures are identified among the sampled seeps: 1) a seep composed of an early/peak oil window component and an intensely cracked condensate/wet gas component, 2) seeps solely containing a peak/late oil window component, and 3) seeps composed of a peak/late oil window component and an intensely cracked condensate/wet gas component. Identified components indicate at least three distinct charges or stages of petroleum generation. Black oil components might indicate actively producing source rock in all regions represented by the seeps. Intensely cracked components indicate petroleum mixing via thermogenic gas infiltration, and suggest an effect on oil quality. Important questions concerning migration pathways and timing, ties to New Zealand’s offshore basins, and potential for reservoir entrapment of these petroleum components remain.

1.1. INTRODUCTION

The key goal of the oil and gas industry’s upstream sector is to find and extract significant amounts of energy resources. Because there are high risks associated with this process, methods for identifying economically-attractive accumulations are critical. Geochemical analysis of oils is one technique available for assessing both the presence and producibility of petroleum accumulations (Hunt, 1979; Tissot and Welte, 1984).

Analytical geochemistry provides inexpensive methods to leverage the limited data available in frontier exploration areas. Geochemical technology can be successfully applied to determine impact of biodegradation on oil quality and to identify petroleum mixtures

(Seifert and Moldowan, 1979; Connan, 1984; Volkman et al., 1984; Peters and Moldowan,

1993). It can be used to assess contribution of shallow versus deep sources and to fingerprint conversion of oil to gas (Jones and Drozd, 1983; Behar et al., 1992; Dahl et al.,

1999). Analysis of geochemical data in conjunction with geological and geophysical data is critical for success in frontier exploration and production (Demaison, 1984).

Oil seeps in areas of New Zealand’s east coast are an ideal target for demonstrating the power and utility of advanced geochemical techniques. This region contains hundreds of seeps, shows, and other petroleum indicators, and has attracted exploration interest for over a century (Uruski, 2010). However, petroleum systems are still poorly understood, and significant accumulations of oil and gas remain undiscovered (Uruski, 2010). The oils examined as a part of this study have been previously characterized as marine oils, however, the source rock units that generated these oils remain uncertain (Johnston et al.,

1992; Murray et al., 1994; Rogers et al., 1994, 1999; Killops et al., 1996).

The east coast region was characterized by Gondwanan subduction lasting until ~85 Ma

(Uruski, 2010). This period of subduction gave way to a prolonged period of passive margin sedimentation, lasting ~85 to 25 Ma (Uruski, 2010). It was during this period of passive margin sedimentation that the inferred source rock intervals were deposited in the area of interest, including within, it is inferred, the East Coast and Pegasus basins (Uruski,

2010). The Upper Cretaceous through Paleocene Whangai Formation and the Paleocene age Waipawa Formation black shale are the two most likely source rock candidates within the region, although, as mentioned above, it is uncertain which of these source rock candidates may be contributing to the oil seeps found up and down the east coast (Uruski,

2010; Johnston et al., 1992; Murray et al., 1994; Rogers et al., 1994, 1999; Killops et al.,

1996).

We present results from our studies of onshore oil seep samples from New Zealand’s North

Island and South Island (Figure 1.1). We studied biomarker and diamondoid compounds to determine the degree of petroleum biodegradation, to assess thermal maturity, and to calculate the extent of cracking of these oils. We bring the results of these studies together to identify mixing of different charges within these oils.

1.2. SAMPLES AND METHODS

1.2.1. SAMPLE SELECTION

Oil seeps provide key insight into petroleum system characteristics and can greatly reduce exploration risk by establishing active source rock presence, by constraining maturity and biodegradation, and by identifying mixed sources (Seifert and Moldowan, 1978; Magoon and Dow, 1994). For this study, we selected five oil seep samples provided by GNS Science

(Sykes et al., 2012). The three “northern” oil samples are from the Rotokautuku, Totangi, and Waitangi oil seeps, located most proximal to New Zealand’s East Coast and

Raukumara basins (Figure 1.1). The “southern” oils are from the Isolation Creek and

Kaikoura seeps, located closest to New Zealand’s deepwater Pegasus Basin (Figure 1.1).

1.2.2. METHODS

1.2.2.1. GC-FID — Whole-oil analysis

Whole-oil chromatograms for the five oil seep samples were provided by GNS Science from Sykes et al. (2012). These chromatograms were obtained via gas chromatography- flame ionization detection (GC-FID) at Applied Petroleum Technology (APT) in Kjeller,

Norway using procedures detailed in Weiss et al. (2000).

1.2.2.2. Liquid chromatography — Sample preparation

Crude oil samples (30-40 mg) were weighed and spiked with 5β-cholane and deuterated diamondoid internal standards for quantitation. Deuterated diamondoids in the spike include D15-1-methyladamantane, D16-adamantane, D3-1-methyldiamantane, D4- diamantane, and D4-triamantane.

Spiked samples were fractionated by sequential elution using a proprietary light hydrocarbon solvent and dichloromethane on silica gel columns to obtain saturate and aromatic fractions. Paraffins were removed from saturated hydrocarbon fractions using a proprietary light hydrocarbon solvent on zeolite columns.

1.2.2.3. GC-MS — Saturated hydrocarbon fraction analysis

Saturated fractions were analyzed by gas chromatography-mass spectrometry (GC-MS) using an Agilent Technologies 7890A/5975C GC-MS system with a 7693A autosampler.

The GC was equipped with a 60 m DB-1 column (0.25 mm i.d.; 0.25 μm phase thickness).

Helium was used as carrier gas at a flow rate of 2.25 mL/min. The temperature program began at 80°C and was raised at a rate of 2°C/min to 320°C where it was held for 15 min.

Calibrated external standards were run alongside the oil samples to account for response factors of different compounds.

Biomarker analysis and quantitative diamondoid analysis was conducted by monitoring ions with specific mass-to-charge ratios (m/z). Biomarkers were quantified using m/z 191 for hopanes and m/z 217, 218, and 259 for steranes. Diamondoids were analyzed and quantified using m/z 135, 136, 149, 150, and 152 for adamantanes, m/z 187, 188, 192, and

201 for diamantanes, m/z 239, 240, and 244 for triamantanes, and m/z 292 for tetramantanes.

Quantitation of biomarker and diamondoid compounds was achieved by integration of compound peak areas and/or measurement of peak heights relative to internal and external standards.

1.2.2.4. GC-MS-MS — Saturated hydrocarbon fraction analysis

Gas chromatography-tandem mass spectrometry (GC-MS-MS) of the saturated hydrocarbon fractions was conducted using an Agilent Technologies 7890A GC interfaced to a 7000 Triple Quad GC-MS with a 7693A autosampler. The GC was equipped with a

60 m DB-1 column (0.25 mm i.d.; 0.25 μm phase thickness). Helium was used as carrier gas at a flow rate of of 2.25 mL/min. The temperature program began at 80°C and was raised at a rate of 2°C/min to 320°C where it was held for 15 min. Oil samples were analyzed alongside calibrated external standards in order to account for variability in chemical compound response differences.

The mass spectrometer was run in parent-to-daughter mode using the first quadruple to focus parent ions and the third quadrupole to focus daughter ions. The transitions monitored were m/z 370→177, 318→191, 330→191, 332→191, 370→191, 384→191,

398→191, 412→191, 426→191, 440→191, 454→191, 468→191, 482→191, and

426→205 for hopane biomarkers and 288→217, 302→217, 330→217, 358→217,

372→217, 386→217, 400→217, and 414→217 for sterane biomarkers.

Biomarkers were quantified by integrating compound peak areas and/or measuring peak heights relative to internal and external standards.

1.3. RESULTS AND DISCUSSION

1.3.1. BIODEGRADATION

Assessment of biodegradation is critical in exploration and production. Biodegradation can lead to removal of the most economically-desirable petroleum compounds and can significantly affect the value of a petroleum accumulation (Connan, 1984; Wenger et al.,

2002). Biodegradation leads to lower API gravity (i.e., higher viscosity) and lower producibility (i.e., increased difficulty of extraction) of petroleum accumulations (Evans et al., 1971; Head et al., 2003).

1.3.1.1. n-Alkanes and isoprenoids

Whole-oil chromatograms reveal key differences between samples (Figure 1.2). Gas chromatograms for the Totangi, Waitangi, and Rotokautuku oil seep samples show good preservation of both n-alkane and isoprenoid compounds (Figure 1.2). In contrast, Isolation

Creek and Kaikoura samples show almost no preservation of n-alkanes and minimal preservation of isoprenoids (Figure 1.2).

The variations in compound class preservation between 1) Totangi, Waitangi, and

Rotokautuku crude oil samples and 2) the Isolation Creek and Kaikoura oil samples highlight distinctions in the degree of biodegradation to which each group of samples has been subjected. The n-alkanes and isoprenoids are the compound classes most susceptible to destruction via biodegradation (Bailey et al., 1973; Atlas, 1981; Leahy and Colwell,

1990). Therefore, the high degree of both n-alkane and isoprenoid compound preservation in the Totangi, Waitangi, and Rotokautuku oil seep samples suggests no more than slight amounts of biodegradation. In contrast, the poor preservation of n-alkanes and isoprenoids

9 in the Isolation Creek and Kaikoura oil samples suggests greater, although still moderate, biodegradation of these samples (Wenger and Isaksen, 2002). A more pronounced unresolved complex mixture of hydrocarbons in the Isolation Creek and Kaikoura oils also suggests greater biodegradation of the n-alkane and isoprenoid compounds (Blumer et al.,

1973; Gough and Rowland, 1990; Killops and Al-Juboori, 1990).

1.3.1.2. Steranes

Despite the contrast in the degree of biodegradation identified via assessment of n-alkane and isoprenoid compounds (Figure 1.2), the GC-MS sterane (m/z 217) distributions of all five oil samples show clean peaks without identifiable evidence for degradation of steranes in any of the samples (Figure 1.3; for peak assignment see Table 1.1). In all samples, both regular steranes and rearranged steranes (diasteranes) are well-preserved. After n-alkanes and isoprenoids, steranes are typically the compounds most susceptible to biodegradation and can be completely removed under heavy biodegradation, whereas diasteranes are much more resistant to destruction (Seifert and Moldowan, 1979; Peters and Moldowan, 1993).

Comparison of sterane and diasterane fingerprints shows that sterane peaks in the Kaikoura oil sample are just as well-preserved relative to diasterane peaks as the steranes in the

Waitangi sample (Figure 1.3; for peak assignment see Table 1.1). This suggests absence or only minimal amounts of sterane biodegradation in all samples and provides a contrast to results from examination of n-alkane and isoprenoid biodegradation discussed above.

Sterane and diasterane concentrations obtained via GC-MS-MS further support the finding of negligible biodegradation of steranes in any of the crude oil samples. Ratios of total C27,

C28, and C29 steranes to total C27 + C28 + C29 steranes are nearly identical to respective ratios of total C27, C28, and C29 βα-diasteranes to total C27 + C28 + C29 βα-diasteranes (i.e., C27 10 sterane / total C27 + C28 + C29 steranes is nearly identical to C27 βα-diasteranes / total C27 +

C28 + C29 βα-diasteranes) (Table 1.2) for all samples except Rotokautuku, which shows a lower C27 sterane ratio versus C27 diasterane ratio, possibly due to thermal maturity effects or the relative activity of clay (van Kaam-Peters et al., 1998). Comparison of ternary diagrams clearly illustrates the corroboration of sterane and diasterane data (Figure 1.4).

This strongly supports the finding that steranes have not been altered by biodegradation and also suggests that maturity is not playing a role in affecting sterane distributions, with the possible exception of slight differences in the C27 steranes of the Rotokautuku oil sample (Seifert and Moldowan, 1979).

1.3.1.3. Hopanes

Hopanes are often more robust to biodegradation than steranes, however, in some instances hopanes may be removed before steranes (Rullkötter and Wendisch, 1982; Peters and

Moldowan, 1991). GC-MS m/z 191 hopane distributions of the five oil samples show no indication of hopane destruction (Figure 1.5). As with sterane distributions, the hopane peaks are qualitatively clean. Detailed examination of the C31-C35 17α(H),21β(H)-

Homohopanes reveals good preservation of compound peaks in all samples (Figure 1.5), and in particular, good preservation of the C35 homohopane peaks, which in some cases are more susceptible to biodegradation (Goodwin et al., 1981; Munoz et al., 1997). The lack of destruction of hopane compounds, as with the lack of destruction of steranes, suggests relatively low degrees of biodegradation in these oil samples.

Biodegradation of hopanes often leads to observation of 25-norhopanes in abundances corresponding to the severity of biodegradation (Blanc and Connan, 1992; Moldowan and

McCaffrey, 1995; Bennett et al., 2006). Measurement of the m/z 398→191 transition 11 reveals the absence of 25-norhopanes in all of these oils (Figure 1.6). This provides further indication of little to no biodegradation of hopane compounds.

1.3.1.4. Discussion of biodegradation

Preservation of steranes and hopanes suggests relatively low levels of biodegradation in these oils, or, at least suggests that the seep samples contain a component of relatively non- biodegraded oil. The distributions of n-alkanes and isoprenoids in the northern samples

(Rotokautuku, Totangi, and Waitangi seep samples) indicate good preservation of these compound classes, and is likely indicative of essentially fresh seep oil for these particular samples i.e., active oil seeps (Peters and Moldowan, 1993). In contrast, the extensive destruction of n-alkanes and isoprenoids seen in the southern oils (Isolation Creek and

Kaikoura samples) suggests greater biodegradation of these compounds. Biodegradation of these two southern oil samples might be related to aerobic microbial activity (Atlas and

Bartha, 1992; Damsté et al., 2002; Hoefs et al., 2002). Aerobic degradation is more common in less active oil seeps i.e., seeps with less frequent fresh oil input. The interpretation of subaerial biodegradation in the southern oils is supported by absence of

25-norhopanes in all oil samples. 25-norhopane is generally associated with anaerobic degradation of petroleum in subsurface reservoirs and is therefore not an expected product of aerobic degradation of subaerial seeps (Bost et al., 2001).

1.3.2. THERMAL MATURITY

Thermal maturity of oils directly affects oil quality. Increasing maturity leads to increasing

API gravities and changes in other bulk oil properties (Orr, 1974). However, traditional bulk properties such as API gravity, sulfur, and saturated hydrocarbon content are not only

12 affected by maturity but also by differences in source-rock facies, migration, and reservoir characteristics (Seifert and Moldowan, 1978; Peters and Moldowan, 1991). Therefore, examination of bulk properties alone tends to yield a less robust measure of maturity

(Peters et al., 2005).

Sterane and hopane biomarker compounds provide a much more robust means of assessing thermal maturity of oils. Ratios of various biomarker compounds can be utilized to bracket the maturity of liquid hydrocarbon samples in the early, peak, and late oil windows, and into the condensate/wet gas stage (van Graas, 1990; Peters et al., 2005).

Concentrations of diamondoid compounds are especially effective in assessing the maturity of late oil window and condensate/wet gas window hydrocarbons (Dahl et al., 1999; Zhang et al., 2005; Wei et al., 2007; Azevedo et al., 2008). Diamondoid analysis can indicate the contribution of a condensate/wet gas-generating source to both determine thermal maturity as well as fingerprint previously unidentified components of petroleum mixtures (Dahl et al., 1999; Wang et al., 2006; Springer et al., 2010).

1.3.2.1. Biomarker parameters

Biomarker isomerization ratios are some of the most commonly used and most robust parameters for indicating thermal maturity (Mackenzie et al., 1980; Mackenzie and

McKenzie, 1983; Seifert and Moldowan, 1986). It should be noted that biomarker maturity proxies are only effective in the main phase of the oil window. Maturity of late- or post- mature oils will not be characterized by biomarkers. Diamondoids show clearly that some of these oils are high-low maturity mixtures (see ahead).

Due to C31–C35 hopane equilibration as early as equivalent vitrinite reflectance of ~0.5%, isomerization ratios of C31–C35 17α-hopanes are particularly useful in assessing immature to early oil generation (Schoell et al., 1983). Hopane 22S/(22S+22R) ratios of ~0.6 indicate equilibration in the early oil window (Seifert and Moldowan, 1980), and do not increase from this ~0.6 endpoint value as thermal maturity increases beyond the early oil window.

Each of the C31, C32, C33, C34, and C35 22S/(22S+22R) ratios for all oil samples examined in this study have reached the ~0.6 endpoint value (Table 1.3). This suggests equilibration in these oil samples and indicates maturities at least as high as the early oil window.

Both 20S/(20S+20R) and ββ/(ββ+αα) isomerization ratios of the C29 steranes increase with increasing thermal maturity (Seifert and Moldowan, 1986), and are effective in assessing maturities ranging from immature to peak oil window maturity (Peters et al., 2005). These

C29 isomerization ratios suggest similar maturities for the Totangi, Waitangi, Isolation

Creek, and Kaikoura oils, whereas both isomerization parameters suggest a somewhat lower maturity in the Rotokautuku oil (Figure 1.7). The 20S/(20S+20R) isomerization ratios have all reached equilibrium in the Totangi, Waitangi, Isolation Creek, and Kaikoura oils, but none of these oils show equilibration of the ββ/(ββ+αα) isomerization ratios

(Figure 1.7). This suggests early to peak oil window generation of at least some component of these oils. In contrast, the Rotokautuku oil has reached neither isomerization endpoint, indicating a lower maturity pulse of oil generated in the early oil window. Slightly higher

ββ/(ββ+αα) isomerization ratios in the Isolation Creek and Kaikoura oils may suggest slightly higher thermal maturity than in the Totangi and Waitangi oils.

In addition to biomarker isomerization ratios, other sterane and hopane compound parameters such as trisnorhopane and norcholestane ratios can provide corroborating information on thermal maturity (Peters et al., 2005).

The ratio of C27 18α-trisnorneohopane to C27 17α-trisnorhopane, here expressed as

Ts/(Ts+Tm), shows a strong dependence on maturity and can be used to assess maturities ranging from immature into the condensate/wet gas window (Seifert and Moldowan, 1978;

Kolaczkowska et al., 1990). Because Ts/Tm ratios can be highly influenced by source-rock facies, these ratios are best interpreted alongside other biomarker parameters (Moldowan et al., 1986). Maturities indicated by the Ts/(Ts+Tm) hopane ratios of these oils show agreement with the maturities indicated by sterane isomerization ratios (Figure 1.7).

Isolation Creek and Kaikoura oils appear to be at least slightly more mature than Totangi and Waitangi oils, and the Rotokautuku oil is less mature than all other oils (Figure 1.7).

Relative concentrations of C26 steranes can also serve as maturity parameters. The abundance of C26 21-norcholestane relative to the total amount of 21-, 24-, and 27- norcholestanes increases with increasing thermal maturity (Moldowan et al., 1991), although source rock lithology can also have an impact on this ratio. The ratio of 21- norcholestane relative to total norcholestanes is more than twice as high in the Isolation

Creek and Kaikoura oils as it is in the Totangi, Waitangi, and Rotokautuku oils (Figure

1.7). This suggests a higher thermal maturity in the Isolation Creek and Kaikoura oils. This finding is consistent with sterane ββ/(ββ+αα) isomerization ratios and with Ts/(Ts+Tm) hopane ratios. The 21-norcholestane ratios do not suggest significant difference in the maturities of the Totangi, Waitangi, and Rotokautuku oils (Figure 1.7).

1.3.2.2 Diamondoid parameters

Analysis of diamondoid concentrations in conjunction with biomarker concentrations allows for both assessment of the degree of oil-to-gas cracking and direct identification of mixed oils (Dahl et al., 1999). Utilizing a high diamondoid baseline of 15 ppm (Moldowan et al., 2015), the southern oils (Isolation Creek and Kaikoura samples) are shown to have extremely high concentrations of 3- + 4-methyldiamantanes at 202 ppm and 163 ppm, respectively. The Rotokautuku oil sample has a high concentration of 36 ppm, while the

Totangi and Waitangi oil samples have lower concentrations (i.e., below the baseline of 15 ppm) of 13 ppm and 9 ppm, respectively. The high diamondoid concentrations of the

Isolation Creek, Kaikoura, and Rotokautuku samples indicate intensely cracked oils

(Figure 1.8), whereas the low concentrations of the Totangi and Waitangi oils likely represent end-member uncracked oil values. Figure 1.8 shows the 3- + 4-methyldiamantane concentrations plotted versus a biomarker concentration, C29-sterane stigmastane

(5α,14α,17α(H),20R-24-ethylcholestane). Stigmastane decreases with increasing maturity and is effectively completely destroyed by the time oil begins to crack. Therefore, any appreciable amount of stigmastane in a sample containing high amounts of diamondoids indicates a mixed oil (Dahl et al., 1999; Sassen and Post, 2008; Moldowan et al., 2015).

The Isolation Creek and Kaikoura oil samples have lower concentrations of stigmastane than the northern oils; however, both oils still contain over 10 ppm stigmastane, indicating a widely mixed maturity range (Dahl et al., 1999). The pairing of both high biomarker

(stigmastane) concentrations and high diamondoid (3- + 4-methyldiamantane) concentrations strongly indicates that Isolation Creek, Kaikoura, and Rotokautuku all represent a mixture of black oil (high biomarkers, negligible diamondoids) and cracked oil

(negligible biomarkers, high diamondoids). The ubiquity of well-preserved sterane and hopane biomarkers, discussed in the sections above, similarly supports the finding that these three oil samples represent mixtures. This is because if samples solely represented unmixed, thermally-cracked hydrocarbons, the sterane and hopane biomarkers observed here would not be preserved.

The extent of cracking of the Rotokautuku, Isolation Creek, and Kaikoura samples can be quantitatively determined using the equation %Cracking = [1-(Co/Cc)] X 100, whereby Co represents the methyldiamantane concentration in the end-member uncracked oil samples and Cc is the methyldiamantane concentration of any cracked sample (Dahl et al., 1999). It should be noted that this assumes that any cracked sample has been derived from the same source rock and source facies. It is also possible that, due to potential evaporative loss of light ends, calculation of %Cracking can cause an overestimate of the degree to which an oil sample has been cracked (Dahl et al., 1999; Wei et al., 2007). To account for these potential uncertainties, we calculate the percentage of cracking based on a Co (diamondoid baseline) of 4 ppm, or about the global average (Moldowan et al., 2015), and based on a very high (i.e., quite conservative) Co of 10 ppm. Even utilizing the less-realistic conservative baseline value of 10 ppm, both Isolation Creek and Kaikoura samples show over 94% conversion of liquid hydrocarbons to gas and the Rotokautuku sample shows over 70% cracking (Table 1.4). Using the more realistic diamondoid baseline value of 4 ppm, Isolation Creek and Kaikoura both show about 98% cracking and Rotokautuku shows nearly 90% cracking (Table 1.4). Destruction of oil via biodegradation may account for some elevation of diamondoid concentrations, however, biodegradation does not realistically explain 70+% cracking in these oils. Larter et al. (2005) demonstrated at most

60% cracking at biodegradation levels of PM 8 based on the Peters & Moldowan (1993) biodegradation scale. A biodegradation level of PM 8 is quite severe, and would surely lead to alteration of sterane and hopane parameters, which we see no evidence for.

Therefore, although biodegradation may account for some elevation of diamondoid concentrations, it cannot fully account for the high concentrations of diamondoids observed here.

Diamondoid results indicate mixed oils containing very intensely cracked components in the southern (Isolation Creek and Kaikoura) oil samples and in the Rotokautuku oil seep sample, while the Totangi and Waitangi oil samples represent uncracked (end-member) oils.

1.3.3. IMPLICATIONS FOR PETROLEUM MIXING AND CHARGE HISTORY

Hopane and sterane biomarker maturity parameters reveal a contrast between more mature

Totangi, Waitangi, Isolation Creek, and Kaikoura black oil components (i.e., components rich in heavy, large, non-volatile hydrocarbons) and a less mature Rotokautuku black oil component. The lower maturity of the Rotokautuku oil indicates that this oil was generated by a source rock at an earlier oil window maturity than the other four oils. This suggests presence of a lower-maturity black oil charge in the Rotokautuku oil versus higher-maturity black oil charges in the Totangi, Waitangi, Isolation Creek, and Kaikoura oils.

Diamondoid concentrations indicate high-maturity condensate/wet gas window components in the Isolation Creek, Kaikoura, and Rotokautuku oils, whereas the Totangi and Waitangi oils are missing the high-maturity charge.

Both biomarker maturity parameters and diamondoid cracking parameters reveal three distinct charge histories represented by the five oil samples studied here. The Rotokautuku oil contains evidence for an early to peak oil window charge as well as a condensate/wet gas charge, the Totangi and Waitangi oils reveal evidence for just a peak to late oil window charge, and the Isolation Creek and Kaikoura oils indicate the presence of a peak to late oil window charge and a condensate/wet gas charge.

1.4. CONCLUSIONS

We have investigated both biomarker compounds and diamondoid compounds to assess oil seep biodegradation, maturity, and oil-to-gas cracking. We used these results to understand petroleum mixing and charge history for petroleum systems in this area.

1. Biodegradation. The northern oils (Rotokautuku, Totangi, and Waitangi) show preservation of n-alkane and isoprenoids, suggesting very slight biodegradation. This is uncommon in subaerial oil seeps, and likely indicates highly active seeps. In contrast, the southern oils (Isolation Creek and Kaikoura) show extensive destruction of these compounds, indicating heavy biodegradation and perhaps less active seeps.

All oils show good preservation of sterane and hopane compounds and show absence of

25-norhopane. A reasonable explanation for the destruction of n-alkanes and isoprenoids in the southern oils could be aerobic degradation of the oils once they were exposed subaerially. This is further supported by the lack of 25-norhopane, which tends to be generated through subsurface anaerobic degradation (e.g., in-reservoir biodegradation).

2. Maturity. Both sterane and hopane maturity parameters identify a component of relatively high-maturity (peak to late oil window) black oil in the Totangi, Waitangi,

Isolation Creek, and Kaikoura oil seeps. In contrast, these parameters suggest a lower- maturity (early to peak oil window) black oil component in the Rotokautuku oil seep.

3. Oil-to-gas cracking. Diamondoid compounds indicate the contribution of intensely cracked components to the Isolation Creek, Kaikoura, and Rotokautuku oil seeps. The conflicting maturities suggested by the coexistence of the lower-maturity component

(identified via biomarker assessment) and the highly cracked component (identified via

20 diamondoid assessment) indicate multiple charges to the Rotokautuku oil. Furthermore,

Isolation Creek, Kaikoura, and Rotokautuku oils all indicate multiple charges due to the coexistence of well-preserved steranes and hopanes with very high diamondoid concentrations. In other words, if any of these oils were solely derived from a charge of intensely cracked oil, sterane and hopane biomarkers would not be present.

4. Implications for oil mixing. The northern Totangi and Waitangi oil seeps contain no evidence of a deep thermal source, and likely represent unmixed (end-member) black oils.

The northern Rotokautuku oil and the southern Isolation Creek and Kaikoura oils are mixed oils containing evidence for at least two charges: one of black oil, and one of intensely cracked hydrocarbons. The identification of this cracked component suggests mixing via gas infiltration from a deep source, and impacts prospectivity of these mixed oils relative to the uncracked Totangi and Waitangi oils.

5. Implications for depth of different charges. Maturity parameters reveal that the end- member Totangi and Waitangi oils and the black oil components of the mixed Isolation

Creek and Kaikoura oils were derived from higher-maturity charges generated in the peak to late oil window. In contrast, the black oil component of the Rotokautuku oil was derived from a lower-maturity charge generated in the early to peak oil window. This suggests that the black oil charge of the Rotokautuku seep likely originated from a shallower depth than that of all other oils.

6. Importance of utilizing advanced geochemical technologies. This study emphasizes the power of applying advanced geochemical technologies and, in particular, diamondoid analysis, in improving understanding of petroleum systems. The high maturity contribution

21 identified in the mixed Isolation Creek, Kaikoura, and Rotokautuku oils would have been overlooked if only conventional biomarker parameters had been investigated.

7. Future work. The findings of this work raise a number of important questions to be addressed through the assessment of petroleum migration pathways in New Zealand’s east coast frontier regions. The prevalence of faulting in the regions where the studied oil seeps are found implies some degree of structural control on the migration of hydrocarbons.

However, as of yet unexamined are questions related to whether these oil seeps originate directly from active source rocks or from leaking reservoirs. Also unexamined would be whether the multiple charges contributing to some seeps represent charges from the same source-rock organofacies or whether they represent charges from different organofacies.

Assessment of the relative volumetric contribution of deep cracked sources versus normal oil window sources to any existing petroleum accumulations will be important. A valuable contribution will be characterizing burial history within each seep’s geological setting to assess more absolute timing of the generation and expulsion of the various charges identified in this study.

ACKNOWLEDGEMENTS

The authors thank Andrew Hanson and three anonymous reviewers for helpful suggestions, which improved the manuscript, and also thank editor Ryan Rodgers. ZB thanks the

Stanford University Basin and Petroleum System Modeling (BPSM) research group, the

U.S. Department of Energy National Energy Technology Laboratory (DOE NETL), and the Stanford McGee/Levorsen Research Grant Program for funding. ZB also thanks

Allegra Hosford Scheirer, Inessa Yurchenko, William Thompson-Butler, Jeremy Dahl, and the staff at Biomarker Technologies Incorporated and GNS Science for helpful comments and assistance. RS thanks the Ministry of Business, Innovation and Employment for funding through the GNS Science-led research program on New Zealand petroleum source rocks, fluids, and plumbing systems (contract C05X1507).

AUTHOR’S NOTE

This study (Chapter 1) is published in Energy & Fuels as:

Burton, Z. F. M.; Moldowan, J. M.; Sykes, R.; Graham, S. A. Unraveling petroleum

degradation, maturity, and mixing and addressing impact on petroleum

prospectivity: insights from frontier exploration regions in New Zealand. Energy &

Fuels 2018, 32 (2), 1287-1296.

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TABLES

Table 1.1. Diasterane and sterane peak assignment for Figure 1.3.

Table 1.2. Sterane and diasterane %C27−29 for all samples, illustrating similarity of the sterane and diasterane ratios.

Table 1.3. C31−35 17α-hopane 22S/(22S + 22R) isomerization ratios; from GC-MS-MS.

Table 1.4. Extent of cracking for the mixed-source samples. Calculated using %cracking

= [1 − (Co/Cc)] × 100 (after Dahl et al., 1999).

FIGURES

Figure 1.1. Area of interest and general setting; white circles indicate oil seep locations labeled as follows: R for Rotokautuku, W for Waitangi, T for Totangi, IC for Isolation

Creek, and K for Kaikoura; black lines indicate active plate boundary; frontier basins most relevant to this study, as well as the established Taranaki Basin, are labeled.

Figure 1.2. Whole-oil gas chromatograms show n-alkane and isoprenoid distributions and illustrate the contrast between 1) the very high preservation of n-alkane and isoprenoid compounds in the Totangi, Waitangi, and Rotokautuku oil seep samples and 2) the poor preservation, i.e., nearly complete destruction of n-alkanes and isoprenoids in the Isolation

Creek and Kaikoura oil samples; GC-FID chromatograms from GNS Science public report

(Sykes et al., 2012).

Figure 1.3. Representative m/z 217 fragmentograms of selected sterane distributions for the Waitangi and Kaikoura samples illustrate the preservation of sterane biomarkers and the similarity of sterane:diasterane ratios for all samples; Obtained via GC-MS of saturate hydrocarbon fractions.

Figure 1.4. Ternary diagrams show high degree of similarity in the relative abundances of

C27, C28, and C29 regular steranes versus βα-diasteranes (rearranged steranes); For steranes,

C27 = Total C27 / Total (C27 + C28 + C29) and so on for C28 and C29, for βα-diasteranes, C27

= C27/(C27 + C28 + C29) βα-diasteranes; Determined by GC-MS-MS of oil sample saturate hydrocarbon fractions.

Figure 1.5. Representative m/z 191 fragmentograms of selected hopane distributions from the Waitangi oil and the Kaikoura oil illustrate the good preservation of biomarkers in all five samples. Enlarged region highlights the preservation of the C31-C35 17α(H),21β(H)-

Homohopanes.

Figure 1.6. GC-MS-MS 398→191 transitions showing the C29 25-Nor-17α(H)-Hopane peak in a reference standard as a contrast to the lack of a discernible 25-norhopane peak in any of the samples analyzed in this study.

Figure 1.7. (a) C29 sterane isomerization and (b) C26 norcholestane versus C27 hopane maturity parameters. “Endpoint” areas indicate equilibration of isomerization ratios. All ratios calculated from GC-MS-MS measurements (C29 steranes monitored with m/z

400→217, C26 norcholestanes with m/z 358→217, and C27 hopanes with 370→191).

Figure 1.8. Concentrations of thermally stable diamondoids (methyldiamantanes) versus much less stable biomarkers (stigmastane) indicate thermal maturity and degree of cracking of oils, and reveal petroleum mixtures (e.g., Dahl et al., 1999).

CHAPTER 2

INTERPRETING SOURCE ROCK DEPOSITIONAL ENVIRONMENT

FROM SEEP OIL ISOTOPE AND ORGANIC GEOCHEMISTRY,

EAST COAST OF NEW ZEALAND

INTERPRETING SOURCE ROCK DEPOSITIONAL ENVIRONMENT

FROM SEEP OIL ISOTOPE AND BIOMARKER GEOCHEMISTRY,

EAST COAST OF NEW ZEALAND

Zachary F. M. Burton1, J. Michael Moldowan1,2, Leslie B. Magoon1, Richard Sykes3, and Stephan A. Graham1

1Department of Geological Sciences, Stanford University, Stanford, CA, USA

2Biomarker Technologies Incorporated, Rohnert Park, CA, USA

3GNS Science, Lower Hutt, New Zealand

ABSTRACT

Biomarker fingerprints of crude oil samples from four onshore East Coast Basin oil seeps were analyzed to assess source rock characteristics including type of organic matter input, redox conditions, sedimentary facies, and age. Results show that samples generally form two groups, correlating with geographic location: a northern and a southern group. Source rocks associated with all seep samples are interpreted to be marine. However, results suggest northern samples had more terrigenous organic matter input to their source rock(s), while southern samples had more marine input. Results suggest northern sample source rock(s) had more oxic depositional environments, whereas southern sample source rock had more reducing environments. A shale source rock sedimentary facies was indicated for all samples. These observations suggest that southern samples may be derived from slightly

46 higher quality source rocks (higher HI, deposited in more reducing conditions), although source rocks in both regions are oil prone. Biomarker age parameters suggest that the northern oil samples are from a younger (Cenozoic) source rock, whereas the southern oil samples are from an older (Cretaceous) source rock. Source rock characteristics

(depositional environment and age) point to the presence of two different source rocks. We postulate that northern oils samples from younger source rock with more terrigenous organic matter input represent the upper Paleocene Waipawa Formation, whereas southern oil samples from older source rock with more marine organic matter input represent the

Upper Cretaceous to Paleocene Whangai Formation.

2.1. INTRODUCTION

New Zealand’s northern east coast contains over 300 oil and gas seeps and shows (Uruski

2010), proving the presence of at least one petroleum system (Magoon and Dow 1994).

This suggests the opportunity for oil and gas exploration not only within the East Coast

Basin, but also within the adjacent Pegasus and Raukumara basins, both of which lie entirely offshore (Figure 2.1). Understanding of petroleum system characteristics of the three basins comprising the East Coast Province is limited. Efforts thus far have focused primarily on the structurally complex East Coast Basin, while the nature of its deep water extensions, the Raukumara and Pegasus basins, is even less constrained.

In frontier basins like the East Coast, Pegasus, and Raukumara basins, data are limited.

Because of this, whatever data are available are quite valuable. Onshore oil seeps such as those examined in this study can provide valuable insight as well as guide future research and petroleum exploration efforts. Biological marker (biomarker) fingerprints of the crude oil composing these onshore seeps provide information on the type of organic matter, the redox conditions, the sedimentary facies, and the age of the source rocks responsible for forming these oil seeps (Peters et al. 2005).

The purpose of this study is to collect biomarker data from four onshore oil seeps to investigate the origin of petroleum in the East Coast and Pegasus basins. The geochemical properties of each oil sample were evaluated in order to provide constraints on the type of organic matter input, the oxicity of the depositional environment, the sedimentary facies

(i.e., shale vs. carbonate), and the age of potential source rock(s). Our study expands upon and revises previous biomarker- and isotope-based investigations (Killops, 1996; Rogers

48 et al. 1999) of East Coast Basin oil seep samples, and establishes a rough oil-source rock correlation.

2.2. BACKGROUND

The East Coast, Pegasus, and Raukumara basins lie off the east coast of New Zealand

(Figure 2.1). At present, these basins lie along or straddle the plate boundary between the

Australian and Pacific plates (Figure 2.1). Subduction of the Pacific Plate beneath the

Australian Plate in the north (adjacent to the Raukumara and East Coast basins) transitions into strike-slip motion along the Alpine Fault of South Island (this transition from subduction to strike-slip occurs adjacent to Pegasus Basin). Although these three basins currently occupy a tectonically active continental margin setting, relative tectonic quiescence characterized much of the history of these basins, and deposition of petroleum source rocks (i.e., sediment capable of generating hydrocarbons) likely occurred in a passive margin setting. After subduction of the Hikurangi Plateau choked off subduction beneath the Chatham Rise (continental margin) of Gondwana in the late Early Cretaceous

(Davy 2014), a prolonged (~75 Myr) period of passive margin sedimentation with minimal tectonic activity began across the East Coast, Pegasus, and Raukumara basins (Bland et al.

2015). Not until ~25 Ma, as late as the earliest Miocene, did the second phase of Hikurangi

Plateau subduction begin (Ballance 1976; van der Lingen 1982; Spörli 1987; Lamb and

Bibby 1989; Nicol et al. 2007). Ongoing subduction highly deformed the East Coast Basin

(Rait et al. 1991; Nicol et al. 2007), whereas the Pegasus Basin remains relatively undeformed (Bland et al. 2015). Formation of the modern Hikurangi margin coincided with inception of the Alpine Fault ~24 Ma (Kamp 1986), which led to accelerated delivery of sediments to Pegasus Basin.

Passive margin Lower Cretaceous to Paleocene shales are the most likely candidates for viable source rock in these basins, both based on organofacies and depth of burial necessary

50 for thermal maturation (Uruski and Bland 2011). Uruski and Bland (2011) speculated about a Lower Cretaceous marine source rock, the Glenburn Formation (Figure 2.2) turbiditic shales containing terrestrially-derived 2 wt.% total organic carbon (TOC—a standard measure of the organic content of a petroleum source rock). However, the Whangai

Formation shale of Upper Cretaceous to Paleocene age and the Waipawa Formation black shale of late Paleocene (Thanetian) age (Figure 2.2) are most commonly invoked as potential source rocks (Schiøler et al. 2010; Uruski and Bland 2011; Hollis et al. 2014).

Both the Whangai and the Waipawa formations, the latter conformably overlying the

Whangai (Leckie et al. 1995), were deposited during the passive margin phase of Pegasus,

East Coast, and Raukumara basin sedimentation (Leckie et al. 1995; Hollis et al. 2014). In outcrop in the southern to central regions of interest (Marlborough to Wairarapa), the

Whangai Formation occurs as a widespread fine-grained mudstone unit truncated by an erosional unconformity as shown in Figure 2.2 (Bland et al. 2015). Deposition of the

Whangai Formation continued post-unconformity into Paleocene time, but as a calcareous mudstone member (and even sandstone member in areas) rather than a shaly mudstone

(Figure 2.2; Bland et al. 2015). At the end of the Paleocene, a change in sea level

(regression according to Schiøler et al. 2010; Hollis et al. 2014, transgression according to

Fuerst 2012) is associated with the deposition of the Waipawa Formation black shale, which occurs as a dark brown-grey mudstone or muddy sandstone (Figure 2.2; Hollis et al.

2014; Bland et al. 2015). Importantly, this sea level change led to increased delivery of terrigenous organic matter offshore, therefore the Waipawa Formation has been found to have elevated concentrations of terrigenous organic matter input relative to the Whangai

Formation (Schiøler et al. 2010; Fuerst 2012; Hollis et al. 2014). Rocks overlying the

Waipawa Formation are more calcareous in nature, and by the Oligocene carbonate deposition was widespread (Bland et al. 2015).

The Whangai and Waipawa formations have TOCs of 0.5-1.5 wt.% (Whangai Formation;

Uruski and Bland 2011) and 0.5-12.0 wt.% (Waipawa Formation; Hollis et al. 2014), and have average Hydrogen Indices (HI, a measure of how much hydrogen is available for conversion to hydrocarbons, used to characterize the “quality,” or petroleum producing potential, of a source rock; Tissot and Welte 1984) of 165 for Whangai Formation (246 rock samples; Sykes et al. 2012) and 212 for Waipawa Formation (153 rock samples; Sykes et al. 2012), though due to lateral variation, the quality of both potential source rocks could be much different offshore i.e., in the East Coast, Pegasus, and Raukumara basins.

The five oil seep samples (Rotokautuku, Totangi, Waitangi, Isolation Creek 1, and

Isolation Creek 2) analyzed in this study have previously been described, on the basis of biomarker and stable carbon isotope analyses, as originating from Cretaceous-Paleocene marine source rocks with minor terrigenous input (Johnston et al. 1991; Murray et al. 1994;

Rogers et al. 1994; Killops, 1996; Rogers et al. 1999). Due to similarities in biomarker distributions, as well as generally light carbon isotope signatures, the five oil seep samples examined in this study have previously been characterized as being derived from the same source rock (Rogers et al. 1999), or as belonging to the same family (Sykes et al. 2012).

A study of the five oil seep samples by Burton et al. (2018) revealed that all oil samples rank low on the Peters & Moldowan biodegradation scale (Peters and Moldowan 1993).

The low degrees of biodegradation seen in these oil samples suggests the suitability of application of biomarker-based techniques in assessing the source rock depositional environment, organic matter input, and sedimentary facies (see Burton et al. (2018) for 52 whole-oil gas chromatograms and complete discussion of biodegradation). This study also revealed that all five oil samples have some component of early to peak oil window oil, and show good preservation of steranes and hopanes (see Burton et al. (2018) for sterane and hopane distributions). Based on Burton et al. (2018) biomarker parameters should be viable, with minimal biasing effect of biodegradation or thermal maturity. Nevertheless, n- alkane and isoprenoid parameters (i.e., those parameters most susceptible to biodegradation) are not included in this study due to non-negligible amounts of biodegradation seen in these oil samples.

2.3. MATERIALS AND METHODS

2.3.1. MATERIALS

Five oil seep samples from East Coast Basin were obtained from GNS Science for use in this study (Figure 2.1). Details concerning the collection and nature of these samples can be found in Sykes et al. (2012). We analyzed three oil samples (the northern oil seep samples from the Rotokautuku, Totangi, and Waitangi seeps) from the Raukumara

Peninsula of North Island and two oil samples (the southern oil seep samples from the

Isolation Creek seep) from the Marlborough Region of South Island. We utilized two oil samples collected at different dates from the Isolation Creek seep as a means of controlling for any temporal heterogeneity in the quality (i.e., apparent respective depositional environment characteristics) of the oil seep.

2.3.2. METHODS

2.3.2.1. Stable carbon isotopes

Stable carbon isotope data for bulk (saturated and aromatic) fractions were provided by

GNS Science. These analyses were conducted by Applied Petroleum Technology (APT),

Norway following methods outlined by Weiss et al. (2000).

2.3.2.2. Sulfur

Wt.% sulfur data for the five oil samples were obtained from GeoMark Research.

2.3.2.3. Liquid chromatography

Liquid chromatography was carried out at Biomarker Technologies Incorporated. Crude oil samples (30-40 mg) were weighed out and spiked. Spiking solutions added in solution

54 with known concentrations by volumetric pipette contained 5β-cholane as the internal standard in quantitative GC-MS and GC-MSMS for analysis of biomarkers. Spiked samples were fractionated by sequential elution. A proprietary light hydrocarbon solvent and dichloromethane were used on silica gel columns to obtain both saturate and aromatic hydrocarbon fractions. A proprietary light hydrocarbon solvent was used on zeolite columns to remove n-alkanes from the saturated hydrocarbon fractions.

2.3.2.4. GC-MS

Gas chromatography-mass spectrometry (GC-MS) was used to analyze saturate hydrocarbon fractions at Biomarker Technologies Incorporated. We used an Agilent

Technologies 7890A/5975C GC-MS system with a 7693A autosampler. The GC was equipped with a 60 m DB-1 column with 0.25 mm i.d. and 0.25 μm phase thickness.

Helium was used as the carrier gas. The temperature program was started at 80°C and was subsequently raised at a rate of 2°C/minute until 320°C. Temperature was then held at

320°C for 15 minutes. In order to account for response factors of different compounds, calibrated external standards were run alongside the oil samples.

Ions with specific mass-to-charge ratios (m/z) were monitored during biomarker analysis.

Biomarkers were quantified using m/z 191 for hopanes and m/z 217, 218, and 259 for steranes.

Compound peak areas were integrated and/or peak heights were measured relative to both internal and external standards in order to achieve biomarker quantitation.

2.3.2.5. GC-MS-MS

Gas chromatography-tandem mass spectrometry (GC-MS-MS) was also used to analyze saturated hydrocarbon fractions at Biomarker Technologies Incorporated. We used an

Agilent Technologies 7890A GC interfaced to a 7000 Triple Quad GC-MS with a 7693A autosampler. The GC was equipped with a 60 m DB-1 column with 0.25 mm i.d. and 0.25

μm phase thickness. Helium was used as the carrier gas. The temperature program was started at 80°C and was subsequently raised at a rate of 2°C/minute until 320°C.

Temperature was then held at 320°C for 15 minutes. Oil samples were analyzed alongside calibrated external standards so as to account for differences in chemical compound response factors.

The GC was run in parent-to-daughter mode using the first quadruple to focus parent ions and the third quadrupole to focus daughter ions. The transitions monitored were m/z

370→177, 318→191, 330→191, 332→191, 370→191, 384→191, 398→191, 412→191,

426→191, 440→191, 454→191, 468→191, 482→191, and 426→205 for hopane biomarkers and 288→217, 302→217, 330→217, 358→217, 372→217, 386→217,

400→217, 414→217, and 414→231 for sterane biomarkers.

Compound peak areas were integrated and/or peak heights were measured relative to both internal and external standards so as to quantify biomarker compounds.

2.4. RESULTS AND DISCUSSION

2.4.1. SOURCE ROCK ORGANIC MATTER INPUT

2.4.1.1. Stable carbon isotopes

Stable carbon isotopic composition of saturated versus aromatic C15+ hydrocarbon fractions often successfully distinguish crude oil generated by source rocks containing predominantly marine or predominantly terrigenous organic matter (Sofer 1984). Isotope

13 compositions for the northern oil samples range from –29.8 to –29.1 for δ Csaturates and

13 from –27.9 to –27.6 for δ Caromatics while isotopic compositions for the southern oil

13 13 samples range from –27.8 to –27.6 for δ Csaturates and from –27.1 to –26.9 for δ Caromatics

(Table 2.1). The δ13C values observed in the northern oil samples are more negative

(isotopically lighter) than the less negative (isotopically heavier) δ13C values of the southern crude oil samples.

13 13 The oil sample δ Csaturates versus δ Caromatics values (Table 2.1) are plotted on the crossplot

(Figure 2.3) first used by Sofer (1984) to discriminate between marine versus terrigenous input. This plot shows a best-fit separation line, based on statistical analysis of 339 oil samples by Sofer (1984), that delineates terrigenous (above line) and marine (below line) source rock organic matter provenance. The northern oil samples all fall above the Sofer line, suggesting that these oil samples were derived from a more terrigenous organic facies

(Figure 2.3), whereas the southern oil samples plot well below the Sofer line, indicating a predominance of marine organic matter input (Figure 2.3).

Stable isotopes tend to offer a more robust quantification of organic matter input than conventional chromatographic assessment of crude oil waxiness, though caution must be

57 exercised if dealing with biodegraded or over-mature fluid (Peters et al. 2005). Sofer (1984) derived a statistical parameter referred to as the canonical variable (CV) and described by

13 13 the equation CV = –2.53 δ Csaturates + 2.22 δ Caromatics – 11.65. Oil samples with CV values above 0.47 are classified as waxy terrigenous oil, whereas oil samples with CV values below 0.47 are classified as non-waxy marine oil. The northern oil samples have CV values above 0.47 (ranging from ~0.7 to 1.8), indicating terrigenous organic matter input (Table

2.1). In contrast, the southern oil samples yield CV values well below 0.47 (–1.5 for both samples), indicating marine input (Table 2.1).

2.4.1.2. Biomarkers and sulfur

Previous work has shown that the presence of certain C30-steroids (24-n-propylsteranes and

24-n-propyldiasteranes) in oil is a robust indication of input of marine organic matter to a source rock (Moldowan et al. 1985; Peters et al. 1986; Moldowan et al. 1990), and that, conversely, a C30 sterane index of zero tends to indicate non-marine oil (e.g., Holba et al.

2000). The C30 sterane index (C30/(C27—C30); e.g., Moldowan et al. 1992; Peters et al.

2005) is greater than zero for all of these oil samples (Table 2.2), indicating at least some component of marine organic matter input to all of the source rocks that produced these oil seeps.

The abundance of C29 steranes tends to be higher in oil from terrigenous source rock organic matter (Moldowan et al. 1985; Peters et al. 2005). The C27-C28-C29 sterane ternary diagram (Figure 2.4) shows that the northern oil samples all have slightly higher C29 than the southern oil samples, suggesting slightly more terrigenous organic matter input to the source rocks that produced the northern oil seeps.

Figure 2.5 shows sulfur content plotted against the oleanane index of these oil samples. At

0.09%, 0.18%, and 0.23%, sulfur content is lower for the northern oil samples than for the southern oil samples, which have sulfur contents of 0.37% and 0.42%, respectively (Figure

2.5). Sulfur content tends to be higher in reducing depositional settings (Doff, 1969), which in turn correlate with more marine depositional settings (Didyk et al. 1978). In contrast, lower sulfur contents correspond to more oxic settings (Doff, 1969), which tend to be more terrigenous in nature (Didyk et al. 1978). Therefore, on the basis of sulfur content, it appears that the northern oil samples all have more terrigenous organic matter input than the more marine southern oil samples. This inference corroborates well with the distributions of C29 steranes (Figure 2.4) and oleanane indices (Figure 2.5) of these samples.

Oleanane is a key indicator of higher plant, particularly angiosperm, input to source rocks

(Whitehead 1974; Grantham et al. 1983; ten Haven and Rullkotter 1988). Oleanane indices for these oil samples are all quite low (e.g., Moldowan et al. 1994; Peters et al. 1999), however, important distinctions are apparent. The northern oil samples all have higher oleanane indices than the southern oil samples, and the northern Totangi and Waitangi oil samples have indices more than twice those of the southern Isolation Creek oil samples.

This suggests a greater degree of higher plant input, i.e., greater terrigenous organic matter input, to the northern oil samples.

The tetracyclic polyprenoid (TPP) ratio is effective in indicating lacustrine organic matter input, whereby lacustrine crude oil samples have high TPP ratios (Holba et al. 2000).

Figure 2.6 shows TPP ratios plotted against %C30 24-propyldiacholestanes, which are indicators of marine Chrysophyte algal input, and thus, are absent in lacustrine oil (Peters

59 et al. 2005). All of the oil samples examined in this study have very low TPP ratios and have non-zero values for C30 24-propyldiacholestane ratios. The negligible TPP ratios preclude any of these oil samples from being lacustrine in nature, while the non-zero C30

24-propyldiacholestane ratios indicate that these oil samples originated from marine organic matter, albeit with varying amounts of terrigenous organic matter input, as discussed above.

2.4.2. SOURCE ROCK DEPOSITIONAL ENVIRONMENT OXICITY

Understanding the relative oxicity of a source rock’s depositional environment is critical to understanding source rock quality, as the relative oxicity of sediment and the water column directly affects the degree to which organic matter is preserved and incorporated into the sediment (Didyk et al. 1978). Biological markers can be used to assess the oxicity at the time of organic matter deposition.

The C35 homohopane index is an indicator of redox potential in marine sediments, whereby high values indicate reducing conditions, and low values indicate oxic conditions (Peters and Moldowan 1991). Figure 2.7 shows that the northern Rotokautuku oil sample has the lowest C35 homohopane index, and the northern Totangi and Waitangi oil samples have lower C35 homohopane indices than the southern Isolation Creek oil samples. This suggests that more oxic conditions prevailed during deposition of the organic matter contributing to the northern oil samples, whereas more reducing conditions were present during deposition of the organic matter contributing to the southern oil samples.

As discussed above for Figure 2.5, sulfur content is lower for the northern oil samples than for the southern oil samples. Because lower sulfur contents tend to correspond to more oxic

60 depositional settings, while higher sulfur contents correspond to more reducing depositional settings (Doff, 1969), the northern oil samples appear to originate from organic matter deposited under more oxic conditions, whereas the southern oil samples appear to originate from organic matter deposited under more reducing conditions (Figure

2.7). The sulfur data compare well with the C35 homohopane indices (Figure 2.7).

Plotting gammacerane indices versus C35/C34 homohopanes yields similar results (Figure

2.8). Elevated amounts of gammacerane are indicative of highly reducing, hypersaline conditions (Moldowan et al. 1985; Jiamo et al. 1986) and indicate a stratified water column

(Damsté et al. 1995). Gammacerane indices for the southern oil samples are about twice those of the northern oil samples. This suggests that organic matter contributing to the southern oil samples was deposited under more reducing conditions than the organic matter contributing to the northern oil samples, which was deposited under more oxic conditions.

Homohopane distributions can also be useful in assessing redox conditions during source rock deposition (Peters et al. 2005). Here C35/C34 homohopanes are plotted (Figure 2.8), and indicate redox conditions whereby elevated C35 homohopanes are indicative of more reducing conditions (Peters et al. 2005). Southern oil samples have higher C35/C34 homohopanes, indicating more reducing source rock depositional conditions, whereas the northern oil samples have lower C35/C34 homohopanes, indicating a more oxic source rock depositional environment (Figure 2.8).

2.4.3. SOURCE ROCK SEDIMENTARY FACIES

Various parameters have been used to distinguish oil originating from either carbonate or siliciclastic (i.e., shale) source rocks (Peters et al. 2005). One such parameter is the

61 diasteranes/steranes ratio (e.g., Mello et al. 1988). Clays catalyze the transformation of steroids to diasteranes, meaning that higher clay content in a source rock will lead to greater conversion of steroids to diasteranes, and thus, a higher diasteranes/steranes ratio. Table

2.3 shows diasteranes/steranes ratios for the oil samples examined in this study. Of note is that ratios for all oil samples are high to very high (>1.0), indicating that these oil samples all originated from clay-rich (i.e., shale) source rock(s). This corroborates quite well the eight parameters examined in Figure 2.9.

Elevated concentrations of C24 tetracyclic terpane tend to indicate carbonate or evaporate depositional environments (Palacas et al. 1984; Connan et al. 1986; Connan and Dessort

1987; Clark and Philp 1989). Ratios of C24 tetracyclic terpane/hopane, along with ratios of gammacerane/hopane (e.g., Wang et al. 2014), can therefore be used to indicate carbonate versus clay-rich source rocks (Figure 2.9A). All oil samples examined in this study have relatively low C24 tetracyclic terpane/hopane ratios and low gammacerane/hopane ratios, indicating clay (shale) depositional environments for all oil samples. Ts/(Ts+Tm) and C27 diasterane/(diasterane+regular sterane) ratios corroborate these findings (Figure 2.9B).

Lower Ts/(Ts+Tm) ratios tend to correlate with carbonate depositional environments

(McKirdy et al. 1983), while lower C27 diasterane/(diasterane+regular sterane) ratios also indicate carbonate depositional environments (Moldowan et al. 1986). Diasteranes occur with increasing clay content in the depositional environment, i.e., with increasingly shaly sedimentary facies. Therefore, the relatively high Ts/(Ts+Tm) ratios and C27 diasterane/(diasterane+regular sterane) ratios indicate a shale depositional environment. A plot of C29/C30 hopane versus C35/(C34+C35) homohopane agrees with these findings

(Figure 2.9C), in that the lower C29/C30 hopane ratios and lower C35/(C34+C35) homohopane

62 ratios indicate a shale depositional environment (ten Haven et al. 1988; Peters et al. 2005;

Wang et al. 2014). A shale depositional environment is also indicated by a plot of C22/C21 tricyclic terpane versus C24/C23 tricyclic terpane (Figure 2.9D), whereby oil samples have low C22/C21 tricyclic terpane and high C24/C23 tricyclic terpane, both characteristic of shale depositional environments (Peters et al. 2005).

The nine parameters examined here (diasterane ratio of Table 2.3 and eight parameters of

Figure 2.9) all indicate a shale source rock facies, that is, a shale depositional environment for all of the oil samples examined in this study.

2.4.4. SOURCE ROCK AGE

A large set of Cretaceous-Eocene New Zealand oil seep and well samples described in

Sykes et al. (2012) shows ages, based on higher plant biomarker work of previous studies

(Killops et al., 1995, 2003; Sykes et al. 2012), correlated with carbon isotope values for these samples (Figure 2.10). This data set shows that younger oil samples have lighter

(more negative) carbon isotope values, whereas older oil samples generally have isotopically heavier (less negative) carbon isotope values (Figure 2.10). Comparison of oil samples described in our study with the larger set of New Zealand oils shows close agreement between the isotopically lighter northern oil samples and the Paleocene-

Eocene/Eocene oil samples (Figure 2.10). Similarly, there is close agreement between the isotopically heavier southern oil samples and the mid-Late Cretaceous/Late Cretaceous oil samples (Figure 2.10). This suggests the northern oil samples originate from younger source rock(s) than the southern oil samples.

The correlation between age and isotopic composition of these New Zealand oils appears relatively robust, though isotopic correlations should be bolstered by other parameters.

Therefore, oleanane indices and C26 nordiacholestane ratios are examined (Figure 2.11). In addition to being a higher plant indicator, as described above, abundance of oleanane has been roughly correlated with source rock age (Moldowan et al. 1994). Most oleanane first appeared in the Upper Cretaceous, and since then (through the Cenozoic) abundances in source rocks have generally increased (Moldowan et al. 1994). Oleanane abundances are somewhat higher in the northern oil samples, whereas they are somewhat lower in the southern oil samples (Figure 2.11). This indicates that the northern oil samples may originate from younger source rock(s), whereas the southern oils samples may originate from older source rock, and agrees well with the above-described isotopic age data for these samples.

A third parameter, the C26 24-nordiacholestane to 27-nordiacholestane ratio, has also been shown to have a strong relationship with source rock age (Holba et al. 1998a, b; Peters et al. 2005). 24-nordiacholestanes increase in the Cretaceous, and increase dramatically in the

Cenozoic, coincident with the proliferation of diatoms (Holba et al. 1998a, b). Empirical relationships using 150 oil samples suggest that %24-nordiacholestane ratios above 25 indicate Cretaceous or younger oil samples, while ratios above 55 indicate Cenozoic

(generally Neogene) oil samples (Holba et al. 1998a, b; Peters et al. 2005). Examination of

24-nordiacholestane ratios for the oil samples discussed here reveals that all oil samples are Cretaceous or younger (ratio above 25), but that the northern oil samples all fall in the realm of Cenozoic oils (ratio above 55) (Figure 2.11). This suggests that the northern oil

64 samples are from source rock that is likely Cenozoic in age, whereas the southern oil samples are from a source rock that is likely Cretaceous in age.

2.4.5. IMPLICATIONS

Our results suggest the possibility of regionally extensive source rocks in the East Coast

Basin and, by extension, in the Pegasus and Raukumara basins. Tetracyclic polyprenoid ratios and C30 24-propyldiacholestane ratios indicate that source rocks present in the region are generally marine in nature. This, paired with the indication (via nine unique parameters) of a shale source rock sedimentary facies, and paired with favorable TOCs and HIs seen in onshore analogues, may suggest the presence of regionally extensive, oil-prone marine source rocks in the East Coast, Pegasus, and Raukumara basins.

Important differences between the northern and southern oil samples examined in this study suggest the presence of at least two distinct marine source rocks, revising interpretations by previous studies (Rogers et al. 1999; Sykes et al. 2012). A key difference is the interpreted age of the source rocks associated with the northern and southern oil samples. Previous studies assumed one Cretaceous-Paleocene source rock to be associated with the northern and southern oil seeps (Rogers et al. 1999; Sykes et al. 2012). However, this study suggests that a younger, Cenozoic source rock is associated with the northern seeps, whereas an older, Cretaceous source rock is associated with the southern seeps. We hypothesize that this indicates that the northern oil samples are sourced from the upper

Paleocene Waipawa Formation, whereas the southern oil samples are sourced from the

Upper Cretaceous-Paleocene Whangai Formation. Furthermore, the southern oil samples potentially originate from the shaly Late Cretaceous member of the Whangai Formation rather than the more calcareous Paleocene Whangai Formation member, as evinced by the

65 nine source rock facies parameters examined here (all of which indicate a shale rather than carbonate source rock facies). The age-based interpretation is further supported by the finding that the northern oils have more terrigenous organic matter input. The Waipawa

Formation is known to have higher terrigenous organic matter content, due to erosion of organic material by regression (Schiøler et al. 2010; Hollis et al. 2014) or transgression

(Fuerst 2012) onshore, and due to relative proximity of the shoreline. The elevated concentrations of terrigenous organic matter associated with the northern oil samples therefore suggest these oils may be associated with the Waipawa Formation, whereas the more marine southern oil samples may be associated with the more marine Whangai

Formation. The northern oils also have more oxic depositional environments associated with their respective source rock(s), whereas the southern oils have more reducing depositional environments associated with their respective source rock. This suggests that the southern oils may represent slightly better quality (higher HI, more reducing depositional environment) source rock. Overall, these results suggest presence of oil- generating Waipawa Formation may be more likely for the northern East Coast Basin and

Raukumara Basin, whereas presence of oil-generating Whangai Formation may be more likely for the southern East Coast Basin and Pegasus Basin.

2.5. CONCLUSIONS

This study presents an assessment of the type of organic matter input, the oxicity of the depositional environment, the sedimentary facies, and the age of the source rocks for five oil seep samples from New Zealand’s East Coast Basin. This assessment reveals important differences between the northern oil seep samples and the southern oil seep samples, and likely indicates the presence of at least two distinct source rocks. Study of the carbon isotopes, oleanane indices, and 24-nordiacholestane ratios suggests that the northern oil samples are associated with a Cenozoic source rock (inferred to be the upper Paleocene

Waipawa Formation), whereas the southern oil samples are associated with a Cretaceous source rock (inferred to be the Upper Cretaceous-Paleocene Whangai Formation). The examination of sulfur content, C27-C28-C29 steranes, and oleanane indices indicate that the northern oil samples have a greater amount of terrigenous organic matter input to their respective source rock(s) than the southern oil samples, which have a greater amount of marine organic matter input. This further supports the interpretation that the northern oil samples may be associated with the Waipawa Formation, which is known to have higher terrigenous organic matter content. Oil samples from both areas are generally marine in nature, as revealed by examination of tetracyclic polyprenoid ratios and C30 24- propyldiacholestane ratios. This is consistent with the generally marine nature of the

Waipawa Formation and Whangai Formation. Study of sulfur content, C35 homohopane indices, gammacerane indices, and C35/C34 homohopanes indicate important differences in the redox potential of the source rock depositional environment of these oil samples. The northern oil samples all tend to have oxic depositional environments, whereas the southern oil samples are associated with more reducing depositional environments. This is

67 consistent with past work suggesting the Waipawa Formation was deposited at shallower, and perhaps more oxic, water depths. Lastly, the source rock sedimentary facies associated with these oil samples was studied via assessment of nine distinct parameters and was found to be a shale depositional environment for both the northern and the southern oil samples examined in this study. This is consistent, in that both the Waipawa Formation and Whangai Formation are described as shale. This study provides important new constraints on source rock characteristics, and presents a tentative oil-source rock correlation for both the northern and southern oil seep samples.

ACKNOWLEDGEMENTS

The authors thank Andrew Hanson and Gary Muscio for helpful reviews, which improved the manuscript, and also thank editors Wolf-Christian Dullo and Peter Kukla. ZB thanks the Stanford University Basin and Petroleum System Modeling (BPSM) Industrial

Affiliates Program, the U.S. Department of Energy National Energy Technology

Laboratory (DOE NETL), the AAPG Grants-In-Aid Program, and the Stanford

McGee/Levorsen Research Grant Program for funding. ZB also thanks Allegra Hosford

Scheirer, Inessa Yurchenko, William Thompson-Butler, and the staff at GNS Science,

Biomarker Technologies, Inc., and GeoMark Research for useful comments and assistance.

RS would like to thank the Ministry of Business, Innovation and Employment for funding through the GNS Science-led research program on New Zealand petroleum source rocks, fluids, and plumbing systems (contract C05X1507).

AUTHOR’S NOTE

This study (Chapter 2) is published in International Journal of Earth Sciences as:

Burton ZFM, Moldowan JM, Magoon LB, Sykes R, Graham SA (2019)

Interpretation of source rock depositional environment and age from seep oil, east

coast of New Zealand. International Journal of Earth Sciences 108:1079-1091

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TABLES

Table 2.1. Saturate and aromatic bulk carbon isotope data for the five oil samples (from

GNS Science), canonical variable (CV value) as calculated according to Sofer (1984), and terrigenous vs. marine designation according to CV value.

Table 2.2. C30 sterane indices (%C30 sterane index = 100 × Total C30/Total

(C27+C28+C29+C30)) indicate at least some component of marine organic matter input to all oil seep samples; calculated via GC-MS-MS.

Table 2.3. Total regular steranes, total diasteranes, and diasterane ratio (diasteranes/regular steranes) illustrating high to very high (>1.0) diasterane ratios in all of these samples; measured via GC-MS-MS.

FIGURES

Figure 2.1. Map of northern New Zealand; oil seep locations are indicated with various polygons (see legend in bottom-right corner for oil seep names); frontier basins most relevant to this study, as well as the established Taranaki Basin, are labeled; black lines indicate active plate boundary; modified after Burton et al. (2018).

Figure 2.2. Chronostratigraphic chart based on outcrop stratigraphy from the central portion of the study area, modified after Bland et al. (2015), plus regional tectonic history.

13 13 Figure 2.3. δ Csaturates versus δ Caromatics with Sofer (1984) line, indicating that the northern oil samples fall above the line in the terrigenous realm, whereas the southern oil samples fall below the line in the marine realm; data from GNS Science.

Figure 2.4. Ternary diagram of relative C27, C28, and C29 sterane abundances for these oil samples, indicating that the northern oil samples have slightly higher C29 steranes, and thus slightly higher terrigenous organic matter input, than the southern oil samples; symbols as in Figure 2.3; data from GC-MS-MS.

Figure 2.5. Plot of wt. % sulfur versus oleanane index, showing that the northern oil samples have higher oleanane ratios, and thus slightly more terrigenous organic matter input than southern oil samples; the southern oil samples have higher wt. % sulfur than the northern oil samples indicating slightly more reducing conditions, and thus more marine conditions in the southern oil samples; symbols as in Figure 2.3; wt. % sulfur from

GeoMark; oleanane index calculated from GC-MS-MS data; oleanane index = (18α- oleanane +18β-oleanane)/(18α-oleanane+18β-oleanane+C30 Hopane).

Figure 2.6. Tetracyclic polyprenoid (TPP) ratios versus C30 24-propyldiacholestane ratios

(after Holba et al., 2000) for these oil samples; all oil samples have negligible TPP ratios, indicating no lacustrine organic matter input, and have %C30 24-propyldiacholestane values above zero, indicating some amount of marine organic matter in all of these oil samples; symbols as in Figure 2.3; calculated from GC-MS-MS data; TPP ratio = (2×TPP-

1)/((2×TPP-1)+C26 ααα-27-nor-20S+C26 αββ-27-nor-20R+C26 αββ-27-nor-20S+C26 ααα-

27-nor-20R); %C30 diasteranes = 100×C30/(C27+C28+C29+C30) βα-diasteranes.

Figure 2.7. Wt. % sulfur versus C35 homohopane index indicates more reducing depositional conditions for source rock of the southern oil samples and more oxic conditions for source rock(s) of the northern oil samples; symbols as in Figure 2.3; wt. % sulfur from GeoMark; C35 homohopane index calculated from GC-MS-MS data; C35 homohopane index = (C35 αβ-Hopane (22S)+C35 αβ-Hopane (22R))/(C35 αβ-Hopane

(22S)+C35 αβ-Hopane (22R)+C34 αβ-Hopane (22S)+C34 αβ-Hopane (22R)+C33 αβ-Hopane

(22S)+C33 αβ-Hopane (22R)+C32 αβ-hopane (22S)+C32 αβ-hopane (22R)+C31 αβ-hopane

(22S)+C31 αβ-hopane (22R)).

Figure 2.8. Gammacerane index versus C35/C34 homohopanes indicates more reducing source rock depositional conditions associated with the southern oil samples and more oxic depositional conditions for the northern oil samples; symbols as in Figure 2.3; data from

GC-MS-MS; gammacerane index = Gammacerane/(Gammacerane+C30 Hopane).

Figure 2.9. Various biomarker parameters all indicate a primarily shale source rock facies for these oil samples; symbols as in Figure 2.3; data from GC-MS and GC-MS-MS.

13 13 Figure 2.10. δ Csaturates versus δ Caromatics of oils discussed herein plus numerous

Cretaceous-Eocene oils from New Zealand, indicating the southern oil falls in line with mid-Late and Late Cretaceous oils, whereas the northern oil samples fit with Eocene oils;

Sofer (1984) line shown as dashed line; data from GNS Science.

Figure 2.11. Abundance of C26 24-Nordiacholestanes relative to C26 27-Nordiacholestanes

(after Holba et al. 1998a, b; Peters et al. 2005) versus oleanane index (after Moldowan et al. 1994); Cretaceous or younger age is indicated for all oil samples, however, a Cenozoic

(and potentially Neogene) age is indicated for the northern oil samples according to C26

Nordiacholestanes; oleanane abundances are also higher in the northern oil samples, suggesting a younger age relative to the southern oil samples; symbols as in Figure 2.3; data from GC-MS-MS.

CHAPTER 3

TECTONIC UPLIFT DESTABILIZES SUBSEA GAS HYDRATE:

A MODEL EXAMPLE FROM HIKURANGI MARGIN, NEW ZEALAND

TECTONIC UPLIFT DESTABILIZES SUBSEA GAS HYDRATE:

A MODEL EXAMPLE FROM HIKURANGI MARGIN, NEW ZEALAND

Zachary F. M. Burton1,2, Karsten F. Kroeger3, Allegra Hosford Scheirer1, Yongkoo

Seol2, Blair Burgreen-Chan4, and Stephan A. Graham1

1Department of Geological Sciences, Stanford University, Stanford, CA, USA

2U.S. Department of Energy National Energy Technology Lab, Morgantown, WV, USA

3GNS Science, Lower Hutt, New Zealand

4ConocoPhillips, Houston, TX, USA

ABSTRACT

Rising ocean temperatures and falling sea level are commonly cited as mechanisms of marine gas hydrate destabilization. More recently, uplift – both isostatic and tectonic – has been invoked. However, the effect of tectonic shortening and uplift on gas hydrate stability zone extent has not been validated via integrated computational modeling. Here, modeling along the Hikurangi margin of New Zealand illustrates the mechanism of tectonic uplift as a driver of gas hydrate destabilization. We simulate how tectonic uplift and shortening affect the presence and decrease the extent of a gas hydrate stability zone. We suggest that resultant gas hydrate destabilization in the marine realm may impact the global carbon cycle and oceanic chemistry over geologic time.

3.1. INTRODUCTION

Gas hydrate is a frozen solid and a reservoir of highly concentrated carbon gases

(Kvenvolden, 1988; Sloan, 2003; Sloan and Koh, 2007). Gas hydrates therefore are an important interface in carbon cycling between geosphere, ocean, and atmosphere

(Kvenvolden, 1993; Kvenvolden, 2002; Kroeger et al., 2011). The reservoir for >98% of global gas hydrate is in marine sediments seaward of continental shelfs (McIver, 1981;

Kvenvolden, 1988; Ruppel, 2014), where low-temperature and elevated-pressure conditions promoting stability of gas hydrate prevail.

Estimates of carbon in gas hydrates equal or exceed carbon in all fossil fuel accumulations

(Kvenvolden, 1988; Milkov, 2004). Destabilization of large amounts of gas hydrate and potential release of methane could affect the global carbon cycle and climate (Kvenvolden,

2002; Nisbet and Chappellaz, 2009). This hydrate-mediated effect is suggested for the early

Toarcian oceanic anoxic event (Hesselbo et al., 2000) and the Paleocene-Eocene Thermal

Maximum (PETM) (Dickens et al., 1995; Kaiho et al., 1996; Dickens et al., 1997; Katz et al., 1999; Thomas et al., 2002), whereby warming destabilized gas hydrate, perhaps releasing methane to the atmosphere and warming global climate. However, it should be noted that the causes of the PETM remain in question (McInerney and Wing, 2011).

Alongside ocean warming or falling sea level, isostatic rebound (Wallmann et al., 2018) and tectonic uplift (Jahren et al., 2005; Maclennan and Jones, 2006) have been recently invoked as regional mechanisms for gas hydrate destabilization and catalysts of global carbon cycle changes. While methane from marine hydrate destabilization is unlikely to reach the atmosphere directly (Ruppel and Kessler, 2017), it may contribute to ocean

94 acidification and the oceanic carbon cycle, ultimately impacting global climate (Biastoch et al., 2011; Boudreau et al., 2015).

We examine uplift as a gas hydrate destabilization mechanism along the Hikurangi subduction margin of New Zealand (Figure 3.1). After a phase of passive margin evolution from the Cretaceous to earliest Miocene (Ballance, 1976; van der Lingen, 1982; Spörli,

1987; Lamb and Bibby, 1989; Nicol et al., 2007; Davy, 2014), the Hikurangi margin has undergone three phases of subduction-related deformation along the Hawke’s Bay transect discussed here (Cole and Lewis, 1981; Pettinga, 1982; Chanier, 1991; Rait, 1992; Barnes et al., 2002; Nicol et al., 2002; Nicol and Beavan, 2003; Bailleul et al., 2007; Nicol et al.,

2007; Burgreen-Chan et al., 2016).

Along the Hikurangi margin, evidence for fluid flow and methane leakage (Lewis and

Marshall, 1996; Pecher et al., 2004; Faure et al., 2006; Faure et al., 2010; Greinert et al.,

2010; Naudts et al., 2010; Schwalenberg et al., 2010; Krabbenhoeft et al., 2013), generation of biogenic gas and thermogenic gas from regional source rocks (Kroeger et al., 2015;

Burton et al., 2018, 2019), and presence of gas hydrate (Katz, 1982; Townend, 1997;

Henrys et al., 2003; Pecher et al., 2004; Faure et al., 2006; Greinert et al., 2010; Pecher et al., 2010; Crutchley et al., 2010) is extensive. Seafloor pockmarks, possible evidence for past gas escape (King and MacLean, 1970; Hovland and Judd, 1988; Kelley et al., 1994;

Andreassen et al., 2017), are identified along (Nelson and Healy, 1984) and outboard (Davy et al., 2010; Collins et al., 2011) of the Hikurangi margin.

Here we assess potential for methane release from gas hydrates along the Hikurangi margin by focusing on gas hydrate stability zone (GHSZ) dynamics. We present a model of GHSZ

95 distributions through time for sequential structural restorations, and illustrate widespread shrinking of the GHSZ due to tectonic uplift.

3.2. METHODS

The PetroMod™ v.2016 2D basin and petroleum system model discussed here is adapted from the East Coast Basin Hawke’s Bay model of Burgreen-Chan et al. (2016) and

Burgreen-Chan and Graham (2018). We utilize this model based on seismic interpretation

(Figure 3.2) (after Barnes et al., 2002) and structural reconstruction (after Barnes and Nicol,

2004) of the 124-kilometer Hawke’s Bay CM05-01 seismic. The structural reconstruction by Burgreen-Chan et al. (2016) was performed using geomechanically-based restoration software, which forward models changing tectonic conditions within a finite element framework (e.g., Maerten and Maerten, 2006). Paleo-basin geometries are defined following structural restorations at time steps 24 Ma, 18.8 Ma, 13.6 Ma, 11.9 Ma, 8.5 Ma,

7.0 Ma, 4.0 Ma, and 0.8 Ma, as well as at present day (Figure 3.2) as in Burgreen-Chan et al. (2016). Reference locations along these modeled time steps were chosen at 20, 40, 60,

80, and 100 km along the transects. With 21 kilometers of estimated shortening since 24

Ma (Burgreen-Chan et al., 2016), the reconstructed CM05-01 line used here represents a conservative estimate for total amount of shortening, as opposed to estimates of 35±7 kilometers of shortening by Nicol et al. (2007). Two end-member scenarios for basal paleo- heat flow were adopted (scenario one with 40 mW/m2 from 100-5 Ma and cooling to 30 mW/m2 from 5 Ma-present; scenario two with 55 mW/m2 from 100-15 Ma and cooling to

30 mW/m2 from 15 Ma-present), as in Burgreen-Chan and Graham (2018). These heat flow scenarios are calibrated to paleothermometer data from East Coast Basin wells, as described in Burgreen-Chan and Graham (2018). Sediment-water interface temperature for the present day is adapted from Ridgway (1969). Paleo-sediment-water interface temperatures are adapted from Burgreen-Chan and Graham (2018) after Wygrala (1989)

97 and Beardsmore and Cull (2001). Paleo-water depth for the time steps was estimated for the model by Burgreen-Chan et al. (2016) after biostratigraphic and paleogeographic work from Field and Uruski (1997).

Because of the high sensitivity of methane hydrates to pressure and temperature conditions

(e.g., Xu and Ruppel, 1999), basin models of methane hydrate are most sensitive to the inputs of basin geometry, basal paleo-heat flow, sediment-water interface temperature, and paleo-water depth, and are above all governed by effects of these inputs on pressure and temperature conditions. Modeling of the gas hydrate stability zone described here uses these base conditions in combination with pre-defined gas hydrate phase properties (i.e., pressure, volume, and temperature conditions) to calculate extent of the gas hydrate stability zone, using the equation of Tishchenko et al. (2005), at each structurally restored model time step. Specific parameters used here are gas hydrate density of 940 kg/m3, thermal conductivity of 0.49 W/m/K, heat capacity of 0.50 kcal/kg/K, and dissociation energy of 54.5 kcal/mol.

A number of assumptions and their attendant uncertainties are embedded within basin modeling (e.g., Hantschel and Kauerauf, 2009). These include: structural components, such as the total amount of shortening (which is especially relevant in considering the magnitude of possible structural influence on the gas hydrate system), as well as spatial and temporal interpolation by the software between structural restorations at discrete times in the basin history (assumptions are detailed in Burgreen-Chan et al., 2016); boundary conditions imposed on the model (heat flow, sediment-water interface temperature, and water depth through time); and the role of faults. For this effort, we did not consider faults to be conduits for gas migration to the GHSZ (Barnes et al., 2010; Crutchley et al., 2010) or catalysts of

98 warm fluid flow and potential gas hydrate destabilization (Suess et al., 1999; Pecher et al.,

2010), nor did we consider the role of folding in hydrate formation (Crutchley et al., 2018).

Keeping the focus of the study simple, we sought to understand the role tectonic uplift may play in destabilizing gas hydrate through geologic time.

Notwithstanding the many assumptions and constraints in model parameters, basin modeling allows us to quantify the lateral extent of the GHSZ along the modeled transect at each structurally restored time step.

3.3. RESULTS

Using this PetroMod™ v.2016 2D restored structural model, we simulated gas hydrate properties through time, focusing on the presence and extent of the GHSZ. We model the extent of the GHSZ since the late Oligocene, through numerous phases of Neogene deformation. We tested two end-member heat flow scenarios (of Burgreen-Chan and

Graham, 2018) and find that results for each are qualitatively similar. As expected, lower heat flow generally yields a slightly thicker GHSZ. Here we describe results from this lower heat flow scenario (Figure 3.3), which yields a better fit to geophysical observations of the bottom-simulating reflector (BSR) (Figure 3.5).

The time step at 24 Ma (Figure 3.3) shows the modeled transect assuming largely undeformed strata before onset of subduction. In the initial model state, gas hydrate is stable along most of the transect (GHSZ is 108 km in extent), with gas hydrate stability occurring from water depths of ~750 m to ~3500 m (Table 3.1). At reference locations at

20, 40, 60, 80, and 100 km along the model transect, GHSZ thicknesses range from ~200 to ~500 m at 20 and 100 km along the transect, respectively (Figure 3.4; Table 3.1).

The 13.6 Ma (Figure 3.3) time step shows faulting and shortening. However, the backstripped model suggests that this tectonic activity, as with the activity at 18.8 Ma

(Figure 3.3), did not result in sufficient uplift and reduction in water depth to significantly affect the GHSZ extent (105 km).

Faulting and shortening continue at 8.5 Ma (Figure 3.3). This tectonism starts to affect the predicted GHSZ, particularly in the region ~28-39 km along the transect. This section of the transect has been uplifted such that the sediment surface along this section is at 500-

100

600 m water depths, resulting in GHSZ destabilization. The modeled GHSZ now extends

89 km. GHSZ thicknesses range from ~100 to ~400 m at 20 and 100 km along the modeled transect (Figure 3.4; Table 3.1).

The structural reconstruction at 7 Ma (Figure 3.3) shows the most extreme effects of faulting and shortening. Uplift here elevated the section of the transect from ~15-61 km to water depths of 500 m or less, and led to complete disappearance of the GHSZ along this section. This time step shows markedly limited presence of the GHSZ in comparison to the

24-8.5 Ma time steps. Here, the predicted GHSZ is almost entirely restricted to deeper water in the range of ~64-107 km along the transect. Extent of the GHSZ has dropped markedly, falling to 54 km. The GHSZ is absent from reference locations at 20, 40, and 60 km along the modeled transect, and is ~200 and ~400 m thick at 80 and 100 km along the transect (Figure 3.4; Table 3.1). This time step represents the climax of degree of uplift and

GHSZ disruption. The time step at 4 Ma (Figure 3.3) is similar to the 7 Ma time step (48 km GHSZ extent).

At present day (0 Ma) (Figure 3.3), our model shows further uplift in the ~70-110 km section of the transect. This resulted in increased disruption of the GHSZ, with two patches of GHSZ persisting in the regions ~73-97 km and ~99-108 km. Extent of the GHSZ has fallen to 33 km, with GHSZ absent in all but the 80 and 100 km reference locations (Figure

3.4; Table 3.1). Generally, there is agreement between location and extent of the base of the predicted (modeled) GHSZ and interpretations of the BSR in seismic (Figure 3.5), though it should be noted that the BSR is highly discontinuous and difficult to identify in parts of the section.

101

3.4. DISCUSSION

Our model shows that, since the Oligocene, tectonic uplift on the Hikurangi margin has directly and significantly reduced the extent of the GHSZ. It should be noted, however, that the GHSZ (but not necessarily entrapped methane) is also pushed seaward of the modeled section to greater water depths. Along the modeled transect, gas hydrates ceased to be stable over a distance of 75 km across the shelf (with GHSZ reduced in extent from

~108 to just ~33 km). This is equivalent to a ~70% loss in area that was potentially occupied by gas hydrates along the modeled transect. The largest degree of gas hydrate destabilization is predicted to have occurred over a relatively short time interval in the late

Miocene between 8.5 and 7 Ma. However, given high sedimentation and related high predicted gas generation and hydrate formation rates in the Plio-Pleistocene (Kroeger et al., 2015), post-Miocene uplift may have quantitatively higher impact on release of methane due to a reduction in the extent of the GHSZ. This is in agreement with large degrees of GHSZ destabilization from 4 Ma onward. Most gas derived from gas hydrate dissolution would ultimately have been released into oceanic and atmospheric systems over time, contributing to and impacting ocean chemistry and the carbon cycle. Importantly, this contribution of carbon and influence on oceanic chemistry would likely be magnified during periods of increased plate convergence globally (e.g., Rea and Duncan, 1986;

Verplanck and Duncan, 1987) and hence could cause a change in the rate of global carbon cycling.

Because the structurally restored line used for our model assumes a minimal amount of shortening (Burgreen-Chan et al., 2016), our model results represent a low shortening end-

102 member. Thus, there was possibly an even greater degree of tectonically-mediated GHSZ disruption along the modeled section.

103

3.5. CONCLUSIONS

We show that tectonic uplift can significantly affect GHSZ evolution. In particular, uplift reduces the areal extent and thickness of the GHSZ. When considering the effects of tectonics on gas hydrate deposits through geologic time, it is critical to incorporate the effect of uplift on the GHSZ along with recently developed understanding of the impact of tectonic factors like faulting (Suess et al., 1999; Barnes et al., 2010; Crutchley et al., 2010;

Pecher et al., 2010) and folding (Crutchley et al., 2011; Crutchley et al., 2018).

Furthermore, we propose that tectonic uplift along continental margins represents a mechanism for long-term (million-year timescales), persistent input of carbon to the carbon cycle. More recent understanding of oceanic sinks, and in particular, disassociation of gas in the water column, suggests gas release from destabilized gas hydrate is oceanically buffered rather than delivered to the atmosphere as a potential contributor to climatic change (Faure et al., 2010; Law et al., 2010; Ruppel and Kessler, 2017). This seems especially likely if gas is gradually released to the water column, as is likely the case for tectonically-mediated destabilization of gas hydrate over geologic time. Nonetheless, we suggest that contribution of tectonic uplift to long-term and significant, albeit regional, carbon release should be considered as part of the global carbon budget, and particularly the oceanic carbon cycle, over geologic timescales. This contribution of carbon and influence on ocean chemistry would likely be magnified during periods of increased global plate convergence and could therefore cause changes in the rate of global carbon cycling.

104

ACKNOWLEDGEMENTS

The authors thank An Yin and an anonymous reviewer for helpful suggestions, which improved the manuscript, and also thank editors Victor Tsai and Christian Huber. An early version of this paper benefited from input and suggestions by Ray Boswell and Gareth

Crutchley. Funding to ZB for this work was provided by the Stanford University Basin and

Petroleum System Modeling Industrial Affiliates Group, the U.S. Department of Energy

National Energy Technology Laboratory, the AAPG Grants-in-Aid Program, and the

Stanford McGee/Levorsen fund. We thank Schlumberger for providing the PetroMod™ license. We thank Best Chaipornkaew, Jared Gooley, and Laura Dafov for assistance with

PetroMod™. Seismic data used in construction of the model discussed herein and in interpretation of the BSR can be accessed through https://data.nzpam.govt.nz. Other model input data and parameters are available at www.aapg.org/datashare as Datashare 91.

AUTHOR’S NOTE

This study (Chapter 3) is published in Geophysical Research Letters as:

Burton, Z.F.M., Kroeger, K.F., Hosford Scheirer, A., Seol, Y., Burgreen‐Chan, B.,

and Graham, S.A., 2020, Tectonic uplift destabilizes subsea gas hydrate: A model

example from Hikurangi margin, New Zealand: Geophysical Research Letters, v.

47, e2020GL087150.

105

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117

TABLES

Table 3.1. Depth and thickness of GHSZ at reference locations along modeled time steps.

118

FIGURES

Figure 3.1. Location of modeled cross section (black line). Inset map shows New Zealand outlined and shaded in dark grey and the Hikurangi margin plate boundary in black, with study area in red box.

119

Figure 3.2. Present-day structural transect. (a) Uninterpreted CM05-01 Hawke’s Bay seismic line. (b) Seismic stratigraphic and structural interpretation from Burgreen-Chan et al. (2016). (c) Present-day 2D basin and petroleum system model indicating predicted

(modeled) GHSZ in red. Note 1.7x vertical exaggeration.

120

Figure 3.3. Nine modeled time steps (24 to 0 Ma). Modeled GHSZ shown in red. Note ~5x vertical exaggeration.

121

Figure 3.3. Continued.

122

Figure 3.4. GHSZ thicknesses (black dashed lines) and water depths (blue polygons) through time for reference locations 20, 40, 60, 80, and 100 km along the transect.

123

Figure 3.5. Comparison of BSR and predicted GHSZ at present day. (a) Seismic line with interpreted bottom-simulating reflector (BSR) indicated by black arrows. (b) Present-day modeled GHSZ shown in red. Note vertical exaggeration.

124

CHAPTER 4

GLOBAL INVENTORY OF BIOGENIC GAS HYDRATE-ASSOCIATED

SEDIMENT ORGANIC PROPERTIES

125

GLOBAL INVENTORY OF BIOGENIC GAS HYDRATE-ASSOCIATED

SEDIMENT ORGANIC PROPERTIES

Zachary F. M. Burton, Allegra Hosford Scheirer, and Kenneth E. Peters

Department of Geological Sciences, Stanford University, Stanford, CA, USA

ABSTRACT

Gas hydrates rival in size all known organic carbon sources and are a significant potential energy resource. Organic matter in shallow sediments is the critical feedstock and prerequisite for microbial metabolism and resultant generation of gas. Understanding of the amounts of organic matter needed for microbial activity is therefore fundamental to understanding the basic requirements for generation and incorporation of biogenic gas into gas hydrate. While the processes of biogenic gas generation are well described, and while tentative modeling work has sought to establish lower organic carbon limits for formation of gas hydrate, no attempt has been made in over two decades to turn to drilling data from global gas hydrate sites to enhance our understanding of the sediment organic properties needed for the formation of gas hydrate. Accordingly, we here compile global shallow (<2 km) sediment organic properties, namely, total organic carbon (TOC) and hydrogen index

(HI), as measurements of the elemental abundances required for the formation of biogenic gas (i.e., the formation of hydrocarbons, e.g., CH4) incorporated into gas hydrate. Across

43 known gas hydrate sites from five continents, average TOC values range from 0.3–3+

126 wt. %, compatible with previous studies suggesting a minimum TOC of 0.3–0.5 wt. % needed for gas hydrate formation. Average TOC for all sites is 1.2 wt. % and median TOC is 1.1 wt. % (n = 2000+ total analyses). HI values range from 50–400+ mg hydrocarbon/g

TOC, suggesting an HI of 50 may result in gas hydrate formation. Average HI for all sites is 170 and median HI is 150 (n = 500+ total analyses). This compilation provides both understanding of average organic properties of hydrate-associated sediment as well as possible empirical evidence of minimum TOC and HI necessary for the formation of gas hydrate. This work is relevant to any future study of gas hydrate systems.

127

4.1. INTRODUCTION

Biogenic gas and gas hydrate represent a significant potential energy source (Collett, 2002;

Makogon et al., 2007; Boswell, 2009; Makogon, 2010; Boswell and Collett, 2011; Chong et al., 2016), as well as a significant fraction of carbon in the context of the global carbon cycle (Kvenvolden, 2002; Nisbet and Chappellaz, 2009). Biogenic gas may comprise over

20% of global gas resources (Rice and Claypool, 1981; Katz, 2011), and gas hydrate may account for more carbon than is present in all other known organic carbon sources, including fossil fuel reserves (Kvenvolden, 1995; Sloan, 2003; Collett et al., 2015).

Biogenic gas and gas hydrate resources are attractive for exploitation not only for the volume of potential resource they represent, but also because natural gas is a cleaner fuel compared to oil and coal (Pacala and Socolow, 2004).

Despite the potential size of the prize represented by these gas sources, exploitation of biogenic gas and gas hydrate has been limited due to the difficulty of resource estimation and production (Katz, 2011; Collett et al., 2015). For shallow gas and gas hydrate, much difficulty lies in predicting the volumes of gas incorporated in gas hydrate accumulations.

Fundamental to successful volumetric assessment of biogenic gas generation and gas hydrate formation is an understanding of sediment organic matter content, which is the critical chemical feedstock for any gas-generating reaction (Rice and Claypool, 1981).

Organic matter supplies both the carbon and hydrogen necessary for the formation of hydrocarbons (e.g., CH4). Furthermore, accumulation of organic matter on the seafloor is considered the primary prerequisite for gas hydrate formation, and a direct indicator of gas hydrate potential (Wallmann et al., 2012; Johnson et al., 2014). Thus, understanding of organic matter content in gas hydrate-associated marine sediment is critical to

128 understanding the controls on gas hydrate formation (e.g., minimum organic requirements), the volumetrics and prospectivity of gas hydrate resources, and carbon budgets.

Nonetheless, understanding of the global distribution of gas hydrate-associated organic matter content is lacking, with the closest attempt at such having been published over 20 years ago on 17 biogenic gas hydrate sites along three continental margins (Waseda, 1998).

With this issue in mind, we here compile a comprehensive global inventory of sediment organic properties – namely, total organic carbon (TOC) and hydrogen index (HI), as representative of carbon and hydrogen content – from 43 sites along five continental margins from which biogenic gas hydrate has been recovered (Figure 4.1). Though we discuss both marine-sediment-hosted and permafrost-associated hydrate, our focus here is primarily on marine gas hydrate as it represents >98% of global gas hydrate (Kvenvolden,

1988; Sloan, 2003; Ruppel, 2014).

129

4.2. BACKGROUND

4.2.1. BIOGENIC GAS GENERATION

Natural gas is generated during the degradation of organic matter via both biogenic and thermogenic processes (Tissot et al., 1974; Rice and Claypool, 1981; Tissot and Welte,

1984). Biogenic gas generation occurs during the degradation of organic matter by living microbes in the relatively shallow subsurface (up to ~1–2 km depth; generally <50°C, excluding some thermophilic microbes), while thermogenic gas generation occurs during the thermal degradation of organic matter at significant depths (generally ~3+ km depth;

90+°C) (Tissot et al., 1974; Zeikus and Winfrey, 1976; Rice and Claypool, 1981; Tissot and Welte, 1984; Katz, 2011). Microbial respiration and resultant biogenic gas generation typically results in CH4 gas, while thermally-induced cracking of carbon-carbon bonds and resultant thermogenic gas generation results in heavier (i.e., C2–5) gases as well as, with increasing depth, CH4, CO2, and carbon-rich residue as the heavier gases themselves are cracked (Tissot et al., 1974; Rice and Claypool, 1981; Tissot and Welte, 1984).

Because most gas incorporated into gas hydrates is considered to be biogenic (e.g.,

Kvenvolden, 1995, 1998; Booth et al., 1996; Kvenvolden and Lorenson, 2001; Sloan,

2003; Koh et al., 2011; Chong et al., 2016), we here focus our discussion on biogenic gas generation. The primary processes of and controls on the biogenic generation of natural gas (predominantly CH4 in marine settings, but including CO2 in nonmarine settings) have been discussed in, among others, Rice and Claypool (1981) and, more recently, Katz

(2011).

130

The succession of low-temperature processes that lead to decomposition and diagenesis of organic matter (of average composition described by the Redfield ratio, i.e., (CH2O)106

(NH3)16 (H3PO4); Redfield, 1958; Geider and La Roche, 2002) and resultant biogenic gas generation in marine sediments begins with consumption of oxygen by benthic organisms during aerobic respiration (i.e., oxidation) of organic matter (which produces CO2) in the water column and uppermost sediment (Claypool and Kaplan, 1974; Froelich et al., 1979;

Rice and Claypool, 1981; Kristensen, 2000). Once all oxygen is consumed, bringing about fully anaerobic conditions, respiration of organic matter by sulfate-reducing microbes becomes the dominant process (e.g., Oremland and Taylor, 1978) and continues until sulfate supply within the sediment pore waters is depleted (Nissenbaum et al., 1972;

Claypool and Kaplan, 1974; Martens and Berner, 1974; Rice and Claypool, 1981; Westrich and Berner, 1984). Once sulfate reduction ceases, methanogens (i.e., CH4-generating microbes, which are exclusively from the domain archaea and include, for example, the orders Methanomicrobiales and Methanosarcinales in marine sediments, e.g., Kendall et al., 2007) are able to thrive, and the generation of methane becomes the prevailing microbial process (Claypool and Kaplan, 1974; Martens and Berner, 1974; Rice and

Claypool, 1981). Methanogens thrive on three major types of substrates: CO2, acetate, and methyl-group containing compounds (Mah et al., 1977; Balch et al., 1979; Liu and

Whitman, 2009). Thus, the only known methanogenic pathways are CO2 reduction (CO2 +

4H2 → CH4 + 2H2O), acetate fermentation (CH3COOH → CH4 + CO2), and methyl group reduction (Whiticar et al., 1986; Liu and Whitman, 2009). Of these, the process of hydrogen-mediated CO2 reduction is cited as the most important mechanism responsible

131 for potential marine gas accumulations (Claypool and Kaplan, 1974; Rice and Claypool,

1981; Whiticar et al., 1986; Katz, 2011).

Following sediment burial (whereby more rapid burial favors preservation of organic matter, e.g., Müller and Suess, 1979), the primary controls on methane generation include, as outlined in Rice and Claypool (1981): an anaerobic (completely anoxic) environment

(e.g., Buswell and Neave, 1930; Barker, 1936; Buswell and Mueller, 1952), a sulfate-poor environment (e.g., Claypool and Kaplan, 1974; Martens and Berner, 1974; Oremland and

Taylor, 1978), relatively low temperatures (generally on the order of 0–75°C, e.g., Buswell and Mueller, 1952; Zeikus and Wolfe, 1972; Zeikus and Winfrey, 1976, though in cases of some hyperthermophiles temperatures can exceed 120°C, e.g., Kurr et al., 1991; Takai et al., 2008), sufficient pore space for microbes to function (e.g., Hassink et al., 1993;

Fredrickson et al., 1997; Phadnis and Santamarina, 2011), and finally – and central to this study – the availability of organic matter (e.g., Söhngen, 1906; Buswell and Neave, 1930).

Katz (2011) notes that additional controls such as pore water salinity (e.g., Waldron et al.,

2007; Zhao et al., 2017) and nutrient availability (e.g., Ergüder et al., 2001; Jones et al.,

2010; Fallgren et al., 2013) also exert influence.

Rice and Claypool (1981) note that the majority of biogenic gas generation likely occurs at burial depths of less than 1,000 meters below the seafloor (mbsf), and may occur over a range as narrow as a few cm to hundreds of m depending on relative rates of biologic processes and sediment accumulation. However, consensus on the peak depth of marine methanogenic activity (i.e., biogenic gas generation) is lacking. Modeling work tuned to sediment core data and laboratory experiments predicts that the reactivity of sedimentary organic matter decreases with increasing depth (i.e., with increasing time) according to a

132 power function (Middelburg, 1989; Wallmann et al., 2006), and that rates of biogenic gas generation depend on this organic matter reactivity and therefore also decrease with depth, with modeled peak rates of biogenic gas generation therefore predicted to occur at quite shallow depths (e.g., 22–28 mbsf in Blake Ridge model scenarios) (Wallmann et al., 2006).

This may be compatible with drilling results from some gas hydrate sites (e.g., offshore

Peru and Hydrate Ridge, Oregon), which suggested abundant methanogenic activity <20 mbsf (Suess et al., 1988; Colwell et al., 2008). As a competing mechanism, laboratory work on anaerobic microbe communities (including those isolated from aquatic sediment) as well as findings from anaerobic waste digestion and treatment suggest that methanogens are most active under mesophilic conditions, with peak activity at an optimum temperature range of 35–45°C (Hungate, 1950; Zeikus and Winfrey, 1976; Kim et al., 2002; Collins et al., 2003; Katz, 2011), such as might be present at relatively deeper (i.e., 1+ km) depths.

Interestingly, despite shallow (<50 mbsf) modeled peak rates of biogenic gas generation, the work by Wallmann et al. (2006) postulates that more than 90% of biogenic methane found in gas hydrate along the Sakhalin slope (Sea of Okhotsk), upon which the organic matter degradation and methane generation model of those authors is based, and at Blake

Ridge is sourced from sediment beneath the gas hydrate stability zone (GHSZ) at depths

>100 m and >300 m, respectively (i.e., sediment beneath the depth of predicted peak biogenic gas generation), and thus, at temperatures greater than the typical <20°C temperatures found within the GHSZ.

4.2.2. GAS HYDRATE FORMATION

Gas hydrates (i.e., gas clathrates) are nonstoichiometric crystalline minerals that occur in the form of a frozen solid comprised of a gas molecule surrounded by a cage or lattice of

133 water molecules (Jeffrey and McMullan, 1967; Kvenvolden, 1988; Sloan, 2003; Sloan and

Koh, 2007). These ice-like solids are a highly concentrated form of natural gas that are stable under specific combinations of elevated pressures and low temperatures

(Kvenvolden, 1988; Sloan, 2003; Sloan and Koh, 2007).

Most gas found in gas hydrate is CH4, though CO2 and C2+ hydrocarbon gases are often present in minor amounts (Kvenvolden, 1995; Milkov, 2005). On the basis of scientific studies of gas hydrate samples recovered from drilling expeditions along continental margins, most recovered samples show that incorporated methane is generated via biogenic rather than thermogenic processes (Kvenvolden, 1995, 1998; Booth et al., 1996;

Kvenvolden and Lorenson, 2001; Sloan, 2003; Koh et al., 2011; Chong et al., 2016).

Provided that a sufficient gas source is present (i.e., an advection-dominated system whereby CH4 flux exceeds some critical value and the amount of CH4 dissolved in liquid exceeds methane solubility; Xu and Ruppel, 1999), the primary control on whether formation of gas hydrate is possible is the prevailing pressure and temperature conditions, whereby relatively high pressures (generally at water depths >300 m for marine gas hydrate) and relatively low temperatures (generally approaching 0°C) are required for stability of gas hydrate (e.g., Hammerschmidt, 1934; Marshall et al., 1964; Jeffrey and

McMullan, 1967; Stoll et al., 1971; Claypool and Kaplan, 1974; Shipley et al., 1979;

Kvenvolden, 1988, 1993; Handa, 1990; Dickens and Quinby-Hunt, 1994; Sloan, 2003;

Sloan and Koh, 2007). Pressure and temperature conditions governing hydrate stability are in turn influenced by climatic and geologic factors including ocean warming (e.g., Mienert et al., 2005; Biastoch et al., 2011; Serov et al., 2017), falling sea level (e.g., Kayen and

Lee, 1991; Lerche and Bagirov, 1998), isostatic rebound (e.g., Wallmann et al., 2018), and

134 tectonic uplift (e.g., Jahren et al., 2005; Maclennan and Jones, 2006; Burton et al., 2020).

It should be noted that additional factors including pore water salinity (e.g., Deaton and

Frost, 1949; Katz et al., 1959; Menten et al., 1981; de Roo et al., 1983; Li and Nghiem,

1986; Handa, 1990; Dickens and Quinby-Hunt, 1997), gas composition (e.g.,

Hammerschmidt, 1934; Katz et al., 1959; Hovland et al., 1995; Sloan and Koh, 2007), and capillarity (e.g., Clennell et al., 1999; Henry et al., 1999; Anderson et al., 2009; Liu and

Flemings, 2011) can affect gas hydrate stability. The depth range over which gas hydrate is predicted to be stable based on appropriate temperature and pressure conditions is referred to as the gas hydrate stability zone (or methane hydrate stability zone; MHSZ), the base of which is sometimes correlated with a bottom-simulating reflector (BSR) seen in marine seismic data which marks the transition from solid overlying gas hydrate to underlying free gas (e.g., Shipley et al., 1979; Yamano et al., 1982), though it is noted that there are numerous scenarios in which the BSR may not correlate with the actual base of the GHSZ (Xu and Ruppel, 1999).

For greater than 98% of global gas hydrate, the requirements for gas hydrate formation, namely an abundant gas source paired with appropriate pressure-temperature conditions, exist in marine deep-water sediments of continental margins, while the remaining ~1–2% of gas hydrate exists in permafrost regions (Kvenvolden, 1988; Sloan, 2003; Ruppel,

2014).

4.2.3. SEDIMENT ORGANIC PROPERTIES, BIOGENIC GAS, AND GAS

HYDRATES

In marine environments, accumulation of organic carbon on the seafloor is generally regarded as the primary prerequisite for the formation of gas hydrates, and a first order

135 indicator for gas hydrate potential (Wallmann et al., 2012; Johnson et al., 2014). Organic carbon content in the context of hydrocarbon generation is quantified through measurement of the mass fraction of total organic carbon (TOC), expressed in terms of percentage by weight (wt. %) (Espitalié et al., 1977; Peters, 1986; Peters and Cassa, 1994; Tyson, 1995).

Previous work suggests a minimum TOC content of ~0.3 wt. % (based on reaction- transport modeling by Burwicz et al., 2011), 0.4 wt. % (based on mass-transfer modeling by Klauda and Sandler, 2005), or 0.5 wt. % (based on numerical calculations by Waseda,

1998) is necessary for gas hydrate formation to occur via hydrate in-filling of sediment pore spaces. Of note is the work by Burwicz et al. (2011), which suggested that gas hydrate formation is controlled by a combination of sediment TOC and sedimentation rate, and that this means the minimum TOC needed for gas hydrate formation ranges from ~0.3 to >1.0 wt. %. These values are consistent with earlier estimations of the minimum TOC needed for biogenic methane generation of about 0.5 wt. % (Claypool and Kaplan, 1974; Rice and

Claypool, 1981).

Methane generation at even lower values of TOC may be possible. Clayton (1992) demonstrated that in normally pressured sediments, free gas can be generated with a TOC as low as 0.12 wt. %, whereas in overpressured sediments, a TOC of 0.2 wt. % can result in free gas generation. Clayton (1992) does however note that a TOC in excess of 0.5 wt.

% is likely still required to generate significant quantities of biogenic gas.

Furthermore, with regard to the concept of some global minimum value of TOC needed for gas hydrate formation, Milkov and Sassen (2003) comment that determining such a threshold is difficult given the wide variability of hydrogen/carbon ratios in sediments that may otherwise contain similar TOC values. This raises a perhaps more important, albeit 136 less frequently cited, consideration in the transformation of sedimentary organic carbon to biogenic gas, namely, the availability of hydrogen in these sediments. Hydrogen has long been known as a limiting factor controlling the generation of oil and gas (e.g., Dow, 1977;

Hunt, 1979; Tissot and Welte, 1984). Perhaps less well known is that hydrogen content in shallow sediments likewise controls the ability of methanogens to produce methane that may ultimately be incorporated into gas hydrate, but this has been known for several decades; for instance, Rice and Claypool (1981) noted that the reduction of CO2 by available hydrogen is the most significant process by which methane is generated.

Similar to the use of TOC to quantify carbon available for hydrocarbon generation, the quantity of available hydrogen is parameterized by the hydrogen index (HI) in units of milligrams of hydrocarbon per gram of TOC (mg HC/g TOC) (Espitalié et al., 1977; Peters,

1986).

Together, TOC and HI control how much gas is generated both biogenically in shallow sediments and thermogenically from petroleum source rocks at depth. Thus, TOC and HI control how much gas is available for the formation of gas hydrates.

137

4.3. METHODS

We reviewed global drilling data to compile an inventory of sedimentary organic properties, namely, TOC and HI. Our global inventory includes TOC and HI data from localities where marine gas hydrate has either been recovered (43 localities), or in areas where the presence of marine gas hydrate is strongly indicated by geochemical, geophysical, or other indicators (9 localities) (Figure 4.1). It should be noted that we have intentionally refrained from compiling TOC and HI data for sites where gas hydrate presence has not been confirmed (via physical sample recovery) or inferred, because such sites could well be characterized by organic contents sufficient for the generation of biogenic gas, but could simply be lacking additional requisite conditions for the formation of gas hydrate (e.g., pressure and temperature conditions). Lastly, for sake of completeness, we also include herein (but do not focus on) discussion of permafrost-associated gas hydrate localities (Figure 4.1).

We obtain TOC and HI data from 6 DSDP legs (9 sites), 4 IODP expeditions (6 sites), 13

ODP legs (25 sites), and various other publications on gas hydrate localities. Where possible, TOC and HI averages are included for sediment intervals above, within, and below the GHSZ. We compile TOC and HI data from the following drill sites (Figure 4.1):

North America

1. Cascadia margin

a. Offshore Vancouver Island, IODP Expedition 311 (sites U1325, 1326,

1327)

b. Offshore Oregon (Cascadia Basin), ODP Leg 146 (site 892) 138

c. Offshore Oregon (Hydrate Ridge), ODP Leg 204 (sites 1244, 1245, 1246,

1247, 1248, 1250, 1251)

d. Offshore California (Eel River Basin), ODP Leg 167 (site 1019)

2. Middle America Trench

a. Offshore Mexico, DSDP Leg 66 (sites 490, 491, 492)

b. Offshore Guatemala, DSDP Leg 67 (site 497), DSDP Leg 84 (sites 568,

570)

c. Offshore Costa Rica, DSDP Leg 84 (site 565), IODP Expedition 344 (site

U1412), ODP Leg 170 (sites 1040, 1041)

3. Gulf of Mexico

a. Garden Banks

b. Green Canyon

c. Orca Basin, DSDP Leg 96 (site 618)

4. Blake Ridge, DSDP Leg 76 (site 533), ODP Leg 164 (sites 994, 995, 997)

South America

1. Offshore Peru, ODP Leg 112 (sites 685, 688), ODP Leg 201 (site 1230)

2. Offshore Chile, ODP Leg 141 (site 859)

3. Offshore NE Brazil, ODP Leg 155 (sites 937, 938, 939)

4. Offshore S Brazil

139

5. South Falkland Basin

Europe

1. W Svalbard margin

2. N Black Sea (Sorokin Trough)

Africa

No TOC/HI data from confirmed or suspected gas hydrate localities

India

1. Mahanadi Basin, IODP Expedition 353 (site U1445), NGHP-01 Expedition

2. Krishna-Godavari Basin, NGHP-01 Expedition

3. Andaman Sea, NGHP-01 Expedition

Asia

1. Okhotsk Sea

2. Japan Sea (Okushiri Ridge), ODP Leg 127 (site 796)

3. Japan Sea (Ulleung Basin)

4. Nankai Trough, ODP Leg 131 (site 808), ODP Leg 190 (site 1178)

5. Taiwan Basin

6. South China Sea (Pearl River Mouth Basin), ODP Leg 184 (site 1146)

Oceania

140

1. Hikurangi margin, IODP Expedition 372A (site U1517)

Antarctica

No TOC/HI data from suspected gas hydrate localities

Permafrost-associated hydrate

1. Alaska North Slope, USA

2. Mackenzie Delta, Canada

3. Qinghai-Tibet Plateau, China

4. Russian arctic

141

4.4. GLOBAL INVENTORY OF SEDIMENT ORGANIC PROPERTIES

4.4.1. NORTH AMERICA

Gas hydrate has been recovered at numerous localities along the Cascadia margin of North

America (Figure 4.2). To the west and offshore of Vancouver Island (Figure 4.2), IODP

Expedition 311 recovered gas hydrate at two sites (Riedel et al., 2006). At site U1326, biogenic gas hydrate was recovered starting at 47 mbsf, while at site U1327, biogenic hydrate was recovered starting at 111 mbsf (Riedel et al., 2016). At site U1325, gas hydrate was inferred via electrical resistivity logs to be present starting at 73 mbsf (Riedel et al.,

2006). The average TOC contents and HI measurements at each of these three sites are given in Table 4.1, and range from a TOC of ~0.5–0.7 wt. % and HI of ~55–80 for sediments overlying sampled or inferred intervals of hydrate, and a TOC of ~0.4–0.8 wt.

% and HI of 60–75 for sediments at depths of recovered or inferred gas hydrate.

Further south along the Cascadia margin, ODP Leg 146 recovered biogenic gas hydrate from Cascadia Basin offshore Oregon (Westbrook et al., 1994) (Figure 4.2). This hydrate was recovered at site 892 between 2 to 19 mbsf, with a BSR documented at 73 mbsf

(Westbrook et al., 1994). Average TOC and HI for this site are recorded in Table 4.1.

Sediments from the inferred GHSZ lying above 73 msbf contain average TOC of 1.4 wt.

% and HI of 135, while underlying sediments contain average TOC of 1.5 wt. % and HI of

175.

Hydrate Ridge (Figure 4.2), also offshore Oregon, is subject of much study centered around gas hydrate. Gas hydrate was sampled extensively at multiple Hydrate Ridge sites (sites

1244, 1245, 1246, 1247, 1248, 1250, 1251) during ODP Leg 204 (Tréhu et al., 2003). At

142 site 1244 gas hydrate was recovered, and the extent of the GHSZ was inferred via Cl and thermal anomalies, as well as resistivity, to occur from 45–124 mbsf (Tréhu et al., 2003).

Using the same proxies, the GHSZ is inferred to occur between 48–131 mbsf at site 1245, from 40–114 mbsf at site 1256, from 45–118 mbsf at site 1247, from 0–115 mbsf at site

1248, from 6–114 mbsf at site 1250, and from 40–200 mbsf at site 1251 (though it should be noted that at site 1251, gas hydrate is inferred to be localized primarily at 90–110 mbsf and 190–200 mbsf, whereas at other sites gas hydrate is inferred to occur throughout the

GHSZ) (Tréhu et al., 2003). Average TOC and HI values for all of these sites are found in

Table 4.1. For these sites, average TOC and HI from sediments within the inferred GHSZ range from 1.0–1.3 wt. % and 110–170, respectively. Sediments overlying the GHSZ contain average TOC contents from 0.9–1.4 wt. % and HI values from 50–200, while sediments underlying the GHSZ contain average TOC of 1.0–1.4 wt. % and HI of 90–155.

Near the southern extent of the Cascadia margin, offshore California, gas hydrate has been recovered in the Eel River Basin (Brooks et al., 1991) (Figure 4.2). TOC and HI data is not available from the site at which hydrate was recovered, so we instead report TOC values

(HI not available) from site 1019 of ODP Leg 167, which is also located within Eel River

Basin, though some 100 km to the north of the sampled gas hydrate locality (Lyle et al.,

1997). This TOC data is listed in Table 4.1, and ranges from 0.5–1.5 wt. % for the ~250 m interval sampled.

Gas hydrate has been recovered at various locations along the Middle America Trench, from offshore Mexico to offshore Costa Rica (Watkins et al., 1981; von Huene et al., 1985;

Kimura et al., 1997; Waseda, 1998; Harris et al., 2013) (Figure 4.2). Offshore Mexico

(Figure 4.2), hydrate (documented as ice releasing abundant gas) was recovered at three

143 sites (sites 490, 491, 492) during DSDP Leg 66 (Watkins et al., 1981; Waseda, 1998). HI data is not available for these sites, but average TOC data is given in Table 4.1, and is equal to 1.7 or 1.8 wt. % for all three sites.

In the Middle America Trench of offshore Guatemala (Figure 4.2), DSDP Leg 67 and 84 recovered biogenic gas hydrate (Waseda, 1998). During DSDP Leg 67, hydrate was recovered at site 497, while during DSDP Leg 84, hydrate was recovered at sites 568

(where drilling stopped above the GHSZ) and 570 (where gas hydrate was recovered between 201–210 mbsf) (Waseda, 1998). Average TOC for these three sites is available in

Table 4.1, though HI was not measured. At DSDP Leg 67 site 497, average TOC for the upper ~400 m of sediment is 2.5 wt. %, while average TOC for the upper 330 and 400 m of sediment of both hydrate sites of DSDP Leg 84 is equal to 1.7 wt. %.

To the south, the Middle America Trench of offshore Costa Rica (Figure 4.2) contains gas hydrate as sampled during DSDP Leg 84 (site 565) and ODP Leg 170 (sites 1040 and

1041), and as inferred to be present via mousse-like sediment, Cl anomalies, and a BSR recorded during IODP Expedition 344 (site U1412) (von Huene et al., 1985; Kimura et al.,

1997; Harris et al., 2013). During DSDP Leg 84, gas hydrate was recovered at site 565 from 285 and 318 mbsf (von Huene et al., 1985), while during ODP Leg 170, gas hydrate at site 1040 was recovered at 257 mbsf and inferred from ~20–360 mbsf and gas hydrate at site 1041 was recovered between 116–338 mbsf and inferred to primarily occur between

100–280 mbsf (Kimura et al., 1997). At IODP Expedition 344 site U1412, a BSR occurs at ~200 mbsf (Harris et al., 2013). Average TOC for all sites is listed in Table 4.1, while average HI is available only for ODP Leg 170 site 1040. Average TOC for sediment within the inferred GHSZ at all sites ranges from 1.0–1.5 wt. %, and average HI for the GHSZ at

144 site 1040 is 300. One site, ODP Leg 170 site 1041, has an average TOC of 1.4 wt. % for sediment overlying the inferred GHSZ, while at the two Leg 170 sites, sediment underlying the inferred GHSZ or zone of hydrate recovery has an average TOC of 0.7 and 1.1 wt. %.

Throughout the northern Gulf of Mexico (Figure 4.2), presence of both near-surface (at or near seafloor) and deeper (up to 500 mbsf) gas hydrate has been confirmed over decades of drilling (e.g., Boswell et al., 2012). Recently, drilling during the Gulf of Mexico Gas

Hydrates Joint Industry Project (JIP) Leg I and Leg II has turned up evidence for deeper hydrates at numerous sites throughout the northern Gulf of Mexico. Unfortunately, associated TOC and HI data is lacking or unpublished. Here, we turn to older studies of at- or near-seafloor hydrates, as well as deeper DSDP data, to obtain sediment organic properties. Biogenic gas hydrate has been recovered from surface samples at Garden Banks and Green Canyon (and has also been inferred or recovered at greater depths here)

(Waseda, 1998; Boswell et al., 2012). Core data in the upper 4 mbsf at Garden Banks yield an average TOC of 0.9 wt. %, while core data in the upper 5 mbsf at Green Canyon yield average TOC of 1.0–1.4 wt. % (Waseda, 1998). At DSDP Leg 96 site 618 in Orca Basin, gas hydrate nodules were recovered from 20–40 mbsf (Bouma et al., 1986). Here, TOC is relatively uniform, ranging from 0.6–1.0 wt. % and averaging approximately 0.6 wt. % in the upper ~93 m of sediment (Bouma et al., 1986). A later study of near-surface sediments yielded average TOC of 0.5–1.5 wt. % and HI of 100–250 for the upper 8 mbsf

(Tribovillard et al., 2008). These data are given in Table 4.1.

Blake Ridge off the southeastern coast of U.S.A. (Figure 4.2) is another site extensively studied for its gas hydrate accumulations. Biogenic gas hydrate has been sampled here during DSDP Leg 76 (site 533) and ODP Leg 164 (sites 994, 995, 997) (Sheridan et al.,

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1983; Paull et al., 1996; Waseda, 1998). During DSDP Leg 76, most evidence for gas hydrate at site 533 occurred between ~150 to 250 mbsf (Sheridan et al., 1983). At this interval, average TOC is 1.1 wt. %, while in the overlying sediments, average TOC is 0.5 wt. %. During ODP Leg 164, highest gas hydrate concentrations across the three sites were found at ~200–450 mbsf (Paull et al., 1996). The average TOC at this interval is 1.3 or 1.4 wt. % at all three sites, with an HI ranging from 110–240, while the average TOC of overlying sediments at this site ranges from 0.7–1.0 wt. % and HI ranges from 85–155.

TOC and HI data for Blake Ridge are available in Table 4.1.

Gas hydrate presence has been inferred in other North American localities, namely, the

Bering Sea (e.g., Scholl and Cooper, 1978; Carlson et al., 1985) and the northeastern

Atlantic margin (e.g., Majorowicz and Osadetz, 2001), however, gas hydrate samples have as of yet not been recovered from these localities (Collett et al., 2015).

4.4.2. SOUTH AMERICA

On the South American Pacific margin of offshore Peru (Figure 4.3), gas hydrate has been recovered during ODP Leg 112 (sites 685 and 688) and 201 (site 1230) (Suess et al., 1988;

D’Hondt et al., 2003). At site 685 of ODP Leg 112, biogenic gas hydrate was recovered at

99 and 164 mbsf and a BSR at ~600 mbsf was observed 4.6 km seaward of the site (Suess et al., 1988). At site 688 of ODP Leg 112, biogenic hydrate was recovered at 141 mbsf and a BSR at ~500 mbsf was observed 2.5 km seaward of the site (Suess et al., 1988). At site

1230 of ODP Leg 201, gas hydrate was recovered at 82 and 148 mbsf, and was inferred to occur from ~70 mbsf to the base of the cored interval at 278 mbsf (D’Hondt et al., 2003).

Average TOC and HI values for these sites are listed in Table 4.2. Average TOC for intervals containing gas hydrate range from 2.3–3.5 wt. % while HI ranges from 330–410.

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For the ~70 m of sediment above the inferred GHSZ at ODP Leg 201 site 1230, average

TOC is 3.3 wt. % and HI is 275, while for the ~200 m of sediment beneath the inferred

GHSZ at ODP Leg 112 site 688, average TOC is 1.7 wt. % and HI is 195.

Offshore Chile (Figure 4.3), presence of biogenic gas hydrate was inferred via chemical composition, wireline log, and a BSR at ~100 mbsf (though gas hydrate is speculated to likely be present below this) at ODP Leg 141 site 859 (Behrmann et al., 1992). Here, the sediment above the BSR contains an average TOC of 0.5 wt. % and an HI of 95, while sediment below the BSR contains a TOC of 0.2 wt. % and an HI of 95. These data are presented in Table 4.2.

On the Atlantic margin of South America, off the coast of northeastern Brazil (Figure 4.3), biogenic gas hydrate has been retrieved from the upper Amazon fan and adjacent continental slope via <6 m-long piston core sampling (Ketzer et al., 2018). No TOC or HI data is available from this sampling campaign, so here we report in Table 4.2 TOC averages from ODP Leg 155 (sites 937, 938, 939), which sampled sediment some 100 km away from, and in deeper water than, the location of retrieved hydrate, though was still located on the upper fan (Flood et al., 1995). Average TOC values are effectively constant at all three sites, regardless of sediment depth. In the first ~50 m of sediment, TOC values are

0.8 or 0.9 wt. %, while in the underlying sediment (up to ~300 mbsf), average TOC values are 0.9 wt. % at all three sites. HI is not reported.

Offshore southern Brazil, in Pelotas Basin (Figure 4.3), gas hydrate has also been recovered via piston core sampling (Miller et al., 2015). Here, the authors report that TOC values are generally less than 1.0 wt. %, while a later study by the authors quote TOC values of

147 shallow sediments retrieved via piston coring as equal to 0.5–0.6 wt. % (Rodrigues et al.,

2019), though HI is not reported.

In South Falkland Basin (Figure 4.3), the presence of gas hydrate is inferred due to the presence of a BSR, though gas hydrate has not been retrieved here (Foschi et al., 2019).

Two shallow samples (<3 mbsf) in the foreland region of the basin yield TOC values of

0.5 and 0.6 wt. %, while HI is not reported.

4.4.3. EUROPE

In Europe, gas hydrate has been retrieved at localities including the western Svalbard margin and the Sorokin Trough of the northern Black Sea (Blinova et al., 2003;

Stadnitskaia et al., 2008; Graves et al., 2017) (Figure 4.4). Unfortunately, neither locality has been sampled for TOC measurement to depths any greater than shallow piston or gravity core capabilities (~5 mbsf), nor is HI reported from either site.

Gas hydrate has been retrieved at two sites – Vestnesa Ridge and the “pockmark site” – along the western Svalbard margin (Graves et al., 2017) (Figure 4.4). A study in shallower water, some 30 km away from the nearest location of retrieved gas hydrate, and at the landward limit of the GHSZ, yielded an average TOC value of 0.6 wt. % from shallow piston and gravity core sampling (Graves et al., 2017). This is recorded in Table 4.3.

In the northern Black Sea (Figure 4.4), biogenic gas hydrate has been retrieved via shallow

(<5 mbsf) sampling at the site of numerous mud volcanoes within the Sorokin Trough

(Blinova et al., 2003; Stadnitskaia et al., 2008). Blinova et al. (2003) report an average

TOC of 2.0 wt. % for the top 5 m of sediment at three mud volcano sites, while Stadnitskaia

148 et al. (2008) report an average TOC of 0.9 wt. % for the top 3 m of sediment at five mud volcano sites. These data are presented in Table 4.3.

Presence of gas hydrate is inferred in additional localities in Europe (discussed in Minshull et al., 2020), including the Iberian and Siberian margins (e.g., Delisle, 2000; Romanovskii et al., 2005; Minshull et al., 2020), however, gas hydrate samples have not been recovered here and TOC and HI data is absent (Collett et al., 2015).

4.4.4. AFRICA

The only location with reported existence of a BSR in the Mediterranean is the Nile Delta

(Merey and Longinos, 2018). However, TOC and HI data from shallow sediments are lacking here, with past efforts instead focusing on the thermogenic petroleum system (e.g.,

Shaaban et al., 2006).

On the deep-water Nigerian margin, gas hydrate has been recovered from pockmarks, and has been detected in shallow (<50 mbsf) sediments (Sultan et al., 2014; Wei et al., 2015).

However, no TOC or HI of shallow sediments is reported from this region.

Generally, the African continental margins remain a frontier with regard to gas hydrate exploration. In addition to the localities mentioned above, inferred localities of gas hydrate occurrence are relatively few (e.g., southwestern African margin as in Ben-Avraham et al.,

2002), and are effectively absent on the eastern African margin (Collett et al., 2015).

4.4.5. INDIA

The 2006 National Gas Hydrate Program (NGHP-01) of India was carried out to assess the presence of gas hydrate accumulations all along the Indian continental margin.

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In the Mahanadi Basin on the eastern margin of India (Figure 4.5), both NGHP-01 and

IODP Expedition 353 inferred that gas hydrate was present, via pressure core analysis, thermal and Cl anomalies, and presence of a BSR during NGHP-01, and via mousse-like sediment texture and thermal and Cl anomalies during IODP Expedition 353 (Johnson et al., 2014; Clemens et al., 2016). At the NGHP-01 site, gas hydrate was inferred to be stable to 220 mbsf, and average TOC for the upper 300 m of sediment was 1.1 wt. %, though HI was not measured (Johnson et al., 2014). At IODP Expedition 353 site U1445, the upper

~660 m of sediment was found to have an average TOC of 1.2 wt. % (Clemens et al., 2016).

These data are included in Table 4.4.

In the Krishna-Godavari Basin of the eastern margin (Figure 4.5), gas hydrate was successfully recovered both during NGHP-01 in 2006 and NGHP-02 in 2015 (Johnson et al., 2014; Boswell et al., 2019). Of all basins studied during NGHP-01, the highest gas hydrate saturations are found within the Krishna-Godavari Basin (Johnson et al., 2014).

For the upper ~180 m of sediment across six boreholes in the basin, average TOC was found to be 1.6 wt. % (Johnson et al., 2014). This is recorded in Table 4.4.

In the Andaman Sea to the east of the Indian continental landmass (Figure 4.5), presence of gas hydrate was inferred via pressure core analysis and thermal and Cl anomalies during

NGHP-01 (Johnson et al., 2014). Here, gas hydrate is inferred to be stable up to 620 mbsf, and the upper ~700 m of sediment contain an average TOC of 0.7 wt. % (Johnson et al.,

2014). These data are found in Table 4.4.

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4.4.6. ASIA

On the eastern and southeastern Asian continental margins, gas hydrate has been recovered in the Okhotsk Sea, the Japan Sea, the notable Nankai Trough, the Pearl River Mouth Basin of South China Sea, and the northwest Borneo Trough (Tamaki et al., 1990; Taira et al.,

1991; Cranston et al., 1994; Wang et al., 2000; Waseda and Uchida, 2004; Kim et al., 2007;

Uchida et al., 2009; Warren et al., 2010; Paganoni et al., 2016) (Figure 4.6).

In the Okhotsk Sea (Figure 4.6), biogenic gas hydrate has been recovered in sediment just

0.3–1.2 mbsf (Cranston et al., 1994). Average TOC measured here from 3-m gravity cores is 1.2–1.7 wt. %, however, HI data was not collected. This is recorded in Table 4.5.

In the Japan Sea, gas hydrate has been recovered both in the north from Okushiri Ridge and in the south from Ulleung Basin (Tamaki et al., 1990; Kim et al., 2007) (Figure 4.6).

ODP Leg 127 (site 796) recovered biogenic gas hydrate at 90 mbsf, and inferred gas hydrate to be present between 0–120 mbsf (Tamaki et al., 1990). Sediment within the inferred GHSZ has an average TOC of 0.9 wt. % and an average HI of 235, while the underlying ~300 m of sediment has an average TOC of 0.4 wt. % and an average HI of 325

(Tamaki et al., 1990). These data are presented in Table 4.5. In Ulleung Basin to the south, biogenic gas hydrate has also been recovered (Ryu et al., 2013). A study of eight piston cores (~8 m maximum depth) throughout the basin yielded an average TOC of 1.8 wt. % and an HI of 145, though this sampling campaign was not associated with the retrieval of gas hydrate (Kim et al., 2007). These data are given in Table 4.5.

Nankai Trough offshore Japan (Figure 4.6) has been a primary focus of gas hydrate studies to date, including the first production test of marine gas hydrate in 2013 (e.g., Yamamoto

151 et al., 2014). Biogenic gas hydrate has been recovered here and documented in numerous efforts, including retrieval during ODP Leg 131, and was inferred to be present (via Cl and temperature anomalies) in abundant quantities during ODP Leg 190 (Taira et al., 1991;

Moore et al., 2001; Waseda and Uchida, 2004; Saito and Suzuki, 2007; Uchida et al., 2009).

At ODP Leg 131 site 808, gas hydrate was recovered between 90–140 mbsf, and the GHSZ was inferred to extend to 190 mbsf (Taira et al., 1991). Sampled sediment lying within the inferred GHSZ has an average TOC of 0.6 wt. %, while the underlying ~1 km of sediment has an average TOC of 0.4 wt. %. At ODP Leg 190 site 1178, initial proceedings inferred gas hydrate to be present (on the basis of Cl and temperature anomalies) from ~90 mbsf to the BSR at 400 mbsf (Moore et al., 2001). The two measurements available above this inferred GHSZ yield an average TOC of 1.3 wt. %, while the sediment within the inferred

GHSZ has an average TOC of 1.1 wt. %. At the same site, a later study inferred that gas hydrate is present from ~120–400 mbsf, with highest concentrations at ~150–200 mbsf, on the basis of Cl anomalies and locally colder temperatures (Saito and Suzuki, 2007).

Sediment above this inferred GHSZ has an average TOC of 0.6 wt. % and HI of 50, sediment within the inferred GHSZ has an average TOC of 0.9 wt. % and an HI of 35, and sediment underlying the inferred GHSZ has an average TOC of 0.7 wt. % and an HI of 25.

Further Nankai Trough work, albeit in somewhat shallower water, retrieved biogenic gas hydrate and found an average TOC of ~0.5 wt. % for the entire ~2 km of drilled sediment

(Waseda and Uchida, 2004; Uchida et al., 2009). All Nankai Trough TOC and, where available, HI data are listed in Table 4.5.

In the Taiwan Basin to the south (Figure 4.6), the presence of gas hydrate has been inferred by numerous studies (e.g., Liu et al., 2006), although to-date no hydrate has been retrieved.

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A study of shallow (<12 mbsf) sediment associated with inferred gas hydrate yields an average TOC of 0.5 wt. % (Lei et al., 2018). This is indicated in Table 4.5.

In the Pearl River Mouth Basin of the South China Sea (Figure 4.6), gas hydrate has been successfully retrieved (e.g., Wu et al., 2011). However, TOC data is not available for the exact sites of retrieved hydrate, which lay at water depths of 1105–1423 m, instead, TOC data reported here are from ODP Leg 184 site 1146 at a water depth of 2091 m (Wang et al., 2000). Here, average TOC for the upper ~50 m of sediment is 0.9 wt. %, while average

TOC for the underlying ~250 m of sediment is 0.3 wt. %. These data are presented in Table

4.5.

It should be noted that gas hydrate has also been successfully retrieved from the northwestern Borneo Trough, however, no TOC or HI data is available from this location and much hydrate is of a presumed thermogenic origin (e.g., Warren et al., 2010; Paganoni et al., 2016).

4.4.7. OCEANIA

On the Hikurangi margin offshore eastern New Zealand, gas hydrate was recently recovered during IODP Expedition 372A (site U1517) (Pecher et al., 2019) (Figure 4.7).

Gas hydrate here is predominantly biogenic (Pecher et al., 2019), though thermogenic contribution to gas hydrate has also been inferred for the region (e.g., Kroeger et al., 2015;

Burton et al., 2018), perhaps due to the presence of widespread regional source rocks

(Burton et al., 2019; Naeher et al., 2019). Gas hydrate presence was inferred at ~135–165 mbsf based on chloride measurements and ~113–146 mbsf based on resistivity anomalies

(Pecher et al., 2019). Sediments including these intervals of inferred GHSZ have an

153 average TOC of approximately 0.5 wt. %, however, HI is not reported (Pecher et al., 2019).

These data are presented in Table 4.6.

Presence of gas hydrate is inferred around the Australian continental margin (e.g., the

Tasman Sea as in Exon et al., 1998 or the North West Shelf as in Paganoni et al., 2019) and other parts of Oceania, but has as of yet not been confirmed via sample retrieval

(Collett et al., 2015).

4.4.8. ANTARCTICA

Presence of gas hydrate is inferred around the Antarctic continental margin (e.g., the South

Shetland margin as in Tinivella et al., 1998 or the Ross Sea as in Geletti and Busetti, 2011), but has not been confirmed via drilling and gas hydrate sample retrieval (Collett et al.,

2015).

4.4.9. PERMAFROST-ASSOCIATED GAS HYDRATE

Permafrost-associated gas hydrate has been confirmed via sample recovery in the North

American arctic (Canada’s Mackenzie River Delta and Alaska’s North Slope) and in the

Qinghai-Tibet Plateau of China, and has been inferred to occur in various parts of the

Russian arctic (Collett, 1993, 2002; Song et al., 2014; Collett et al., 2015) (Figure 4.8).

Interestingly, a predominantly thermogenic gas source has been inferred for all confirmed permafrost-associated gas hydrate sites (Collett, 1993; Dallimore and Collett, 2005; Collett et al., 2011; Lorenson et al., 2011; Lu et al., 2013; Wang et al., 2015; Zuo et al., 2016).

Gas hydrate occurrence was inferred at the Mallik site, located in the Mackenzie Delta region of the Northwest Territories (Figure 4.8), as early as 1972 (Bily and Dick, 1974).

The first confirmed recovery of gas hydrate in the North American arctic occurred here in

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1993 (Dallimore and Collett, 1995), while gas hydrate production tests were conducted at

Mallik in 2002 (Dallimore and Collett, 2005). The permafrost at the Mallik 5L-38 well occurs from 0–650 m below the surface, while a number of layers containing gas hydrate occur from 896–1100 m (Dallimore and Collett, 2005). As mentioned above, analyzed gas hydrate samples suggest a predominantly thermogenic gas source for hydrate occurrences at the Mallik site (Dallimore and Collett, 2005). Because only immature sediment of at most late Oligocene age has been penetrated at the site, the TOC and HI associated with the deeper source rock for Mallik thermogenic gas is unknown (Haberer et al., 2006).

Gas hydrate occurrence (the Eileen gas hydrate deposit) in the Prudhoe Bay and Kuparuk

River oil fields of the Alaska North Slope (Figure 4.8) was confirmed via drilling in 1972 and inferred in dozens of wells via well log analysis (Collett, 1993). Gas hydrate production tests were conducted in 2007 at the Mount Elbert site (Hunter et al., 2011) and in 2012 at the Iġnik Sikumi site (Collett et al., 2015; Boswell et al., 2017). Permafrost occurs to a depth of 536.4 m at the Mount Elbert site and up to 600 m in the region, while the GHSZ at Mount Elbert spans ~200–870 m (Hunter et al., 2011; Lorenson et al., 2011). Permafrost at the Iġnik Sikumi site occurs to a similar depth as at Mount Elbert, as does the base of the GHSZ (Boswell et al., 2017). A predominantly thermogenic source is inferred for gas hydrate occurrences of the North Slope (Collett, 1993; Collett et al., 2011; Lorenson et al.,

2011). However, the rock hosting this gas hydrate is immature (with a vitrinite reflectance of ~0.4%), meaning the gas must have been sourced from deeper source rock (Collett et al., 1990; Collett, 1993).

Gas hydrate was successfully recovered in 2008–2009 during scientific drilling beneath the

Qilian Mountain permafrost of the Qinghai-Tibet Plateau of China (Zhu et al., 2010; Lu et

155 al., 2011) (Figure 4.8). Permafrost is estimated to be just ~95 m thick at the drilling sites

(far thinner than the North Slope, Mackenzie Delta, or the West Siberian Basin of Russia;

Song et al., 2014), with the base of permafrost located at ~110–120 m depth, while gas hydrate occurs beneath the permafrost at depths of 133–369 m (Zhu et al., 2010; Lu et al.,

2011). As with the North American arctic hydrate occurrences, gas hydrate of the Qilian

Mountains is inferred to primarily originate from a thermogenic gas source (Lu et al., 2013;

Wang et al., 2015; Zuo et al., 2016). Similarly, the rock hosting gas hydrate of this region is deemed to be of insufficient thermal maturity to act as a source rock for thermogenic gas

(Lu et al., 2013), thus, some authors speculate that the Upper Triassic Galedesi Formation

(with a TOC of ~1.6 wt. %, type II kerogen, and a vitrinite reflectance of 1.2–3.0%) is a likely source rock candidate (Zuo et al., 2016).

In the Russian arctic, gas hydrates are known to be present in the West Siberian Basin

(Figure 4.8), and are inferred to occur throughout Russian permafrost regions (Makogon,

1972; Cherskiy et al., 1985). Permafrost thickness ranges from a few m to 580 m from the south to the north of the West Siberian Platform, while the top of the GHSZ is estimated to occur at depths of 200–500 m and the base of the GHSZ at depths of 400–1000 m

(Cherskiy et al., 1985). Workers have suggested that >30% of gas produced from the

Messoyakha gas field of the northern West Siberian Basin is directly supplied by the depressurization-induced dissociation of gas hydrate within the field (Makogon, 1981;

Makogon and Omelchenko, 2013), although this has been called into question (Collett and

Ginsburg, 1998). Organic properties for gas-producing sediment or rock are not given for the Messoyakha field (e.g., Makogon and Omelchenko, 2013).

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4.4.10. SUMMARY OF GLOBAL DRILLING DATA FROM GAS HYDRATE

LOCALITIES

Sediment organic properties (i.e., TOC and HI) are compiled for 43 localities from five continents from which gas hydrate has been recovered. Data is collected for hydrate localities of North and South America, Europe, Asia, and Oceania, while data is not available for confirmed gas hydrate localities in Africa or for inferred localities in

Antarctica.

Figure 4.9 and 4.10 provide kernel density estimation functions (e.g., Vermeesch, 2012) and histograms that describe the distribution of TOC data (Figure 4.9) and HI data (Figure

4.10), while Figure 4.11 and 4.12 provide box plots describing the TOC data (Figure 4.11) and HI data (Figure 4.12) from localities where gas hydrate has been recovered. For the 76 average TOC values of sediment intervals at sites of recovered gas hydrate (which represent 2000+ total TOC analyses), the average TOC value is ~1.2±0.6 wt. %, while the median value is 1.1 wt. % (Figure 4.11). The minimum TOC value is 0.3 wt. % while the maximum value is 3.5 wt. %, and the first quartile is ~0.8 wt. % while the third quartile is

1.4 wt. % (Figure 4.11). Outliers are 2.3, 2.5, 2.6, 3.3, and 3.5 wt. % TOC (Figure 4.11).

For the 42 average HI values of sediment intervals at sites of recovered gas hydrate (which represent 500+ total HI analyses), the average HI value is ~170±85, while the median value is 150 (Figure 4.12). The minimum HI value is 50 while the maximum value is 410, and the first quartile is ~115 while the third quartile is ~195 (Figure 4.12). Outliers are 325,

330, 340, 360, and 410 HI (Figure 4.12).

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Figure 4.13 contains box plots displaying average TOC and HI values of sediment known or inferred in the literature to be located below, within, and above the GHSZ. Sites where the position of analyzed sediment relative to the GHSZ is unknown are therefore excluded.

For sediment below the GHSZ, average TOC is ~1.0±0.4 wt. %, while the median TOC is

1.1 wt. % (281 total TOC analyses from 13 sites). The minimum TOC value is 0.4 wt. % while the maximum value is 1.5 wt. %, and the first quartile is 0.7 wt. % while the third quartile is ~1.4 wt. % (Figure 4.13). There are no TOC outliers. Average HI for sediment underlying the GHSZ is ~180±85, while median HI is 150 (62 total HI analyses from 11 sites). The minimum HI value is 90 while the maximum value is 360, and the first quartile is 135 while the third quartile is 195 (Figure 4.13). Outliers are 325 and 360 HI (Figure

4.13).

For sediment within the GHSZ, average TOC is ~1.3±0.7 wt. %, while the median TOC is

1.2 wt. % (500+ total TOC analyses from 24 sites). The minimum TOC value is 0.4 wt. % while the maximum value is 3.5 wt. %, and the first quartile is ~0.9 wt. % while the third quartile is 1.4 wt. % (Figure 4.13). Outliers are 2.3, 2.6, and 3.5 wt. % TOC (Figure 4.13).

Average HI for sediment within the GHSZ is ~185±100, while median HI is 155 (~250 total HI analyses from 18 sites). The minimum HI value is 60 while the maximum value is

410, and the first quartile is ~115 while the third quartile is 255 (Figure 4.13). There are no

HI outliers.

For sediment above the GHSZ, average TOC is ~1.1±0.7 wt. %, while the median TOC is

1.0 wt. % (170+ total TOC analyses from 13 sites). The minimum TOC value is 0.5 wt. % while the maximum value is 3.3 wt. %, and the first quartile is 0.6 wt. % while the third quartile is ~1.4 wt. % (Figure 4.13). There is one outlier at 3.3 wt. % TOC (Figure 4.13). 158

Average HI for sediment overlying the GHSZ is ~130±70, while median HI is 135 (80 total

HI analyses from 11 sites). The minimum HI value is 50 while the maximum value is 275, and the first quartile is 60 while the third quartile is 180 (Figure 4.13). There are no HI outliers.

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4.5. CONCLUSIONS

Sediment organic properties are critical in any effort to understand and quantitatively characterize biogenic shallow gas and gas hydrate systems. Sediment organic matter richness, as defined by carbon and hydrogen content, is a necessary parameter for improved prediction of biogenic gas and gas hydrate volumes, hydrate saturations, resource accumulations, and marine carbon budgets. We report total organic carbon (TOC) and hydrogen index (HI) values from confirmed gas hydrate localities on continental margins across the globe (North America, South America, Europe, Asia, India, and Oceania). Data from confirmed hydrate localities on the African margin and inferred localities on the

Antarctic margin are lacking. Across the five continents for which data are available, average TOC values at or near confirmed gas hydrate localities range from 0.3 to >3 wt. % and average HI values range from 50 to >400. Average global hydrate-associated TOC is

1.2 wt. % (median 1.1 wt. %) and average HI is 170 (median 150). In general, surveyed sites are compatible with previous studies establishing a minimum required TOC of 0.3–

0.5 wt. % for gas hydrate formation, while surveyed sites suggest that HI values as low as

50 may lead to gas hydrate formation.

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ACKNOWLEDGEMENTS

Funding to ZB was provided by the Stanford University Basin and Petroleum System

Modeling Industrial Affiliates Group, the U.S. Department of Energy National Energy

Technology Laboratory, the Gordon Research Conference Carl Storm Fellowship, the

AAPG Foundation Grants-in-Aid Program, the Stanford Levorsen Fund, and the Shell

Foundation Grant Fund. We thank Yongkoo Seol for support and helpful discussion and

Anatoly Aseev for assistance with data visualization. ZB particularly acknowledges the support of Steve Graham.

161

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TABLES

Table 4.1. Surveyed gas hydrate sites associated with the North American continental margin.

North Expeditio Wate Gas Sub- TO n HI n Reference

America n and Site r hydrate botto C

Locality depth recovered m

(mbsl ? depth

) (mbsf)

Cascadia IODP Exp. 2200 Inferred ~0–70 0.7 24 80 2 Riedel et al., margin 311, Site 4 2006

(offshore U1325 ~70– 0.4 45 65 4 Vancouver 300* 5 Island) IODP Exp. 1828 Yes ~0–50 0.5 21 55 2

311, Site 1

U1326 ~70– 0.4 41 60 4

270* 1

1321 Yes ~0– 0.5 25 60 2

110 5

190

IODP Exp. ~110– 0.8 45 75 4

311, Site 300* 5

U1327

Cascadia ODP Leg 675 Yes 4–68* 1.4 9 13 6 Westbrook et margin 146, Site 5 al., 1994

(offshore 892 80– 1.5 10 17 5 Oregon, 167 5 Cascadia

Basin)

Cascadia ODP Leg 890 Yes 3–38 1.3 5 13 2 Tréhu et al., margin 204, Site 5 2003

(offshore 1244 47– 1.3 9 13 4 Oregon, 125* 0 Hydrate 146– 1.3 19 13 7 Ridge) 326 5

ODP Leg 870 Yes 4–51 0.9 1 18 1

204, Site 0

1245 51– 1.0 3 17 2

130* 0

140– 1.1 19 15 9

540 5

191

ODP Leg 850 Yes 4–37 1.3 5 50 2

204, Site 46– 1.1 8 11 3 1246 114* 0

122– 1.4 3 15 2

134 0

ODP Leg 835 Yes 3–38 1.2 5 10 1

204, Site 0

1247 47– 1.1 9 15 2

117* 5

138– 1.1 7 12 4

215 5

ODP Leg 830 Yes 1– 1.1 9 11 2

204, Site 111* 5

1248 121– 1.0 4 15 4

148 0

ODP Leg 792 Yes 7– 1.3 10 15 3

204, Site 112* 0

1250 121– 1.3 3 14 2

145 5

192

ODP Leg 1210 Yes 0–32 1.4 3 20 1

204, Site 0

1251 41– 1.3 12 14 2

186* 0

207– 1.4 23 90 4

440

Cascadia ODP Leg 980 GH 100 km ~0– 0.5– 79 – – Lyle et al., margin 167, Site to S, within 250 1.5 1997

(offshore 1019 basin

California,

Eel River

Basin)

Middle DSDP Leg 1761 Yes ~40– 1.8 30 – – Waseda,

America 66, Site 610 1998

Trench 490

(offshore DSDP Leg 2883 Yes 7–435 1.7 5 – – Watkins et Mexico) 66, Site al., 1981

491

DSDP Leg 1935 Yes 9–210 1.8 4 – –

66, Site

492

193

Middle DSDP Leg 2347 Yes ~0– 2.5 36 – – Waseda,

America 67, Site 400 1998

Trench 497

(offshore DSDP Leg 2010 Yes ~0– 1.7 16 – – Guatemala 84, Site 400 ) 568

DSDP Leg 1698 Yes ~0– 1.7 13 – –

84, Site 330

570

Middle IODP Exp. 1876 Inferred 3– 1.4 58 – – Harris et al.,

America 344, Site 195* 2013

Trench U1412

(offshore ODP Leg 4177 Yes ~0– 1.0 77 30 2 Kimura et al., Costa 170, Site 360* 0 8 1997 Rica) 1040 ~360– 0.7 40 36 3

653 0

ODP Leg 3306 Yes ~0– 1.4 32 – –

170, Site 100

1041 ~100– 1.5 34 – –

280*

194

~280– 1.1 16 – –

380

DSDP Leg 3099 Yes 0–286 1.0– 11 – – von Huene et

84, Site 1.1 al., 1985

565

Gulf of – 850 Yes 0–4 0.9 8 – – Waseda,

Mexico 1998

(Garden

Banks)

Gulf of – 800– Yes 0–5 1.0– 16 – –

Mexico 880 1.4

(Green

Canyon)

Gulf of DSDP Leg 2412 Yes 0–93* 0.6– 16 – – Bouma et al.,

Mexico 96, Site 1.0 1986

(Orca 618

Basin) – 2240 0–8 0.5– – 10 – Tribovillard

1.5 0– et al., 2008

195

Blake ODP Leg 2798 Yes ~0– 0.8 18 14 8 Paull et al.,

Ridge 164, Site 200 0 1996

994 ~200– 1.4 13 15 1

450* 5 1

ODP Leg 2779 Yes 0–119 0.7 – 15 4

164, Site 5

995 119– 1.3 – 24 3

686* 0 0

ODP Leg 2770 Yes ~0– 1.0 17 85 9

164, Site 200

997 ~200– 1.4 11 11 8

450* 0

DSDP Leg 3191 Yes ~0– 0.5 31 – – Sheridan et

76, Site 155 al., 1983;

533 Waseda, ~155– 1.1 45 – – 1998 400*

*indicates, where possible, the complete depth interval of the GHSZ; in some cases, indicates the interval over which gas hydrate was recovered, or where most evidence for

GH exists

196

Table 4.2. Surveyed gas hydrate sites associated with the South American continental margin.

South Expeditio Wate Gas Sub- TO n HI n Reference

America n and Site r hydrate botto C

Locality depth recovered m

(mbsl ? depth

) (mbsf)

Offshore ODP Leg 5071 Yes 3–451 2.3 15 33 1 Suess et al.,

Peru 112, Site 0 5 1988

685

ODP Leg 3820 Yes 6– 3.5 17 41 1

112, Site 452* 0 7

688 521– 0.7 8 19 8

750 5

ODP Leg 5086 Yes 0–68 3.3 6 27 6 D’Hondt et

201, Site 5 al., 2003

1230 85– 2.6 18 34 1

252* 0 8

Offshore 2741 Inferred 0–90 0.5 19 95 1 Behrmann et

Chile 9 al., 1992

197

ODP Leg 107– 0.2 28 95 2

141, Site 468 8

859

Offshore ODP Leg 2760 GH >100 ~0–50 0.8 15 – – Flood et al.,

NE Brazil 155, Site km away, 1995 ~50– 0.9 24 – – (upper 937 in 180 Amazon shallower ODP Leg 2804 ~0–50 0.8 10 – – fan) water 155, Site ~50– 0.9 49 – – 938 300

ODP Leg 2784 ~0–50 0.9 28 – –

155, Site ~50– 0.9 5 – – 939 90

Offshore S – 500– Yes 0–6 0.5– – – – Miller et al.,

Brazil 600 (piston 0.6 2015;

(Pelotas and core) Rodrigues et

Basin) 1000– al., 2019

1500

South – – Inferred <3 0.5 2 – – Foschi et al.,

Falkland and 2019

Basin 0.6

198

*indicates, where possible, the complete depth interval of the GHSZ; in some cases, indicates the interval over which gas hydrate was recovered, or where most evidence for

GH exists

199

Table 4.3. Surveyed gas hydrate sites associated with the European continental margin.

Europe Expeditio Wate Gas Sub- TO n HI n Reference

Locality n and Site r hydrate botto C

depth recovered m

(mbsl ? depth

) (mbsf)

W – 400 GH 30 km Shallo 0.6 – – – Graves et al.,

Svalbard away, in w 2017 margin deeper piston/

water gravity

core

N Black – 2050– Yes 0–5 2.0 30 – – Blinova et

Sea 2110 al., 2003

(Sorokin – 1800– Yes 0–3 0.9 48 – – Stadnitskaia Trough) 1900 et al., 2008

200

Table 4.4. Surveyed gas hydrate sites associated with the Indian continental margin.

India Expeditio Wate Gas Sub- TO n HI n Reference

Locality n and Site r hydrate botto C

depth recovered m

(mbsl ? depth

) (mbsf)

Mahanadi IODP Exp. 2502 Inferred 1–661 1.2 69 – – Clemens et

Basin 353, Site al., 2016

U1445

NGHP-01 1422 Inferred ~0– 1.1 18 – – Johnson et

Exp. 300* 8 al., 2014

Krishna- NGHP-01 895– Yes ~0– 1.6 35 – –

Godavari Exp. 1253 180 0

Basin

Andaman NGHP-01 1344 Inferred ~0– 0.7 41 – –

Sea Exp. 700* 8

*indicates, where possible, the complete depth interval of the GHSZ; in some cases, indicates the interval over which gas hydrate was recovered, or where most evidence for

GH exists

201

Table 4.5. Surveyed gas hydrate sites associated with the Asian continental margin.

Asia Expeditio Wate Gas Sub- TO n HI n Reference

Locality n and Site r hydrate botto C

depth recovered m

(mbsl ? depth

) (mbsf)

Okhotsk – 710 Yes 3 1.2– 14 – – Cranston et

Sea (gravit 1.7 al., 1994

y core)

Japan Sea ODP Leg 2571 Yes 2– 0.9 7 23 7 Tamaki et

(Okushiri 127, Site 111* 5 al., 1990

Ridge) 796 128– 0.4 14 32 1

430 5 4

Japan Sea – variou Data from <8 1.8 – 14 – Kim et al.,

(Ulleung s throughout (piston 5 2007

Basin) basin in core)

which GH

was

recovered

Nankai 4676 Yes 3– 0.6 15 – – Taira et al.,

Trough 191* 1991

202

ODP Leg 202– 0.4 11 – –

131, Site 1280 5

808

ODP Leg 1742 Inferred 47 and 1.3 2 – – Moore et al.,

190, Site 91 2001

1178 91– 1.1 10 – –

366*

14–75 0.6 3 50 3 Saito and

Suzuki, 2007 130– 0.9 9 35 9

366*

518– 0.7 4 25 4

645

– 945 Yes 0– 0.5 – – – Waseda and

2000+ Uchida,

2004; Uchida

et al., 2009

Taiwan – variou Inferred 0–12 0.5 10 – – Lei et al.,

Basin s 2018

2091 ~0–50 0.9 6 – –

203

South ODP Leg GH ~50– 0.3 16 – – Wang et al.,

China Sea 184, Site recovered 300 2000

(Pearl 1146 at 1105–

River 1423 mbsl

Mouth

Basin)

*indicates, where possible, the complete depth interval of the GHSZ; in some cases, indicates the interval over which gas hydrate was recovered, or where most evidence for

GH exists

204

Table 4.6. Surveyed gas hydrate sites from the Oceania region.

Oceania Expeditio Wate Gas Sub- TO n HI n Reference

Locality n and Site r hydrate botto C

depth recovered m

(mbsl ? depth

) (mbsf)

New IODP Exp. 723 Yes 0–190 0.5 99 – – Pecher et al.,

Zealand 372A, Site 2019

(Hikurangi U1517 margin)

205

FIGURES

Figure 4.1. Global sites from which TOC and HI data were compiled. Red square indicates marine site from which gas hydrate was successfully recovered; white square indicates marine location where presence of hydrate is inferred; yellow circle indicates permafrost- associated site from which hydrate was recovered; white circle indicates permafrost site where presence of hydrate is strongly indicated. Note: some symbols represent multiple sites.

206

Figure 4.2. North American sites from which TOC and HI data were compiled. Red square indicates marine site from which gas hydrate was successfully recovered; white square indicates marine location where presence of hydrate is inferred. Note: some symbols represent multiple sites. CV = Cascadia margin (offshore Vancouver Island), CO =

Cascadia margin (offshore Oregon: Cascadia Basin and Hydrate Ridge), CE = Cascadia margin (offshore California, Eel River Basin), MM = Middle America Trench (offshore

Mexico), MG = Middle America Trench (offshore Guatemala), MC = Middle America

Trench (offshore Costa Rica), GM = Gulf of Mexico, BR = Blake Ridge.

207

Figure 4.3. South American sites from which TOC and HI data were compiled. Red square indicates marine site from which gas hydrate was successfully recovered; white square indicates marine location where presence of hydrate is inferred. Note: some symbols represent multiple sites. OP = offshore Peru, OC = offshore Chile, BA = offshore northeastern Brazil (upper Amazon fan), BP = offshore southern Brazil (Pelotas Basin),

FB = South Falkland Basin.

208

Figure 4.4. European sites from which TOC data were compiled. Red square indicates marine site from which gas hydrate was successfully recovered. Note: some symbols represent multiple sites. SM = western Svalbard margin, BS = northern Black Sea (Sorokin

Trough).

209

Figure 4.5. Indian sites from which TOC data were compiled. Red square indicates marine site from which gas hydrate was successfully recovered; white square indicates marine location where presence of hydrate is inferred. Note: some symbols represent multiple sites. MB = Mahanadi Basin, KG = Krishna-Godavari Basin, AS = Andaman Sea.

210

Figure 4.6. Asian sites from which TOC and HI data were compiled. Red square indicates marine site from which gas hydrate was successfully recovered; white square indicates marine location where presence of hydrate is inferred. Note: some symbols represent multiple sites. OS = Okhotsk Sea, JO = Japan Sea (Okushiri Ridge), JU = Japan Sea

(Ulleung Basin), NT = Nankai Trough, TB = Taiwan Basin, SC = South China Sea (Pearl

River Mouth Basin).

211

Figure 4.7. New Zealand site from which TOC data were compiled. Red square indicates marine site from which gas hydrate was successfully recovered. NZ = New Zealand

(Hikurangi margin).

212

Figure 4.8. Permafrost-associated gas hydrate sites. Yellow circle indicates permafrost- associated site from which hydrate was recovered; white circle indicates permafrost site where presence of hydrate is strongly indicated. Note: some symbols represent multiple sites. NS = North Slope (Alaska), MD = Mackenzie River Delta (Northwest Territories,

Canada), QM = Qilian Mountain (Qinghai-Tibet Plateau, China), WS = West Siberian

Basin.

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Figure 4.9. Kernel density estimation function (black line) and histogram (black rectangles) describing the distribution of average TOC values from intervals of sediment

(e.g., sediment below, within, and above the GHSZ) associated with sites where gas hydrate has been recovered.

214

Figure 4.10. Kernel density estimation function (black line) and histogram (black rectangles) describing the distribution of average HI values from intervals of sediment

(e.g., sediment below, within, and above the GHSZ) associated with sites where gas hydrate has been recovered.

215

Figure 4.11. Box plot displaying the distribution of average TOC values from intervals of sediment (e.g., sediment below, within, and above the GHSZ) associated with sites where gas hydrate has been recovered.

216

Figure 4.12. Box plot displaying the distribution of average HI values from intervals of sediment (e.g., sediment below, within, and above the GHSZ) associated with sites where gas hydrate has been recovered.

217

Figure 4.13. Box plots of average TOC and HI of sediment below, within, and above the

GHSZ. n is the number of sites (with hydrate recovered) for which values are available.

218

CHAPTER 5

GLOBALLY DISTRIBUTED PALEOGENE DEEP-WATER SAND-RICH SYSTEMS

DESPITE HIGH SEA LEVEL DURING WARM CLIMATES

219

GLOBALLY DISTRIBUTED PALEOGENE DEEP-WATER SAND-RICH

SYSTEMS DESPITE HIGH SEA LEVEL DURING WARM CLIMATES

Zachary F. M. Burton1, Tim McHargue1, Tyler Kukla1, Christopher H. Kremer2,

Roger B. Bloch3, Jared T. Gooley1, Chayawan Jaikla1, Jake Harrington4, and

Stephan A. Graham1

1Department of Geological Sciences, Stanford University, Stanford, CA, USA

2Dept. of Earth, Environmental & Planetary Sci., Brown University, Providence, RI, USA

3PO Box 2388, New London, NH, USA

4Matador Resources, Dallas, TX, USA

ABSTRACT

It is widely accepted that coarse deep-marine siliciclastic deposition is controlled by the balance between sediment supply and shelf accommodation. The prevailing paradigm suggests this balance is disrupted and deposition is promoted during periods of falling sea level when shelf accommodation is significantly reduced, and conversely, deposition is inhibited during high sea level when accommodation is maximized. Due to high global temperatures and minimal polar ice, the early Eocene was a time of exceptionally high eustatic sea level, presenting an opportunity to test this paradigm. Here, we compile an inventory of documented early Paleogene sand-rich deep-sea systems. If sea-level fall is indeed a prerequisite for major deep-water turbidite deposition, lower Eocene strata should 220 be characterized by a paucity of sandy turbidites (except on tectonically active margins).

We find 59 instances of early Eocene turbidite systems from all continental margins except

Antarctica. Sand-rich systems were widespread on active margins (42 instances), but were also present on passive margins (17 instances). Along passive margins, 13 of 17 (~76%) turbidite systems are associated with significant Eocene-age fluvial systems. We find that sediment from active margins and sediment delivered to passive margins by integrated drainages and elevated hothouse denudation cause sediment supply to exceed shelf accommodation despite globally-elevated sea level. We therefore suggest that—in addition to the paradigm of eustasy—tectonism and climate-driven change (e.g., drainage integration, global rainfall, denudation) may significantly influence global distribution of coarse-grained deep-marine depositional systems, and may overwhelm eustatic control entirely, especially during periods of high sea level.

221

5.1. INTRODUCTION

5.1.1. THE PARADIGM OF DEEP-SEA SILICICLASTIC DEPOSITION

Sequence stratigraphic models of coarse-grained siliciclastic sedimentation require that sediment volume exceed shelf accommodation for shoreline progradation and voluminous deep-sea deposition to occur (Jervey, 1988). Falling eustatic or relative sea level with resultant decreases in accommodation is the prevailing paradigm for delivering coarse clastics to deep-water passive margin settings (Vail et al., 1977; Catuneanu et al., 2011).

Nevertheless, coarse-grained deep-marine deposition may occur at a local level via: (1) highstand turbidite deposition along narrow shelves of tectonically active margins (e.g.,

Covault et al., 2007; Covault and Graham, 2010; Bernhardt et al., 2015; Zhang et al., 2017);

(2) turbidite deposition during any sea-level cycle if shelf-penetrating canyons intersect the shoreline (e.g., Covault et al., 2007; Covault and Graham, 2010; Bernhardt et al., 2015;

Sweet and Blum, 2016); (3) episodes of exceptionally voluminous sediment supply (e.g.,

Burgess and Hovius, 1998; Carvajal and Steel, 2006; Zhang et al., 2017); and (4) other factors such as turbidite initiation by shelf undercurrents, longshore drift, orbital forcing, subglacial meltwater, and monsoons (e.g., Covault and Graham, 2010; Cantalejo and

Pickering, 2015; Bernhardt et al., 2016). Despite such exceptions, and despite debate on applicability of sequence stratigraphy in general (e.g., discussed by Madof et al., 2019), the prevailing paradigm predicts that during high sea level, one would expect minimal development of sand-rich deep-water systems along passive margins (because resultant high accommodation would restrict most deposition to shelves). To test this hypothesis, we look to one of the most notable periods of elevated eustatic sea level of the Cenozoic: the early Eocene interval (Miller et al., 2005, 2020).

222

5.1.2. THE EARLY EOCENE HOTHOUSE

The early Eocene (~56–48 Ma) was the warmest extended climate interval of the past 65 million years (Zachos et al., 2001). It was characterized by largely ice-free conditions, eustatic sea level at least 70–100 m above present, dampened, significantly lower- amplitude fluctuations in sea level (e.g., versus Quaternary glacial-interglacial fluctuations) (Miller et al., 2005, 2020), shallow equator-to-pole temperature gradients

(Greenwood and Wing, 1995; Wolfe, 1995), and modifications to ocean (Weber and

Thomas, 2017) and atmospheric circulation and hydroclimate (Carmichael et al., 2018).

Consequently, the early Eocene would have been an especially unfavorable time for deep- sea sand-rich turbidite deposition on passive margins if falling sea level is a prerequisite.

However, it has been proposed that this period was also characterized by global-mean rainfall (a dominant control on marine sediment supply) that was higher than today and by more frequent and episodic high-intensity rainfall events (Held and Soden, 2006;

Carmichael et al., 2018). These rainfall changes, along with elevated CO2 levels, contributed to increased chemical weathering (Robert and Kennett, 1997; Ehrmann, 1998;

Smith et al., 2008) and physical weathering and denudation (Smith et al., 2008; John et al.,

2012).

Thus, while early Eocene warm climate and high sea level imply particularly unfavorable conditions for passive margin sand-rich deposition under the prevailing paradigm, other climate drivers such as precipitation changes may have provided a competing control on deposition. Our compilation reveals abundant examples of both active and passive margin sand-rich early Eocene deep-sea systems. We present a conceptual model linking global climate and passive margin sand-rich turbidite systems, and suggest tectonism and climatic

223 conditions such as precipitation and integrated drainages provide the means to overwhelm eustasy on a global scale and result in abundant sand-rich deep-marine deposition despite exceptionally high sea level.

224

5.2. METHODS

We reviewed existing literature to compile published examples of deep-marine sediment deposits (especially submarine fans, sand-rich turbidite systems) of Paleocene, early

Eocene, and middle Eocene age from 114 locations (9 in Africa, 0 in Antarctica, 15 in Asia,

14 in Oceania, 17 in Europe, 6 in India, 34 in North and Central America, 19 in South

America) (Figures 5.1 and 5.2).

Most compiled examples come from onshore outcrop and offshore well and seismic data.

It must be noted that this synthesis inherently assumes the accuracy of the surveyed work, including age dating methods (e.g., geochronology, biostratigraphy), interpreted sedimentology and lithology, and environmental interpretations.

Because our compilation relies on work that is published, it represents a lower limit of the actual number of global turbidite systems of this age. We also acknowledge that we likely missed examples published in non-English language sources. In addition, there are likely numerous sand-rich systems (e.g., reservoirs encountered during petroleum exploration) that have not been reported in the literature, as well as those not preserved in the rock record (e.g., due to subduction or erosion).

Detailed methods are given in Appendix C.

225

5.3. ABUNDANT PALEOGENE HOTHOUSE TURBIDITE SYSTEMS

Despite a prevailing early Eocene eustatic high, our survey identified 59 locations where early Eocene deep-sea sedimentation included sand-rich turbidite systems (Figures 5.1 and

5.2B). These locations span all continents except Antarctica. Most systems are in active margin settings (42 examples), but 17 examples of sand-rich passive margin turbidite systems were also identified (Figures 5.1 and 5.2B). There are slightly more early Eocene examples than Paleocene (28 on active margins, 14 on passive margins; Figures 5.1 and

5.2A) or middle Eocene (39 on active margins, 16 on passive margins; Figures 5.1 and

5.2C) examples, although the proportion of passive vs. active margin turbidite systems is similar. These deep-marine depositional systems are described in detail in Appendix C.

Sand-rich lower Eocene deep-sea deposits are found in active margin settings in Asia,

Europe, India, and North and South America (Figures 5.1C and 5.2B). In Europe and Asia,

Alpine-Zagros-Himalayan collisional tectonism resulted in widespread and abundant foreland basin turbidite deposition. Similarly, the Pyrenean orogeny of Europe, Andean orogeny of South America, and Cordilleran orogenesis of North America contributed to turbidite deposition. Deposition also occurred during periods of rifting (e.g., southeast

Borneo; the Tasmanian shelf; west Ireland) and in borderlands settings (e.g., southern

California; Ireland’s Porcupine Basin).

In both active and passive margin settings, deep-marine siliciclastic systems are often associated with noteworthy early Eocene fluvial systems. Of the 59 locations with record of early Eocene turbidite systems, 19 (~32%) locations are explicitly documented as being associated with substantial fluvial systems. Of these examples, six are from active margin

226 settings. Notably, 13 of 17 (~76%) locations in passive margin settings are associated with integrated ancestral fluvial systems.

It should be noted that some workers cite possible influence of local tectonism (e.g., hinterland uplift, salt movement) in five of the 17 passive margin locations (Tanzania

Coastal Basin, Gulf of Mexico, and three Brazilian basins, as detailed in Appendix C).

227

5.4. DISCUSSION

5.4.1. THE PARADIGM OF DEEP-SEA SILICICLASTIC DEPOSITION

Widespread sand-rich early Eocene turbidite systems along both active and passive margins indicate that the falling sea-level paradigm (e.g., Vail et al., 1977; Catuneanu et al., 2011) need not be invoked for sediment supply to overwhelm shelf accommodation at a global scale. In other words, prevalence of such systems during a eustatic high indicates that the tenet of this paradigm whereby deep-sea deposition is inhibited by maximized accommodation is not always applicable.

5.4.2. MECHANISMS OF SANDY TURBIDITE DEPOSITION DURING EUSTATIC

HIGHS

Alongside widespread evidence for active margin deposition during a eustatic high (e.g., as in Quaternary-age examples from California and Chile; Covault et al., 2007; Bernhardt et al., 2015), distribution of early Eocene turbidite systems suggests that in passive margin settings during eustatic highs, an integrated drainage system may be a prerequisite for substantial sand-rich deep-sea deposition (e.g., Sømme et al., 2009). Activity of such drainages may have been magnified by intense precipitation events in the early Eocene

(Held and Soden, 2006; Carmichael et al., 2018), while sediment transport along drainages may have been magnified by intensified weathering (Robert and Kennett, 1997; Ehrmann,

1998; Smith et al., 2008; John et al., 2012), both of which would act to increase sediment supply. Increased sediment supply as a means to overwhelm shelf accommodation even during a highstand is consistent with work on high-supply and supply-dominated systems, whereby significant deep-marine turbidite deposition occurs despite eustatic highs (e.g.,

228 modern river estimates by Burgess and Hovius, 1998; Maastrichtian shelf of Wyoming in

Carvajal and Steel, 2006; modeling work by Zhang et al., 2017).

5.4.3. CONCEPTUAL MODEL LINKING GLOBAL CLIMATE AND SAND-RICH

TURBIDITE SYSTEMS

Based on our results, we present a conceptual model linking global climate to sandy turbidite occurrence via climatic controls on sediment supply and accommodation (Figures

5.3 and 5.4). Our model is compatible both with the falling sea-level paradigm for deposition (Figure 5.4A) and with deposition during eustatic highs (Figure 5.4B). For simplicity, Figure 5.3 shows only relative changes. We assume sediment supply is a linear function of precipitation intensity (a conceptually useful but oversimplified assumption, e.g., McInerney and Wing, 2011), accommodation is a function of sea level, and both precipitation intensity and sea level can be linked directly to global climate. We justify these assumptions below.

As the planet warms, precipitation intensifies and becomes increasingly episodic (Held and

Soden, 2006), resulting in a direct relationship between temperature and sediment supply.

Starting from a cold climate, warming leads to a relatively rapid increase in eustatic sea level (thus, in accommodation) due to ice sheet melt. However, after loss of large continental glaciers, additional warming leads to smaller thermal-expansion-driven increases in sea level (Li et al., 2018). Turbidite deposition occurs in our model when, for a given climate state, sediment supply (orange line; Figure 5.3) exceeds accommodation

(blue curve; Figure 5.3).

229

Numerous factors vary the slope and intercept of the supply and accommodation lines without violating the basic conceptual framework. For example, greater drainage size

(capturing more sediment) or drainages with more active canyon incision can increase the y-intercept of the sediment supply line. The slope of the supply line may decrease in very hot climates because precipitation is theorized to decline with extreme warming (Schneider et al., 2010). Meanwhile, the accommodation curve is sensitive to factors that modify relationships among sea level, ice sheet, and climate. If continental water storage can drive high-amplitude sea-level variability without ice sheets (Li et al., 2018), this will modify the slope of the accommodation curve at warmer temperatures. The inflection point of the accommodation curve may vary depending on whether the planet is warming or cooling because the ice-albedo feedback can cause differences in the climate state of glaciation versus deglaciation (North et al., 1981; Dortmans et al., 2019). Taken together, we expect that the shape and relative locations of the supply and accommodation curves in our model vary over space and time. Therefore, under our framework it is possible to account for future constraints on climate, erosion, and accommodation dynamics when inferring turbidite occurrence.

5.4.4. SUGGESTIONS FOR FUTURE RESEARCH

Our study invites investigation related to both deep-sea turbidite deposition during eustatic highs and the model that we propose for linking climate and sand-rich systems. For instance, the question arises of whether we can indeed use passive margin turbidite occurrence to predict the nature of sediment supply and the potential for (or necessity of) drainage integration. Our results and model invite turbidite system compilations for other hothouse and warm periods, as in the Cretaceous (Norris et al., 2002) or Paleozoic (Frakes

230 et al., 1992), to further query the potential for widespread sandy, deep-marine systems during eustatic highs, as well as compilations for periods of falling or low sea level (e.g., cooling or cold climates) to further test our conceptual model.

231

5.5. CONCLUSIONS

The early Eocene was a time of elevated temperatures, and represents a high global sea- level endmember relative to present day and most of the Cenozoic. Thus, early Eocene hothouse conditions and the associated eustatic high provide an opportunity to test the hypothesis that eustasy (as a control on shelf accommodation) is the dominant mechanism for sand-rich passive margin deep-sea turbidite deposition on a global scale. Accordingly, our data show that, in considering the balance between sediment supply and accommodation as the control on deep-water siliciclastic deposition, sediment supply is at least as important as global eustasy during high sea level. This is consistent with high- supply or supply-dominated turbidite deposition documented by previous authors. We suggest that, while tectonics can overwhelm eustasy on active margins, both large integrated fluvial drainages on a local to regional scale and intensification of the hydrological cycle on a global scale can overcome the influence of global eustasy and deliver sufficient coarse-grained sediment to passive margins to overwhelm shelf accommodation and result in widespread sand-rich deep-marine turbidite systems despite a eustatic high. Therefore, we suggest that, in considering whether turbidite systems should be globally abundant, likelihood of deposition must be thought of in terms of active margin tectonics, passive margin activity of integrated drainages, and global rainfall and denudation in addition to the traditional paradigm of eustatic control. Eustasy need not be the dominant control on sand-rich deep-sea turbidite deposition on a global scale.

232

ACKNOWLEDGEMENTS

The authors thank Andrew Madof, Kevin Pickering, Piret Plink-Bjorklund, and Mike

Sweet for insightful reviews of this manuscript, and also thank Geology editor James

Schmitt. We thank the Stanford Project on Deep-water Depositional Systems (SPODDS) and Basin and Petroleum System Modeling (BPSM) programs for support, and thank

Zachary Sickmann for early input.

233

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FIGURES

Figure 5.1. (A) Cenozoic sea-level curves of Miller et al. (2005) in cyan and Miller et al.

(2020) in black. Bar charts with counts of (B) passive and (C) active margin sand-rich deposits for the Paleocene (Pal.), lower Eocene (lo. Eo.), and middle Eocene (m. Eo.) by continental margin. N. Am.—North America, S. Am.—South America. Detailed descriptions and references are in Appendix C.

239

Figure 5.2. (A) Paleocene, (B) early Eocene, and (C) middle Eocene locations studied, including with/without recorded sandy deep-sea systems. Paleomaps after Scotese (2001).

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Figure 5.3. Conceptual model of passive margin sandy turbidite deposition given climate controls on sediment supply (orange line) and accommodation (blue curve). Sediment supply increases with global temperature via its role in intensifying precipitation.

Accommodation increases rapidly with initial warming as ice melts and sea level rises.

Later warming (after ice melt) has a much weaker effect on sea level and accommodation.

Sand-rich deposition (yellow polygons) happens when, for a given climate state, sediment supply exceeds accommodation.

241

Figure 5.4. Schematic illustration of sand-rich passive margin deep-sea deposition under

(A) falling sea level, where a drop in shelfal accommodation drives deep-water turbidite deposition, and (B) high sea level, where climatically enhanced sediment supply overwhelms high accommodation.

242

CHAPTER 6

GLOBAL EOCENE-OLIGOCENE UNCONFORMITY DUE TO CLIMATIC

COOLING-DRIVEN SUBMARINE EROSION IN SEDIMENTARY BASINS

243

GLOBAL EOCENE-OLIGOCENE UNCONFORMITY DUE TO CLIMATIC

COOLING-DRIVEN SUBMARINE EROSION

IN CLASTIC SEDIMENTARY BASINS

Zachary F. M. Burton, Tim McHargue, and Stephan A. Graham

Department of Geological Sciences, Stanford University, Stanford, CA 94305, USA

ABSTRACT

Global sedimentary hiatuses are well-documented in ancient pelagic sediment (based almost exclusively on results from the Deep Sea Drilling Program), and include Paleocene,

Eocene-Oligocene boundary, and Miocene hiatuses. Less clear is the extent of these hiatuses into continental margin settings. We test the hypothesis that global hiatuses evident in pelagic sections are also manifested in deep-marine siliciclastic basins of continental margins globally. We chose to focus on the Eocene-Oligocene boundary and surveyed literature on siliciclastic basins to produce a global inventory of unconformities of this age. The Eocene-Oligocene boundary is a period of extreme global climatic transition characterized by dramatic cooling and the highest magnitude sea-level fall of the

Cenozoic (aside from the Quaternary). We find evidence for an Eocene-Oligocene unconformity in sedimentary basins along the continental margins of every continent. We document 93 locations where an Eocene-Oligocene unconformity has been recorded, including 28 locations along the African continental margin, 1 along the Antarctic margin,

244

12 along Asian margins, 12 in the Oceania region, 12 along European margins, 19 along the North American margin, and 9 along the South American margin. Of the 93 total locations, 37 are in passive margin settings and 56 are in active margin settings or passive margin settings with local tectonism. Diverse oceanographic and sedimentary processes may be responsible for the development of widespread unconformities. Hiatuses in pelagic sections have been attributed to nondeposition, increased corrosiveness of bottom waters and attendant carbonate dissolution, and erosion associated with the intensification of ocean circulation. On clastic continental margins, additional mechanisms must be considered. For example, along the west African margin, sea-level lowering and attendant subaerial shelf exposure have been invoked to explain widespread Eocene-Oligocene erosional unconformities. However, we find that submarine erosion surfaces in deep-water settings are also common in clastic sections at the Eocene-Oligocene boundary. We speculate that one likely mechanism results from the lowered global temperatures and expansion of polar ice at this time, which accelerated ocean thermohaline currents and likely enhanced the potential for submarine erosion.

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6.1. INTRODUCTION

6.1.1. THE EOCENE-OLIGOCENE TRANSITION

6.1.1.1. Climatic change

The Eocene-Oligocene boundary marks a significant transition from the warm, relatively ice-free hothouse and greenhouse global climatic regimes of the Eocene to the cool icehouse conditions that have typified the Oligocene to recent times (Kennett, 1977; Miller et al., 1987, 1991, 2005, 2020; Zachos et al., 2001; Liu et al., 2009; Westerhold et al.,

2020). Global deep-marine benthic foraminifer oxygen and carbon isotope records demonstrate that the Eocene-Oligocene Transition (EOT) was the most significant, highest-magnitude transition of the entirety of the Cenozoic (Zachos et al., 2001; Miller et al., 2020; Westerhold et al., 2020). This transition is marked by the rapid (e.g., ~300 kyr from mostly deglaciated to significantly glaciated; Coxall et al., 2005) appearance of large, well-developed Antarctic ice sheets and a significant increase in continental ice volume

(Miller et al., 1991, 2020; Zachos et al., 2001; DeConto et al., 2008; Liu et al., 2009; Spray et al., 2019; Westerhold et al., 2020).

6.1.1.2. Sea-level change

An episode of global sea-level fall coeval with the EOT and ice sheet growth has been widely documented by workers over the past several decades (e.g., sequence stratigraphic

“cycle chart” work of Vail et al., 1977; Haq et al., 1987, 1988; oxygen isotopic work of

Miller et al., 1987, 1991), and more recent work on quantifying amplitudes of sea-level change has demonstrated that—outside of the Quaternary—the EOT marked the largest

246 fall (~55 m) in global mean sea level in the whole Cenozoic (e.g., backstripping work of

Miller et al., 2005; oxygen isotopic and Mg/Ca work of Miller et al., 2020).

6.1.1.3. Ocean circulation changes

Major changes in deep ocean circulation also occurred at or near the EOT, with plate reorganizations (i.e., gateway openings) and climatic change and Antarctic glaciation itself competingly invoked in longstanding, ongoing debate regarding the driving forces behind these major oceanographic changes (e.g., Benson, 1975; Kennett and Shackleton, 1976;

Kennett, 1977; Wright and Miller, 1993; Diester-Haass and Zahn, 1996; Davies et al.,

2001; Huber and Nof, 2006; Via and Thomas, 2006; Cramer et al., 2009; Miller et al.,

2009; Borrelli et al., 2014; Goldner et al., 2014; Elsworth et al., 2017; Coxall et al., 2018;

Hutchinson et al., 2019; Toumoulin et al., 2020).

Though consensus is lacking on the relative influence of tectonically-driven versus climatically-driven (i.e., cooling-driven) forcing of the global ocean circulation changes at the Eocene-Oligocene boundary, it bears mentioning that shifts in global climate are often associated with significant changes to deep ocean circulation at other epochs in Earth history, with climate (particularly temperature change) carrying the ability to modify deep- water circulation, and vice versa (e.g., Shackleton et al., 1983; Duplessy and Shackleton,

1985; Boyle and Keigwin, 1987; Oppo and Fairbanks, 1987; Duplessy et al., 1988; Kennett and Stott, 1991; Flower and Kennett, 1994; Marlow et al., 2000; Rahmstorf , 2002;

Piotrowski et al., 2004, 2005; Schott et al., 2009).

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6.1.2. DEEP-MARINE HIATUSES

6.1.2.1. Deep-marine hiatuses in pelagic sections

Early biostratigraphic work in the 1970s (e.g., Johnson, 1972; Kennett et al., 1972; Rona,

1973; Moore et al., 1978) and a reasonably large body of biostratigraphic work by Keller and coworkers concentrated in the 1980s (e.g., Barron and Keller, 1982; Keller and Barron,

1983, 1987; Keller, 1985; Keller et al., 1987; MacLeod and Keller, 1991) documented particular periods (mostly in the Cenozoic, but including the Cretaceous as well) of abundant globally-distributed deep-sea hiatuses in (almost exclusively biogenic) pelagic sediment investigated during the Deep Sea Drilling Project (only Johnson (1972) of the aforementioned studies was not based on DSDP data, instead relying on free-fall/drop core sampling of at most several m of subseafloor sediment).

Of these studies, several document regional or global deep-marine hiatuses of specifically

Eocene-Oligocene boundary age or at least including this age interval—Kennett et al.

(1972) discusses hiatus development in the southwest Pacific, while others discuss somewhat more globally-distributed (but still of a somewhat concentrated/focused paleogeographic distribution) hiatuses (Rona, 1973; Moore et al., 1978; Keller et al., 1987).

6.1.2.2. Mechanisms for hiatus development in pelagic sections

Various mechanisms have been proposed to explain the occurrence of these deep-marine hiatuses distributed throughout the Cenozoic, including, broadly speaking, ocean circulation-driven erosion, chemical dissolution, nondeposition, and sea-level fall.

Initiation and/or intensification of ocean circulation (e.g., more vigorous, higher-velocity deep-sea currents/bottom waters) and resultant increased capacity for mechanical erosion

248 of deep-sea sediment are invoked by numerous of the aforementioned studies to explain, at least in part, these deep-marine hiatuses. Moore et al. (1978) note that numerous clusters of major hiatuses throughout the Cenozoic seem especially related to changes affecting high-latitude regions where bottom waters form. Keller and coworkers invoke flow paths of and changes in major currents, generally speaking, as the primary driver of most major hiatuses in the middle Eocene through earliest Miocene (Keller et al., 1987) and in the

Miocene (Keller and Barron, 1983). Keller and Barron (1987) document “megahiatus” development during the Neogene, which they conclude was driven by circulation-driven erosion. Rona (1973) cites an existing view that erosion by bottom currents tends to be stronger/focused along western boundaries of ocean basins. Kennett et al. (1972) and

Moore et al. (1978) call upon changes in the shape of ocean basins driven by plate movements (e.g., gateway openings mentioned in section 6.1.1.3, above) as a primary driver of changes to circulation dynamics, while Moore et al. (1978) also note that such plate movements may result in physical displacement of a particular locality from an area of deposition to one of erosion. Numerous authors invoke glaciation as a potential driver of hiatus-forming ocean circulation changes, including discussion of Antarctic glaciation in the Eocene and Oligocene (Kennett et al., 1972), Northern hemisphere continental glaciation in the Pliocene (Moore et al., 1978), and glacial stages in the Pleistocene

(Johnson, 1972).

Chemical changes, which may themselves in some cases be influenced or driven by changes to ocean circulation and changes in sea level, can influence hiatus development, particularly by inducing the dissolution (i.e., the corrosion) of calcite- or silica-based planktonic tests. More specifically, “old” bottom waters tend to be more corrosive to calcite

249 tests of foraminifera and coccoliths because such waters are relatively CO2-rich (Moore et al., 1978; Keller and Barron, 1987), while “young” (recently formed at/near the sea surface) bottom waters tend to be more corrosive to opaline silica tests of diatoms, Radiolaria, and silicoflagellates because such waters are generally poor in dissolved silica (Moore et al.,

1978). Similarly, shifting carbonate compensation depths also affect the dissolution of calcite and aragonite (Rona, 1973).

Nondeposition can also explain some hiatuses. Nondeposition may be caused by, for example, low levels of primary productivity in near-surface waters, such as at centers of ocean gyres (Moore et al., 1978; Keller and Barron, 1987), focusing of carbonate deposition on shelves and resultant sediment starvation in the deep sea (Keller et al., 1987), and relatedly, nondeposition and decreases in depositional rates in the deep ocean due to retreat of depositional centers across the continental shelf during sea-level rise (e.g., as invoked for the latest Cretaceous–earliest Paleocene hiatuses; MacLeod and Keller, 1991), and ocean currents themselves as a mechanism for nondeposition, without needing invoke erosion (Moore et al., 1978).

Eustatic sea-level fall is also cited as a potential mechanism contributing to these observed deep-marine hiatuses. Rona (1973) note that sea-level fall may drive physical changes (e.g., continental emergence and erosion on continents, shelves, and slopes—i.e., not so much as an explanation for any of the aforementioned deep-sea hiatuses in pelagic sections) as well as chemical changes (e.g., due to depth-dependent parameters such as the carbonate compensation depth; terrestrial chemical weathering and dissolved load/solute transport to ocean systems). Keller and Barron (1983) invoke cool periods with attendant lower eustatic

250 sea levels and widely-distributed deep-marine erosion as a dominant process in the second half of the Miocene.

6.1.2.3. Mechanisms for hiatus development in pelagic sections at the Eocene-Oligocene boundary

Authors documenting hiatuses at or near the Eocene-Oligocene boundary provide some speculation on and discussion of potential mechanisms for the occurrence of such hiatuses.

Rona (1973) speculates that the hiatus may be partly due to initiation of present-day patterns of thermohaline circulation due to plate motions and climatic change, and, partly due to physical and chemical changes due to eustatic sea-level fall driven by decreases in global mid-ocean ridge system volumes. Moore et al. (1978) state that the hiatuses (which are most prominent throughout the southwest Pacific and eastern Indian ocean with additional, somewhat high hiatus abundance in the northern but not southern Atlantic, and with decreased areal extent of hiatuses in the eastern Pacific and western Indian Ocean) are apparently related to plate movement and resultant changes to ocean basin geometries

(especially gateway/passage opening and/or deepening, including specific mention of the opening of the Australia-Antarctic seaway, the Norwegian Sea-Arctic passage, the Drake

Passage, and the Tasman Passage, as well as subsidence of the Greenland-Faroe sill—it must be noted, of course, that the timing of some or all of these specific tectonic and paleoceanographic events has often been the focus of continued investigation, refinement, and contention, as alluded to in section 6.1.1.3, above). Kennett et al. (1972) speculate that the widespread regional unconformity in the southwest Pacific was caused by circulation changes and resultant deep-marine erosion brought about by plate movement in Australia and by Antarctic glaciation. Keller et al. (1987) mention a hiatus at the Eocene-Oligocene

251 boundary in their survey of numerous late Paleogene hiatuses, but state that this hiatus is less extensive (with regard to geographic distribution) than others found throughout the late Paleogene, and additionally note that this Eocene-Oligocene hiatus is more restricted to higher latitudes than the other surveyed hiatuses, and finally note that faunal changes during this time are less pronounced than during other late Paleogene hiatuses.

6.1.2.4. Deep-marine hiatuses in clastic sections at the Eocene-Oligocene boundary

While the development of Eocene-Oligocene marine hiatuses within pelagic sedimentary successions has been explored on a globally integrated scale (as described above), there is a relative dearth of work focusing on the documentation of unconformities within deep- marine clastic successions at an interregional scale, let alone at a global scale. This gap in our understanding remains, despite calls by workers as long as 45 years ago for the investigation of the effects of dramatic sea-level fall on sedimentation and erosion at the

Eocene-Oligocene boundary (e.g., Ingle et al., 1976; Miller et al., 1987)

Examples of the relatively few cases in which workers document somewhat widespread

Eocene-Oligocene unconformities developed within deep-marine clastic successions—and invoke changes in climate and sea-level at the EOT to explain these unconformities— include the work of Séranne (1999), which documents Eocene-Oligocene unconformities observed along the offshore western African continental margin, from Ghana to Angola.

Séranne (1999) suggests that these features developed as a widespread response to the

EOT, and more specifically, as a result of sea-level fall and resultant subaerial shelf exposure. Similarly, while a widespread Eocene-Oligocene unconformity in the North Sea had initially been attributed to tectonic uplift (Jordt et al., 1995), later workers proposed that the unconformity is instead mostly attributable to changes in climate and sea level at 252 the EOT (Clausen, 1998; Clausen et al., 1999; Huuse, 2002). Huuse (2002) in particular notes that the importance of changes to global climate and sea level may have been drastically underappreciated by previous work on the Cenozoic in the North Sea.

In response to this apparent gap in our understanding of the deep-marine clastic depositional system response to extreme cooling, glaciation, sea-level fall, and intensification of ocean circulation at the EOT, we here undertake the effort to compile and synthesize published examples of Eocene-Oligocene unconformities within clastic successions. This work provides critical insight into the distribution of such erosional features in the deep-marine realm, and requires global mechanistic controls be called upon to explain the globally widespread distribution of these unconformities.

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6.2. METHODS

We reviewed existing literature to compile and describe published examples of unconformities of Eocene-Oligocene age within deep-marine clastic sections, predominantly within sedimentary basins. We gather and document information including descriptions of bounding sedimentary sections and regional stratigraphy, regional to local paleogeographic and tectonic setting, and specific documentation of or speculation on climate, tectonics, paleoceanography, and other mechanisms invoked or potentially responsible for development of the published stratigraphies, including any record of

Eocene-Oligocene unconformities. This information is discussed, where relevant as context or for more granular understanding of observations, in the text that follows, while this synthesized information is supplied in full detail within Appendix D.

Much of the compiled data described herein come from offshore drilling studies and from onshore outcrop studies, wherein sedimentological assessment of rock (including unconformity identification) was performed directly, while the remaining examples come from published analysis of seismic data (given that the sequence stratigraphic community has for decades refined criteria for identification of unconformities in seismic datasets, e.g.,

Vail et al., 1980). It must be emphasized that this synthesis inherently assumes the accuracy of the publications and reports that we have surveyed, including age dating methods (e.g., geochronology, biostratigraphy) employed, quality of data collection and analysis methodologies utilized, proposed stratigraphic interpretations, and interpretations of paleoenvironment, to include depositional settings, paleowater depth estimates, local to regional paleoceanographic conditions and climatic parameters, and local to regional influence of tectonism, onshore fluvial systems, and offshore sedimentological processes.

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Because our compilation has been conducted by synthesizing data and findings from published work, the number of deep-marine unconformities we document here must be treated as a lower bound on the actual number of global Eocene-Oligocene unconformities.

We also acknowledge that we likely missed examples published in non-English language sources. Importantly, there are likely a substantial number of examples either not yet encountered (e.g., due to being located in remote frontier settings, or settings that have seen little deep-sea drilling of thick clastic sections in sedimentary basins—for instance, the

Atlantic margin of the United States, as discussed below and as detailed in Appendix D) or that have been encountered but not published or otherwise publicly reported (e.g.,

Eocene-Oligocene unconformities encountered during drilling by oil and gas exploration companies). Finally, there are likely examples that simply were not preserved in the rock record (e.g., due to subduction and destruction).

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6.3. RESULTS

In total, we document 93 distinct locations—along the margins of all seven continents— where unconformities of Eocene-Oligocene age have been recorded in the literature (Figure

6.2). These include 28 locations along the African continental margin, 1 along the Antarctic margin, 12 along Asian margins, 12 in the Oceania region, 12 along European margins, 19 along the North American margin, and 9 along the South American margin (Figure 6.2).

Of the total 93 locations, 37 are in passive margin settings and 56 are in active margin settings or passive margin settings with active tectonism documented for this time (Figure

6.2).

Full descriptions of documented locations and geologic context are provided in Appendix

6.3.1. PASSIVE MARGIN UNCONFORMITIES

Passive margin unconformities documented here are concentrated along the southern

Atlantic margin (4 locations along eastern South America, 14 locations along western

Africa, and 1 location on the Weddell Sea margin of Antarctica), the northernmost margin of the Indian Ocean (4 locations), the Australian continental margin (9 locations), and the

South China Sea (2 locations) (Figure 6.2). The rifted Atlantic margin of the continental

United States, the Arctic margin, and the Antarctic margin are notable in their dearth of documented Eocene-Oligocene unconformities (Figure 6.2).

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6.3.2. ACTIVE MARGIN UNCONFORMITIES

Unconformities in active margin and regionally tectonically-active passive margin settings are abundant along the western margin of North America (17 locations from southern

California to Alaska’s Bering Sea), the northern African margin and southern European margin (6 locations and 7 locations, respectively), the general Barents, Norwegian, and

North seas region (5 locations), the northeastern Arabian plate (5 locations), southeastern

Africa’s passive margin (3 locations), and the northwestern-most South American margin

(3 locations) (Figure 6.2). Additional localities include those in Oceania (3 locations), the

Bengal Basin (1 location), the Labrador Sea (1 location), northern Hispaniola (1 location),

Espírito Santo Basin on Brazil’s eastern passive margin (1 location), and the Fuegian

Andes (1 location) (Figure 6.2). The active western Pacific margin, Indonesian margin, and much of the Andean margin lack documented Eocene-Oligocene unconformities (Figure

6.2).

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6.4. DISCUSSION

6.4.1. REGIONS LACKING UNCONFORMITIES

As stated in section 6.2, above, an apparent lack of unconformities (i.e., lack of documented unconformities) in a particular region is not wholly conclusive evidence for an actual lack of unconformities (i.e., for an actual absence of the processes that would be responsible for creating such unconformities, if present) in that region.

Thus, the apparent lack of passive margin unconformities, particularly along swaths of the rifted western Atlantic, Arctic, and Antarctic margins, as noted above, must be interpreted with some caution. Namely, one might feasibly construct a defensible claim asserting that such a dearth is indicative of the distribution of (i.e., regional absence of) processes driving submarine erosion or nondeposition along these margins, and yet, one might just as feasibly realize that, in reflecting on the nature of our current state of knowledge in these regions

(particularly when compared critically to the distribution of regions where presence of passive margin unconformities has been confirmed in the literature), that these regions are drastically underexplored with regard to the drilling of deep-sea sedimentary basins and thick clastic sections in these regions (as detailed further in Appendix D)—and, this dearth of drilling correlates more specifically with a lack of petroleum exploration activity along both the Atlantic margin of the United States (mostly due to lack of exploration success or perceived prospectivity) and Antarctica (due to a ban on resource extraction) as well as with the remoteness of both the Antarctic and Arctic margins.

This cautionary note on interpretation of a lack of unconformities may also be applied to active margins, however, in examination of the three active margins broadly lacking in

258 unconformities of Eocene-Oligocene age, it is likely immediately apparent to the reader that these three margins make up major subduction zones. Thus, one could well imagine that absence of Eocene-Oligocene unconformities—and a dearth of clastic sections in general—along these margins may be explained by a general lack of preservation of clastic sections of this age due to subduction-driven destruction.

6.4.2. COMPARISON WITH SEISMIC SEQUENCE STRATIGRAPHIC WORK

Seismic-based and sequence stratigraphic assessment by workers from the Exxon

Production Company (e.g., Figure 6.3) provided the foundations for much of our understanding of Cenozoic sea level (e.g., Vail et al., 1977; Haq et al., 1987)—as well as, of course, foundational models describing and predicting the nature of deposition and erosion in sedimentary basins. This early work is still considered useful as a chronology of changes to sea level, but, significant issues—and therefore, inaccuracies—have been noted with regard to estimations of absolute sea level and of the amplitude of sea-level changes

(e.g., as discussed at some length by Miller et al., 2020).

Relevant to the work at hand is that these early workers interpret the largest falls in

Cenozoic eustatic sea level to have occurred in the middle Oligocene and late middle

Miocene, yet more recent, multiproxy-based developments in our understanding of

Cenozoic sea level strongly indicate that the highest amplitude events occurred at the EOT

(a ~55 m fall in sea level) and during the Quaternary glacials (Miller et al., 2005, 2020).

Relatedly, particularly relevant is that these early workers from Exxon note that major interregional unconformities are associated with the major falls in sea level that they interpret throughout the Phanerozoic, and, do not record such an interregional

259 unconformity for the EOT (Vail et al., 1977). Furthermore, a study from these workers specifically focused on unconformities in the North Atlantic found no regional EOT-age unconformity—either along the western Atlantic margin at the Blake Escarpment (which we discuss in Appendix D) or in the eastern Atlantic along the African margin (Vail et al.,

1980).

We propose that, in contrast to the conclusions put forth by early workers at the Exxon

Production Company (e.g., Vail et al., 1977, 1980), a major global unconformity did develop at the EOT. We are confident in this conclusion for numerous reasons.

Firstly, the aforementioned and above-discussed fundamental changes in our understanding of Cenozoic sea-level and consensus around newer sea-level curves (e.g.,

Miller et al., 2005, 2020) developed over the past 15 years provide a much-improved quantification of fluctuations in sea-level and the amplitude/relative significance thereof, and, provide (alongside the ever-evolving body of EOT-related work, e.g., as discussed in our section 6.1, above) a strong theoretical basis for major changes (i.e., possible unconformity-driving processes) at the EOT.

Our conclusions also benefit from the intentional design and scope of our work. For instance, this includes the geographic breadth of our study (i.e., a truly global survey), whereby we document unconformities at 93 locations across all continents (and have investigated many more locations than this, as detailed in part in Appendix D, during our survey)—this geographic distribution is more expansive, for example, than that of the 49 locations which serve as the basis for the early Exxon sea-level work of Vail et al. (1977), as seen in comparing Figure 6.2 and Figure 6.3 (e.g., note that over 50% of locations used by Vail et al., 1977, are in North America). Furthermore, our work focuses specifically on 260 the EOT, while the work of Vail et al. (1977) intentionally focused on the entirety of the

Phanerozoic—that is, our 93 locations on Figure 6.2 are locations of EOT-age unconformities, while the 49 locations on Figure 6.2 from Vail et al. (1977) are for sections chosen to together capture 540 million years of stratigraphy. A point hopefully of benefit to the broader scientific community also lies in the feasible replicability of our work, in that all literature references—which form the basis for our compiled locations, as well as our syntheses thereof—are available in Appendix D, whereas, for example, an issue raised by Miller et al. (2020) of the early Exxon work is the lack of transparency or access to the data that forms the basis for the conclusions of these workers (however, is not surprising, given the proprietary nature of most data gathered in furtherment of hydrocarbon exploration efforts).

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6.5. CONCLUSIONS

The Eocene-Oligocene Transition represents one of the most severe—and by many metrics, the most severe—transition in global climate during the Cenozoic. This transition was marked by severe cooling, establishment of significant Antarctic glaciation, one of the two highest-amplitude sea-level falls in the Cenozoic, and a time of major changes to ocean circulation (namely, intensification of deep-marine currents). We undertake the compilation of published work to synthesize and document evidence for changes to deep- marine siliciclastic systems during this time—as reflected by the presence of deep-marine unconformities in clastic successions of this age. Accordingly, we document 93 locations where an Eocene-Oligocene unconformity has been documented (28 locations along the

African continental margin, 1 along the Antarctic margin, 12 along Asian margins, 12 in the Oceania region, 12 along European margins, 19 along the North American margin, and

9 along the South American margin). Of these 93 total locations, 37 occur in Eocene-

Oligocene passive margin settings and 56 are in active margin settings or passive margin settings with local tectonism. Diverse oceanographic and sedimentary processes may be responsible for the development of globally widespread unconformities. We speculate that these unconformities are clear manifestation of major oceanographic change at the Eocene-

262

ACKNOWLEDGEMENTS

We acknowledge the Stanford Project on Deep-water Depositional Systems (SPODDS) and the Stanford Basin and Petroleum System Modeling (BPSM) industrial affiliates programs for funding this research. Additional financial support to Z.F.M.B. was provided by the AAPG Grants-in-Aid Program Lawrence W. Funkhouser Named Grant, the Stanford

Earth McGee/Levorsen Graduate Student Research Grant Program, and the Shell

Foundation Grant Program. We thank Colin White and Matthew Malkowski for helpful discussion of Bering Sea and Alaskan stratigraphy and of Gulf of Mexico stratigraphy.

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272

FIGURES

Figure 6.1. Cenozoic sea-level curves based on backstripping analysis (Miller et al., 2005), in cyan, and oxygen isotopic and Mg/Ca analysis (Miller et al., 2020), in black.

273

Figure 6.2. Eocene-Oligocene boundary unconformities documented in deep-marine clastic sections (this study). Black circles indicate passive margin setting. Red triangles indicate active margin setting or passive margin setting with active tectonism. Map constructed using GeoMapApp v. 2019.

274

Figure 6.3. Locations of the seismic datasets used as the basis for the seismic sequence stratigraphic work by Vail et al. (1977) to build their global sea-level cycle charts. Note that these locations were used to construct cycle charts from 541 Ma to the present (i.e., for the entirety of the Phanerozoic). Figure after Fig. 4 of Vail et al. (1977).

275

APPENDIX A

SUPPLEMENTARY MATERIAL FOR CHAPTERS 1 AND 2

APPENDIX A-1. ANALYTICAL DATA: TABLES

Table A-1.1. Results of biomarker compound quantitation (in ppm) from gas chromatography−mass spectrometry (GC−MS) saturated hydrocarbon fraction analysis.

Table A-1.2. Selected biomarker compound ratios from gas chromatography−mass spectrometry (GC−MS) saturated hydrocarbon fraction analysis.

Table A-1.3. Results of biomarker compound quantitation (in ppm) achieved via compound peak area integration and/or compound peak height measurement relative to internal and external standards. From gas chromatography−tandem mass spectrometry

(GC−MS-MS) saturated hydrocarbon fraction analysis.

Table A-1.4. Selected biomarker compound ratios from gas chromatography−tandem mass spectrometry (GC−MS-MS) saturated hydrocarbon fraction analysis.

Table A-1.5. Results of diamondoid compound quantitation (in ppm) from gas chromatography−mass spectrometry (GC−MS) saturated hydrocarbon fraction analysis.

276

APPENDIX A-2. ANALYTICAL DATA: CHROMATOGRAMS

Figure A-2.1. m/z 191 fragmentogram illustrating the distribution of hopanes in the

Rotokautuku oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

Figure A-2.2. m/z 191 fragmentogram illustrating the distribution of hopanes in the

Totangi oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

Figure A-2.3. m/z 191 fragmentogram illustrating the distribution of hopanes in the

Waitangi oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

Figure A-2.4. m/z 191 fragmentogram illustrating the distribution of hopanes in the

Isolation Creek 1 oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

Figure A-2.5. m/z 191 fragmentogram illustrating the distribution of hopanes in the

Kaikoura/ Isolation Creek 2 oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

Figure A-2.6. m/z 217 fragmentogram illustrating the distribution of steranes in the

Rotokautuku oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

277

Figure A-2.7. m/z 217 fragmentogram illustrating the distribution of steranes in the

Totangi oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

Figure A-2.8. m/z 217 fragmentogram illustrating the distribution of steranes in the

Waitangi oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

Figure A-2.9. m/z 217 fragmentogram illustrating the distribution of steranes in the

Isolation Creek 1 oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

Figure A-2.10. m/z 217 fragmentogram illustrating the distribution of steranes in the

Kaikoura/ Isolation Creek 2 oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

278

APPENDIX A-1. ANALYTICAL DATA: TABLES

Table A-1.1. Results of biomarker compound quantitation (in ppm) from gas chromatography−mass spectrometry (GC−MS) saturated hydrocarbon fraction analysis.

279

Table A-1.1. Continued.

280

Table A-1.1. Continued.

281

Table A-1.2. Selected biomarker compound ratios from gas chromatography−mass spectrometry (GC−MS) saturated hydrocarbon fraction analysis.

282

(GC−MS-MS) saturated hydrocarbon fraction analysis.

283

Table A-1.3. Continued.

284

Table A-1.3. Continued.

285

Table A-1.4. Selected biomarker compound ratios from gas chromatography−tandem mass spectrometry (GC−MS-MS) saturated hydrocarbon fraction analysis.

286

Table A-1.4. Continued.

287

Table A-1.4. Continued.

288

Table A-1.5. Results of diamondoid compound quantitation (in ppm) from gas chromatography−mass spectrometry (GC−MS) saturated hydrocarbon fraction analysis.

289

APPENDIX A-2. ANALYTICAL DATA: CHROMATOGRAMS

Figure A-2.1. m/z 191 fragmentogram illustrating the distribution of hopanes in the

Rotokautuku oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

290

Figure A-2.2. m/z 191 fragmentogram illustrating the distribution of hopanes in the

Totangi oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

291

Figure A-2.3. m/z 191 fragmentogram illustrating the distribution of hopanes in the

Waitangi oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

292

Figure A-2.4. m/z 191 fragmentogram illustrating the distribution of hopanes in the

Isolation Creek 1 oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

293

Figure A-2.5. m/z 191 fragmentogram illustrating the distribution of hopanes in the

Kaikoura/ Isolation Creek 2 oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

294

Figure A-2.6. m/z 217 fragmentogram illustrating the distribution of steranes in the

Rotokautuku oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

295

Figure A-2.7. m/z 217 fragmentogram illustrating the distribution of steranes in the

Totangi oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

296

Figure A-2.8. m/z 217 fragmentogram illustrating the distribution of steranes in the

Waitangi oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

297

Figure A-2.9. m/z 217 fragmentogram illustrating the distribution of steranes in the

Isolation Creek 1 oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

298

Figure A-2.10. m/z 217 fragmentogram illustrating the distribution of steranes in the

Kaikoura/ Isolation Creek 2 oil seep sample. Obtained via GC−MS analysis of the saturate hydrocarbon fraction.

299

APPENDIX B

SUPPLEMENTARY MATERIAL FOR CHAPTER 2

APPENDIX B-1. CORRECTION TO INTERNATIONAL JOURNAL OF EARTH

SCIENCES PUBLICATION

Below is the text of a correction to the International Journal of Earth Sciences article that forms the body of Chapter 2. This correction was published at the request of GNS Science coauthor Richard Sykes.

300

APPENDIX C

SUPPLEMENTARY MATERIAL FOR CHAPTER 5

APPENDIX C-1. DETAILED METHODOLOGY

APPENDIX C-2. DETAILED DESCRIPTION OF TURBIDITE OCCURRENCES

APPENDIX C-2.1. Africa

APPENDIX C-2.2. Antarctica

APPENDIX C-2.3. Asia

APPENDIX C-2.4. Australia/Oceania

APPENDIX C-2.5. Europe

APPENDIX C-2.6. India

APPENDIX C-2.7. North America

APPENDIX C-2.7.1. Northern America

APPENDIX C-2.7.2. Central America and The Caribbean

APPENDIX C-2.8. South America

APPENDIX C-3. REFERENCES CITED

301

APPENDIX C-1. DETAILED METHODOLOGY

We compiled a global database of published examples of marine (primarily deep marine) sedimentary deposits of Paleocene, early Eocene, and middle Eocene age. We report records of deposition at 114 global sites (including 9 sites from Africa, 0 sites from

Antarctica, 15 sites from Asia, 14 sites in Australia/Oceania, 17 sites in Europe, 6 sites in

India, 34 sites in North and Central America, and 19 sites in South America), though it should be noted that we surveyed many more sites that we have here excluded because they contained no record of Paleocene, lower Eocene, and/or middle Eocene deep-water deposits (e.g., the margins of Antarctica). Furthermore, a number of the 114 sites we did survey themselves represent numerous locations, basins, and/or geographic regions (e.g., the rift basins of eastern India are grouped together), as detailed below in Appendix C-2.

Our database relies on published examples of marine turbidites and deep-water sediments of Paleocene through Eocene age. These examples almost entirely consist of onshore outcrop studies and offshore drilling studies, in which sedimentological/lithological assessment (e.g., identification of turbidites) has been performed based on direct access to

Paleocene, lower Eocene, and/or middle Eocene rock, although in rare instances we include examples of inferred turbidite sediment occurrence based on published analysis of seismic data (e.g., poorly explored regions such as the East Siberian Shelf).

Because our study is in essence a compilation of previously published data and interpretations, it inherently relies on the validity and accuracy of this previously published work. We rely on the interpretations of previous authors in aspects including the interpreted age of sediment, the interpreted sedimentology and lithology of sediment, the interpreted

302 depositional environment of sediment, and some aspects of associated local and regional tectonic and climatic associations.

Our study is also impacted purely by limitations in the availability and comprehensiveness of published examples of deep marine sedimentation, including examples of turbidites. An assumption inherent to compiling a global database of such examples is that existing published examples are representative of the true global distribution of examples of such marine sediments. However, our record, and the published record in general, is most certainly incomplete, and most certainly underrepresents the total number of turbidite and deep-water sediment occurrences of any given age. Furthermore, we acknowledge that we likely missed examples of turbidites of this age published in non-English language sources.

Thus, our compilation provides a lower bound to the total number of turbidite deposits from the Paleocene, early Eocene, and middle Eocene.

One limitation to our study is that many regions of the world, including remote onshore terrains and difficult-to-access offshore basins, are currently poorly explored (e.g., much of the Arctic). Some of these regions may have further examples of Paleocene- through middle Eocene-age turbidite deposits. Similarly, there may be turbidites of this age that have been encountered during drilling and exploration efforts by oil and gas companies across the world, but are not reported in the published literature.

Poor preservation of sedimentary deposits from the Paleogene may pose a further limitation to the number of recorded Paleogene turbidites. For instance, some Paleocene- and Eocene- age turbidites may have been subducted along active margins and thus lost to the rock record. Similarly, sediment of this age may have been eroded subsequent to deposition.

Any lack of preservation, as with a lack of exploration, would potentially yield 303 underestimation of the total number of Paleocene, lower Eocene, and middle Eocene turbidites.

304

APPENDIX C-2. DETAILED DESCRIPTION OF TURBIDITE OCCURRENCES

The following supplement details a region-by-region account of the presence or absence of

Paleocene, lower Eocene, and middle Eocene turbidite occurrences. This compilation includes brief descriptions of the paleogeographic setting (including regional and local tectonics, presence of integrated fluvial drainages, relative sea level, and climate, where available) of each region, and includes descriptions of the nature of sediment deposition during the Paleocene, early Eocene, and middle Eocene time intervals. Early Eocene paleomaps (with surveyed locations labeled) are provided for each region as follows:

Africa—Figure C-2.2, Asia—Figure C-2.3, Australia/Oceania—Figures C-2.4A and C-

2.4B, Europe—Figure C-2.5, India—Figure C-2.6, North (and Central) America—Figure

C-2.7, and South America—Figure C-2.8.

Figure C-2.1. Figure legend explaining symbols for all paleomap figures in this Data

Repository.

305

306

APPENDIX C-2.1. Africa

Figure C-2.2. Locations of Paleogene sedimentary deposits described for the African continental margin (early Eocene paleomap after Scotese, 2001). SB—Senegal Basin,

NDF—Nile delta fan, EG—Equatorial Guinea, OG—offshore Gabon, OA—offshore

Angola, OB—Orange Basin, MB—Mozambique Basin, RB—Rovuma Basin, TCB—

Tanzania Coastal Basin, LB—Lamu Basin, SA—southern Arabia, NA—North Africa.

Symbols as in Figure C-2.1.

307

Table C-2.1A. Locations of lower Eocene turbidites associated with the African continent, as well as descriptions of deposition during the neighboring time periods.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin Zambezi Delta (Mozambique very low same as same as same as Basin) clastic flux Paleocene Paleocene Paleocene Yes No

N Mozambique (Rovuma Basin) ? deepwater fan ? ? Yes No Tanzania turbidite same as same as same as Coastal Basin channels/lobes Paleocene Paleocene Paleocene ? No* shelfal SE Kenya marine/unconfo fluvial/deltaic/re same as early same as early (Lamu Basin) rmity stricted shelf Eocene Eocene Yes No

*indicates passive margin, but local tectonics

Table C-2.1B. African locations surveyed but lacking lower Eocene turbidites, including descriptions of deposition during the Paleocene through late Eocene.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin W African margin (Senegal Basin to Orange same as same as same as Basin) carbonate Paleocene Paleocene Paleocene No deltaic/delta- same as Niger delta fan ? shale front middle Eocene Yes No Equatorial minor Guinea reservoirs N/A N/A N/A ? No muddy W South Africa prograding same as same as same as (Orange River) wedge? Paleocene Paleocene Paleocene Yes No E & central Sahara (Egypt, Libya, Chad, Niger, Algeria, carbonate/phos same as Tunisia) phate Paleocene evaporites No? No

308

The west African margin rifted from South America in the Early Cretaceous with the breakup of the Gondwana supercontinent and opening of the Atlantic Ocean, and since then has existed as a passive margin (Rabinowitz and LaBrecque, 1979; Lavier et al.,

2001).

From the Early Cretaceous (Albian) to the Eocene-Oligocene transition, the margin was characterized by sediment-starved, shallow-water carbonate platform and carbonate ramp sedimentation at a rate matching post-rift subsidence (Brice et al., 1982; Seranne et al.,

1992; Lavier et al., 2001). Offshore sediments all the way from Senegal Basin in the north to Orange Basin off the coast of western South Africa are characterized by limestone and shale (Seranne and Abeigne, 1999; Brownfield and Charpentier, 2003; Monnier et al.,

2014). Paleogene transgressions combined with low-lying topography and low seasonality climate onshore led to low denudation rates, which, alongside drainage and sediment capture by lacustrine rift basins (Genik, 1992; McHargue et al., 1992), contributed to the dearth of clastic sedimentation along the Paleogene West African margin (Macgregor,

2012). This is reflected in the lack of Eocene-age petroleum reservoirs, and the presence of but a few minor Paleocene reservoirs off Equatorial Guinea (Macgregor, 2012; Thornton et al., 2013). A possible exception exists in the sedimentary thick beneath the present-day

Niger delta fan (Macgregor, 2012). Much of the sediment here seems to have developed in the Neogene (Damuth, 1994), however, Tuttle et al. (1999) report development of a Niger delta system as early as the Eocene, with deposition of delta-front and deltaic clastic sediments of the Agbada Formation inferred to have begun in the middle Eocene. The early

Eocene, however, is inferred to have been characterized by deposition of marine shales of the Akata Formation (Tuttle et al., 1999). The Congo deep-sea fan to the south was not

309 initiated until the Oligocene at the earliest (Savoye et al., 2009). The southern margin of

South Africa exists as an Atlantic-type shear margin characterized by very limited rates of sedimentation since the breakup of Gondwanaland, with the exception of large volumes of detritus delivered offshore western South Africa by the Orange River during the Cretaceous

(Emery et al., 1975; Scrutton and Dingle, 1976) and a muddy prograding wedge during the

Tertiary (Paton et al., 2010).

The southeastern to eastern margin of Africa also has existed as a passive margin since the

Late Cretaceous or earlier breakup of Gondwanaland (Norton and Sclater, 1979).

In the Mozambique Basin, transgression and low rates of clastic sediment flux are reported for Paleocene through Eocene intervals dominated by shales and marls with middle to upper Eocene limestone (Salman and Abdula, 1995; Said et al., 2015; Castelino et al.,

2017). The Zambezi Delta system contributed little sediment to the Mozambique Basin during this time, with Paleocene and Eocene clastic sediment flux of the delta the lowest for the entirety of the Early Cretaceous through Present (Walford et al., 2005). This contrasts sharply with the Rovuma Basin of northern Mozambique where hydrocarbon exploration has revealed a lower Eocene deep-water fan complex associated with progradation of the Ruvuma River delta (Palermo et al., 2014, 2015). The so-called lower

Eocene Coral reservoir, deposited over an estimated six to eight million years, is comprised of sand-rich turbidite channels and lobes resulting in extraordinarily thick (greater than 400 m thick, 15 km wide), clean, and homogenous sands, much like reservoir units of the Gulf of Mexico (Palermo et al., 2014). In the Tanzania Coastal Basin to the north, turbidite channel and lobe systems were deposited from the Paleocene to Oligocene in response to sediment influx due to onshore uplift in the African hinterland (Sansom, 2018). In the Lamu

310

Basin of southeast Kenya, Paleocene shelf and marine settings gave way to Eocene and

Oligocene fluvial, deltaic and delta-front, and restricted-shelf settings (Nyagah, 1995;

Zongying et al., 2013). Late Paleocene and/or early Eocene transgression, as well as warm, humid climates, prevailed in eastern Somalia and southern Arabia, leading to widespread calcareous deposition of the lower Eocene Auradu limestone (Bosellini, 1992; Bolle et al.,

2000). Overall, it should be noted that much is still unknown about deep-water East Africa, much of which remains an area of frontier exploration (Zhixin et al., 2015).

Warm, humid climates, as well as higher sea level, are similarly inferred for Late

Cretaceous to early Eocene eastern and central Sahara (Egypt, Libya, Chad, Niger, Algeria,

Tunisia), reflected by widespread deposition of carbonate and phosphate (Swezey, 2009).

By contrast, upper Eocene evaporitic strata reflect falling sea level and cooler climates

(Swezey, 2009). The north of Africa, namely Egypt and Libya, and the Eastern

Mediterranean Basin to the north, evolved along the northern passive margin of

Gondwanaland from the early Permian onward (Garfunkel, 1998). Since the mid-

Cretaceous, convergence between the African and Eurasian plates has largely affected the

Eastern Mediterranean Basin, and has resulted in structural complexity in North Africa

(Garfunkel, 1998). Nile sediments are thought to be no older than upper Miocene (Gaullier et al., 2000). The western and central Mediterranean basins opened as late as the Miocene through Quaternary in response to back-arc extension associated with North-South convergence between the African and Eurasian plates (Auzende et al., 1973; Rehault et al.,

1984; Dewey et al., 1989; Carminati et al., 1998).

311

Appendix C-2.2. Antarctica

Almost no pre-glacial, lower Cenozoic sedimentary rock has been recovered anywhere on Antarctica’s continental margin (Cooper et al., 1991), and documented instances of turbiditic deposition tend to be late Eocene or younger (Kuvaas and

Leitchenkov, 1992; Escutia et al., 2000; Close, 2010). Antarctica has occupied high latitudes since at least the Cretaceous, during which time it was connected with Australia

(Kennett, 1977). Starting in the early Eocene, Australia separated and drifted toward lower latitudes, forming an ocean adjacent to Antarctica, albeit one lacking circum-Antarctic flow

(Kennett, 1977). For much of the Eocene (and particularly the early Eocene; Pross et al.,

2012), warm temperatures meant Antarctica was largely nonglaciated, with widespread glaciation absent until global cooling at the Eocene-Oligocene boundary (Kennett, 1977;

Zachos et al., 2001). Interglacial times since the late Eocene are characterized by little to no clastic sedimentation on the Antarctic shelfs, with progradational shelf sequences, including fan systems, almost exclusively related to glacial activity (Cooper et al., 1991;

Kuvaas and Kristofferson, 1991).

312

Appendix C-2.3. Asia

Figure C-2.3. Locations of Paleogene sedimentary deposits described for the Asian continental margin (early Eocene paleomap after Scotese, 2001). TB—Thrace Basin,

CTB—central Turkey basins, LB—Levantine Basin, Z—Zagros flysch, Th—Thailand,

SCS—South China Sea Basin, Pa—Palawan Island, Sh—Shimanto Supergroup, SS—East

Siberian Shelf. Symbols as in Figure C-2.1.

313

Table C-2.3A. Locations of lower Eocene turbidites associated with the Asian continent, as well as descriptions of deposition during the neighboring time periods.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin turbidites & N/A carbonate NW Turkey (continental turbidites, mix of early & (shallow to (Thrace Basin) basement) fluvial, deltaic late Eocene deep marine) No Yes central Turkey basins (Haymana, same as same as same as Tuzgölü, etc.) turbidites Paleocene Paleocene Paleocene No Yes Zagros (SE Turkey, N same as Syria & Iraq, carbonate & same as same as Paleocene, but W & S Iran) flysch Paleocene Paleocene waning clastics No Yes central turbidites/muds same as same as same as Yes (syn-rift, Palawan Island tone Paleocene Paleocene Paleocene ? block faulting) N/A S Palawan (Cretaceous turbidites/muds same as early same as early Yes (syn-rift, Island basement) tone Eocene Eocene ? block faulting) Shimanto Supergroup (outboard of East China same as same as same as Sea) turbidites Paleocene Paleocene Paleocene No Yes Shimanto Supergroup (Okinawa same as trench fill same as Island) turbidites Paleocene turbidites middle Eocene No Yes

NW East Siberian Shelf (bordering same as same as same as Makarov Basin) turbidites? Paleocene Paleocene Paleocene Yes? No

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Table C-2.3B. Asian locations surveyed but lacking lower Eocene turbidites, including descriptions of deposition during the Paleocene through late Eocene.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin Levantine Basin (offshore Turkey, Syria, Lebanon, deepwater same as same as same as Israel, Egypt) chalk/marl Paleocene Paleocene Paleocene No Yes? around same as same as same as Arabian Plate carbonate Paleocene Paleocene Paleocene No ? Thailand ? carbonate? carbonate? carbonate? ? ? S & central Vietnam (South China Sea Basin) ? nonmarine? nonmarine? nonmarine? ? No (rift) N Vietnam & S China (South China Sea nonmarine/lacu same as early same as early Basin) ? strine Eocene Eocene ? No (rift) Taiwan (South China Sea nonmarine same as same as same as Basin) shale Paleocene Paleocene Paleocene ? No (rift) E China intracontinental basins, China eastern marginal seas basins (incl. East China Sea Basin, Yellow Sea nonmarine basins), Sea (alluvial & same as same as same as of Japan lacustrine) Paleocene Paleocene Paleocene ? No (rift)

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From the Mesozoic and Cenozoic to present, convergence and collision between the

Eurasian Plate and the northward-moving African, Arabian, and Indian plates has led to development of the extensive Alpine-Himalayan orogenic belt (Dewey and Bird, 1970).

In northwestern Turkey, the Thrace Basin evolved as a foreland basin from the Late

Cretaceous through Paleocene, and then as an intermontane molasse basin in the Eocene to present (Siyako and Huvaz, 2007). Earliest sedimentation in the basin was tectonically controlled and consisted of lower to middle Eocene molasse sediments deposited atop continental basement (Siyako and Huvaz, 2007). These lower to middle Eocene deposits consist of fluvial and deltaic sediments and proximal to distal turbidites making up some two to three kilometers of thickness in the Thrace Basin (Siyako and Huvaz, 2007) The middle to upper Eocene section of Thrace Basin is up to 1.5 kilometers in thickness, and consists of shallow and deep marine facies, including carbonates and proximal to distal turbidites, whereas post-Eocene sediments consist of terrestrial facies (Siyako and Huvaz,

2007). The basins of central Turkey or Anatolia evolved either as forearc basins in the Late

Cretaceous or collisional (peripheral foreland) basins beginning in the early Eocene (Görür et al., 1998). In both settings, marine turbidite sediment filled basins from the Late

Cretaceous or early Eocene until the Oligocene (Görür et al., 1998). For instance, sub- basins of the Tuzgölü Basin complex of central Turkey evolved as forearc basins from the

Late Cretaceous to late Paleocene, and were deformed during late Paleocene to Eocene block collisions (Görür et al., 1984). Molasse sedimentation occurred from the Late

Cretaceous to late Oligocene in both the Haymana and Tuzgölü sub-basins, although the

Late Cretaceous to late Eocene interval is characterized by some eight kilometers of turbiditic sediments whereas the post-Eocene consists entirely of terrestrial red-beds and

316 evaporites (Görür et al., 1984). The Adana Basin of south-central Turkey is Neogene in age (e.g., Williams et al., 1995).

To the south, opening of the Levantine Basin off the coast of Turkey, Syria, Lebanon,

Israel, and Egypt occurred in response to early Mesozoic Pangaea break-up, and was followed by the start of basin closure in the mid- to Late Cretaceous (Gardosh and

Druckman, 2006). Late Cretaceous to early Eocene sedimentation within the basin is inferred to have consisted of pelagic to hemipelagic deep-water chalk and marl (Gardosh and Druckman, 2006). Middle Eocene carbonates are followed by a depositional hiatus represented by an abrupt transition to Oligocene to Miocene siliciclastics (Gardosh and

Druckman, 2006). Further south, deposition around the Arabian Plate was similarly carbonate-dominated for much of the Cenozoic, including the Paleogene interval of interest

(Alsharhan and Nairn, 1995; Brannan et al., 1997; Beavington-Penney et al., 2006;

Garzanti et al., 2013).

The Early Cretaceous to recent Zagros fold-thrust belt constitutes a portion of the Alpine-

Himalayan orogenic belt and stretches some 2000 kilometers across southeastern Turkey, northern Syria and Iraq, and western and southern Iran (Alavi, 2004). The latest

Maastrichtian to late Eocene here forms one of three Late Cretaceous to recent proforeland megasequences (Alavi, 2004). This period is characterized by latest Cretaceous to middle

Eocene progradational sediments deposited during a period of reduced tectonic activity, and middle to upper Eocene retrogradational sediments deposited during intensified tectonism (Alavi, 2004). These sediments were deposited both as widespread carbonates and as flysch-system siliciclastics, with clastic supply waning toward the late Eocene

(Koop and Stoneley, 1982; Alavi, 2004).

317

In Thailand, most Cenozoic basin development occurred in the late Eocene through

Oligocene with the opening of a series of half grabens and grabens (Polachan et al., 1991;

Charusiri and Pum-Im, 2009). The early Eocene is poorly documented, with oil exploration wells rarely penetrating sediments older than the Oligocene exploration targets (Polachan et al., 1991; Charusiri and Pum-Im, 2009), however, some occurrence of Eocene-age carbonates is mentioned in the literature, such as the Eocene Tampur shallow marine carbonates (Andreason et al., 1997).

Rifting in the Atlantic-type South China Sea Basin began sometime in the latest Cretaceous or Paleocene, with sea-floor spreading active in the early and middle Eocene (Taylor and

Hayes, 1983; Ru and Pigott, 1986). Throughout much of the basin, syn-rift deposition was characterized by nonmarine facies. Offshore southern and central Vietnam, syn-rift sediments consist of nonmarine sediments of possible Eocene to Oligocene age (Lee and

Watkins, 1998; Lee et al., 2001b). Off northern Vietnam and southern China, syn-rift sediments of Eocene age consist of nonmarine (largely lacustrine) sediments (Chen et al.,

1993). In the Taiwan region, syn-rift sediments of late Paleocene to late Eocene age are comprised primarily of shales, with minimal sandstone beds and volcaniclastics intercalated, deposited largely in nonmarine systems (Lin et al., 2003; Huang et al., 2012).

Sediments in the central South China Sea Basin primarily consist of Oligocene and younger sediments atop oceanic basement (Li et al., 2015). Central Palawan Island of the Philippine island arc drifted from the Southern China block during opening of the South China Sea, and currently lies on the southeastern margin of the South China Sea (Holloway, 1982;

Suzuki et al., 2000). A Late Cretaceous to Eocene sedimentary section currently exposed on central Palawan Island consists of clastics likely derived from the continental material

318 of the South China block, and contains quartz- and volcaniclastic-rich sandstone turbidites

(along with massive sandstones, pebbly mudstones, and mudstones) inferred to have been deposited in deep-water submarine fans and basin plains (Suzuki et al., 2000). These sandstone beds are generally intercalated with mudstone, are medium- to fine-grained, and are some 5 to 20 cm in thickness (Suzuki et al., 2000). Syn-rift Eocene turbidites are also documented in southern onshore Palawan, where they occur as interbedded arkosic sandstone, mudstone, and siltstone, and are inferred to be present in some offshore half- grabens (Steuer et al., 2013; Aurelio et al., 2014).

Intracontinental basins in eastern China, basins of China’s eastern marginal seas (including the East China Sea Basin and Yellow Sea Basin), as well as the Sea of Japan off the Korean and Japanese coasts, developed as extensional basins in the Cenozoic (Ren et al., 2002).

These rift basins are largely filled with nonmarine (usually alluvial and lacustrine) sediments (Ren et al., 2002). For instance, the Paleocene through Eocene in the Yellow Sea

Basin contains pollen indicative of fluvial paleoenvironments (Yi et al., 2003). Outboard of East China Sea in the Shimanto Supergroup, however, mélange contains sedimentary rock deposited in a forearc basin setting including accretionary complex, trench-slope, and trench settings (Taira et al., 1982). Paleogene sediments here are primarily made up of turbidite sands and shales (Taira et al., 1982). On Okinawa Island, for example, Cretaceous to postulated lower Eocene accretionary complex mélange contains terrigenous sedimentary rocks including turbidites likely deposited in deep-water environments (Ujiie,

2002). These rocks are overlain by middle to upper Eocene trench fill turbidites (Ujiie,

2002). The rifted Sea of Japan and the rifted Okhotsk Sea off far eastern Russia both lack

319

Eocene sediment, with fill comprised of Oligocene and younger sediments (Chough and

Barg, 1987; Lee et al., 2001a; Baranov et al., 2002; Karp et al., 2006).

The Amerasia and Eurasia basins of the Arctic Ocean are rimmed by rifted continental margins (Franke et al., 2004). The age of rifting remains poorly constrained, though a late

Mesozoic/early Cenozoic timing is often invoked (e.g., possible initiation of rifting in the

Makarov/Podvodnikov basins of the Amerasia Basin in the Late Cretaceous; Alvey et al.,

2004; possible initiation of rifting in the Eurasia Basin in the late Paleocene to early

Eocene; Franke et al., 2004 and references therein). The East Siberian Shelf is virtually unexplored, is up to 800 km wide, and contains sediments deposited on a subsiding epicontinental platform continuously since the Late Cretaceous (Franke et al., 2004). Even less understood is the deep-water Amerasia Basin, which encompasses both the Canada and Makarov basins, and occupies water depths up to nearly 4 km (Franke et al., 2004).

On the northwestern portion of the East Siberian Shelf, bordering the Makarov Basin, sedimentary cover of Cretaceous and Cenozoic age is estimated at anywhere from tens of meters to up to 9 km in areas (Sekretov, 2001). Here, clinoforms of inferred Paleogene and early Miocene age are interpreted to represent a deltaic sequence consisting of shallow marine and nonmarine facies (Sekretov, 2001). Parallel reflectors at greater than 1 km water depth along the East Siberian Shelf continental slope are in this region postulated to perhaps indicate intercalated turbidite sandstones and clays of Late Cretaceous to Miocene age (Sekretov, 2001). The Laptev Shelf is also virtually unexplored, hundreds of kms wide, and contains thick successions of sediment of inferred Late Cretaceous to recent age

(Franke and Hinz, 2005). Here, numerous rift basins comprise a graben system filled with

Late Cretaceous to recent sediment in excess of 4 sec (TWT) of thickness, estimated to

320 represent from 4 to 10 km of sediment thickness (Drachev et al., 1998; Franke and Hinz,

2005).

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Appendix C-2.4. Australia/Oceania

Figure C-2.4A. Locations of Paleogene sedimentary deposits described for the Oceania region, except New Zealand (early Eocene paleomap after Scotese, 2001). Bo—Borneo,

PNG—Papua New Guinea, NC—New Caledonia, Ta—Tasmania, EB—Eucla Basin,

WA—western Australia, NW—Northwest Shelf, GP—Gulf of Papua, QP—Queensland

Plateau, CB—Capricorn Basin. Symbols as in Figure C-2.1.

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Figure C-2.4B. Locations of Paleogene sedimentary deposits described for New Zealand

(early Eocene paleomap after Scotese, 2001). RR—Rotorua and Raukumara regions, KC—

Kaikoura and Christchurch regions. Symbols as in Figure C-2.1.

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Table C-2.4A. Locations of lower Eocene turbidites associated with the Oceania region, as well as descriptions of deposition during the neighboring time periods.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin same as NW Borneo turbidites Paleocene volcanics ? Yes Yes same as early Eocene, but SE Borneo capped by Yes? (active (Kutai Basin) N/A turbidites? marine shale ? Yes (deltaic) rift) New Zealand (Rotorua, glauconitic Raukumara mudstone & same as same as same as regions) sandstone Paleocene Paleocene Paleocene No? No

New Zealand same as early (Kaikoura area) limestone turbidites Eocene marl No? No New Zealand (Kaikoura, mid-shelf glauconitic mix of early Christchurch sandstone/mud mudstone & and late shallow water regions) stone sandstone Eocene sandstone No? No

same as Paleocene, but transitioning to W Tasmania conglomerate, same as same as carbonate Yes? (active shelf sandstone, silt Paleocene Paleocene system No rift)

324

Table C-2.4B. Oceania region locations surveyed but lacking lower Eocene turbidites, including descriptions of deposition during the Paleocene through late Eocene.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin NE Papua New Guinea (Finisterre volcanic same as Yes? (volcanic Range) N/A N/A turbidites middle Eocene No arc) same as same as New Caledonia carbonate Paleocene turbidites/flysch middle Eocene No Yes S Australia same as (Eucla Basin) N/A nonmarine? carbonate middle Eocene No No carbonate/hiatu W Australia carbonate s hiatus ? No No NW shelf, same as same as same as Australia limestone/marl Paleocene Paleocene Paleocene No No NE Australia (Gulf of Papua) limestone? limestone limestone limestone? No ? (rifting) NE Australia calcareous (Queensland ooze/terrigenou same as Plateau) ? ? s sandstone middle Eocene ? ? (rifting) NE Australia (Capricorn same as same as same as Basin) claystone? Paleocene? Paleocene? Paleocene? ? ? (rifting)

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In northwestern Borneo, Late Cretaceous to lower Eocene deep-water marine turbidites

(Embalah Group/Rajang Group) crop out (Moss, 1998). Moss (1998) proposed that these strata were deposited as fill in a trapped remnant oceanic basin atop a subducting oceanic plate outboard of the Schwaner Mountains of Borneo. These turbidites are quartz-rich, though they contain lithic clasts, and are generally centimeters to slightly over a meter in thickness and are interbedded with siltstones and shales (Moss, 1998). To the southeast, the Kutai Basin is generally inferred to have initiated in the middle Eocene and contain fill no older than middle Eocene in age (Moss et al., 1997; Moss and Chambers, 1999), though a recent study suggests deep-water turbidites of early to middle Eocene age (Sambakung

Formation) were encountered in two exploration wells (Camp et al., 2009).

In Papua New Guinea, Eocene deposition is generally dominated by limestone, though some areas include shallow water sandstones (Davies and Smith, 1971; Home et al., 1990).

For instance, the Sarawaget beds of northeastern-most Papua New Guinea are composed of middle Eocene to lower Miocene volcaniclastics, with turbidite and conglomerate deposits inferred to have been deposited as clastic aprons adjacent to volcanic edifices of the Finisterre arc (Abbott et al., 1994).

In New Caledonia, the Paleocene through middle Eocene is dominated by micrites, marls, and limestone, although, in areas, middle to upper Eocene fine-grained turbidites of the

Bourail flysch are deposited in an active foreland basin setting (Maurizot, 2011; Maurizot and Cluzel, 2014).

Subsidence of most of New Zealand from the early Cenozoic to Oligocene meant high relative sea level for much of this time (Loutit and Kennett, 1981). Basins off the northeastern, eastern, and southeastern coasts of New Zealand, as well as onshore 326 stratigraphy, reflect a dominance of Paleogene mudstone and calcareous mudstone to limestone deposition along a passive margin during a time of transgression (Schiøler et al.,

2010; Bland et al., 2015; Kamp et al., 2015). Nonetheless, studies of onshore geology document minor glauconitic sandstone-mudstone intervals in the 57-37 Ma mid-bathyal to lower bathyal Wanstead Formation of the North Island’s Rotorua and Raukumara areas, turbiditic sandstones in the mid-bathyal Woodside Formation (coeval with the 55-41 Ma

Lower Marl of the Amuri Limestone) of South Island’s Kaikoura region, glauconitic sands in the late middle Eocene mid-bathyal Fells Greensand Member of the Kaikoura area, and glauconitic, calcareous mudstones and sandstones in the ~54-39 Ma lower bathyal Ashley

Mudstone of the Kaikoura and Christchurch areas (Morris, 1987; Kamp et al., 2015). This period of relative quiescence and passive margin-style deposition also dominated deposition along New Zealand’s western margin, including in the distal portions of the

Taranaki Basin, where marine mudstone was deposited in the Paleocene through Eocene

(Higgs et al., 2012).

From the middle Paleocene to middle Eocene, tectonism was limited around much of the

Australian continental landmass (Müller et al., 2000), which largely existed as a passive margin, as it does today (e.g., von Rad and Exon, 1983; Seidl et al., 1996). The Australian landmass itself was stationary with respect to the mantle from the Late Cretaceous through the late Eocene (Müller et al., 2000). Along the southern ocean, a carbonate shelf was established in the early Eocene following Paleocene transgression (Hobbs et al., 2019).

Ocean circulation to the south of Australia was restricted until the opening of the Tasman

Gateway at either 49-50 Ma (Bijl et al., 2013) or about 43 Ma (Müller et al., 2000). .

327

In the region of Tasmania, however, rifting between Australia and Antarctica began in the

Late Cretaceous and produced numerous rift basins (Exon et al., 2004). Paleocene and

Eocene sedimentation within the area is largely characterized by shelf and deep-water mudstones, except on the shelf of west Tasmania where Late Cretaceous to Eocene conglomerates, sands, and silts were deposited rapidly in a nearshore fault-bounded basin

(Exon et al., 2004). A general transition from siliciclastic to carbonate systems occurred in the late Eocene (Exon et al., 2004).

Southern Australia’s Eucla Basin contains middle Eocene and younger carbonate sediments, though adjacent paleovalleys may include some nonmarine (fluvial) sediments of early Eocene age (Hou et al., 2006). At the confluence of the Arctic and Indian Oceans near the Naturaliste Plateau, the Paleogene is dominated by carbonate ooze (Hobbs et al.,

2019).

Along Australia’s passive western margin, most Cenozoic sedimentation has consisted of carbonates (Quilty, 1977). Here, a late Paleocene to early Eocene depositional cycle occurs, peaking in the late Paleocene and tailing off into an early to middle Eocene depositional hiatus (Quilty, 1977).

Off Australia’s Northwest Shelf, the Late Cretaceous to present has been dominated by limestone and marl deposition, with the early Eocene interval characterized by both limestone and marl deposition in response to Paleocene transgression (Butcher, 1990;

Baillie et al., 1994).

The northeast of Australia was subject to Late Cretaceous through early Cenozoic rifting

(Davies et al., 1989). Eocene deposition in the Gulf of Papua consisted largely of

328 limestone, as temperatures during the Paleocene through middle Eocene throughout the northeast Australian margin were suitable for reef growth and stability and carbonate build- up (Davies et al., 1989; Feary et al., 1991). On the Queensland Plateau, calcareous ooze and terrigenous sand of middle and late Eocene age are present, while the Capricorn Basin potentially contains claystone of early Cenozoic age (Davies et al., 1989).

329

Appendix C-2.5. Europe

Figure C-2.5. Locations of Paleogene sedimentary deposits described for the European continental margin (early Eocene paleomap after Scotese, 2001). CB—Central Basin,

SB—Sørvestsnaget Basin, TH—Tromsø and Hammerfest basins, NS—North Sea, FSB—

Faroe-Shetland Basin, KB— Kangerlussuaq Basin, PB—Porcupine Basin, P—Pyrenees,

BB—Bay of Biscay, IM—Iberian margin, A—Alps. Symbols as in Figure C-2.1.

330

Table C-2.5A. Locations of lower Eocene turbidites associated with the European continent, as well as descriptions of deposition during the neighboring time periods.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin Spitsbergen/Sv albard (Central No (numerous Basin) clastics? submarine fans ? ? deltas) Yes SW Barents Sea (Tromsø prograding & Hammerfest (deltaic?) same as basins) sediment Paleocene ? ? No? Yes? N & central North Sea (adjacent Shetland Platform, between platform & Norwegian same as same as Yes (Shetland Shield) submarine fan Paleocene some fans middle Eocene No uplift)

same as North Sea Paleocene, but Viking Graben capped by late Yes (Shetland (Frigg fan) submarine fan Ypresian shale No uplift) Faroe- turbidite same as Yes? Shetland Basin sandstone limited fans? submarine fan? middle Eocene No? (subsidence) W Ireland delta-front (Porcupine sandstones & same as same as same as Yes? (rift Basin) submarine fans Paleocene Paleocene Paleocene No? basins) Southern Pyrenean submarine fan same as early foreland basin ? & turbidites Eocene ? No Yes Pyrenees (Ainsa, Pamplona, Basque basins, coast of Bay of Biscay) turbidites No Yes

offshore central W Iberia (Peniche hemipelagites same as same as Basin) & turbidites? Paleocene Paleocene ? No? Yes S Iberian margin (SE same as same as same as Spain) sandy turbidites Paleocene Paleocene Paleocene No? Yes W, central, & same as same as same as E Alps flysch Paleocene Paleocene Paleocene No Yes Carpathians (example: Poland Ciezkowice same as same as same as Sandstone) flysch Paleocene Paleocene Paleocene No Yes 331

Table C-2.5B. European locations surveyed but lacking lower Eocene turbidites, including descriptions of deposition during the Paleocene through late Eocene.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin transgression, tectonic quiescence, waning overall North sediment Yes (Shetland Sea siliciclastics supply No uplift) overall central North Sea finer-grained, (Shetland mud- Platform coarse dominated Yes (Shetland source area) siliciclastics deposition No uplift) SW Barents Sea (Sørvestsnaget same as Basin) marine shale Paleocene submarine fans shallow marine No? Yes? SE Greenland (Kangerlussuaq No? Basin) turbidites ? ? ? No? (extensional) offshore Bay of Biscay/French Yes Atlantic coast turbidites? No? (tectonism)

332

In Europe, the opening of the Atlantic Ocean (Pitman and Talwani, 1972) and the failed rifting of the North Sea were responsible for a major transgressive phase from the late

Paleocene to early Eocene, whereas the middle Eocene to early Oligocene represent a general regressive phase (Hardenbol et al., 1998).

In Spitsbergen along the Svalbard margin, transpression associated with opening of the

Norwegian-Greenland Sea led to Paleocene to early Eocene development of the Central

Basin (Steel et al., 1985; Grundvåg et al., 2014). In the late Paleocene through early

Eocene, the basin was filled with >1.5 km of clastic sediment, including Eocene submarine fans, deposited by high-supply prograding deltas fed by a rising orogenic belt despite potentially high sea level (Steel et al., 1985; Grundvåg et al., 2014).

The continental margin of Svalbard and the western Barents Sea extends some 1000 km at present (Ryseth et al., 2003). The Barents Sea also evolved in a transform setting during

Late Cretaceous through Paleogene rifting and early opening between Eurasia and

Greenland, and during opening of the Norwegian-Greenland Sea and accompanying

Eocene subsidence (Faleide et al., 1993, 2008). In the Barents Sea, basinal areas in the west and southwest were generally filled by Paleocene to lower Eocene marine shales, and record late Paleocene and early Eocene transgressive phases, while platform areas

(particularly the eastern Barents Sea) are devoid of Cenozoic strata (Vorren et al., 1991;

Faleide et al., 1993; Henriksen et al., 2011). In the southwestern Barents Sea in the

Sørvestsnaget Basin, early Paleocene to late Eocene transgressive deep-water conditions persisted, with thicker sections of Paleocene to lower Eocene shelfal marine shales present and overlain by middle Eocene deep-water strata, including middle Eocene sandy submarine fans, all of which is overlain by upper Eocene to Miocene condensed shallow

333 marine strata (Ryseth et al., 2003; Henriksen et al., 2011; Safronova et al., 2014). In the southwestern Barents Sea just to the southeast of Sørvestsnaget Basin, the Tromsø Basin and Hammerfest Basin just to its east both record late Paleocene to early Eocene progradation of sediment inferred to be derived from the Loppa High, and possibly associated with deltaic activity (Knutsen and Vorren, 1991; Knutsen et al., 1992).

Major marine flooding in the early Eocene is evident throughout the Norwegian Sea, likely in response to continental breakup and subsidence (Henriksen et al., 2005). This is manifested in the ubiquity of marine shales and paucity of sands throughout the Eocene section of the Norwegian Sea (Eidvin et al., 1998; Henriksen et al., 2005).

The North Sea Basin experienced extension and widened by some 70 km during Middle

Jurassic to Early Cretaceous time, and since has experienced prolonged thermal subsidence

(Barton and Wood, 1984). During the early Eocene, the North Sea experienced transgression consistent with global sea-level rise which, in conjunction with relative tectonic quiescence in the Eocene, resulted in an Eocene waning of sediment supply to the basin (Neal, 1996; Anell et al., 2012). Depositional patterns in the Eocene northern North

Sea were consistent with those of the Paleocene, whereby the majority of Eocene sediment was sourced from the Shetland Platform to the west (>800 m Eocene sediment thickness) and Sognefjorden to the east (>400 m thickness) (Anell et al., 2012). Sediment thicknesses in the central North Sea exceed 1000 m (Anell et al., 2012). In the Shetland Platform source area, Late Cretaceous uplift created topographic highs that contributed to coarse siliciclastic deposition in the Paleocene and continued albeit finer-grained and mud- dominated deposition in the Eocene (Harding et al., 1990). Adjacent Shetland Platform and between the platform and the Norwegian Shield (in the northern and central North Sea),

334 delta fan complexes developed and prograded southeastward during the Paleocene through early Eocene (presumably in response to the uplifted Shetland Platform source area to the northwest and west), though fans were also active in the middle and late Eocene (Heritier et al., 1979; Morton, 1982; den Hartog Jager et al., 1993). Despite marine transgression in the Ypresian, thick sands were deposited in a number of fan complexes including, most notably, the deltaic and delta-front Frigg submarine fan consisting of especially sand-rich sediment of latest Paleocene to late Ypresian age in Viking graben (Heritier et al., 1979;

McGovney and Radovich, 1985). The Frigg fan is of presumed deep-water (Heritier et al.,

1979) or shallow marine (Morton, 1982) origin and is inferred to have predominantly been deposited during sea level fall, such as in the late Ypresian (McGovney and Radovich,

1985). The Frigg fan is capped by late Ypresian marine shale (Heritier et al., 1979). Slightly to the south and adjacent Shetland Platform, the Alba fan system of Witch Ground graben was active in the middle Eocene in response to the uplifted source area and lower sea levels, however, presence of sand in the sediment-starved early to early-middle Eocene is relatively scarce (Harding et al., 1990). Nonetheless, a deltaic pre-Alba sandy system in

Witch Ground graben may represent the southernmost Viking graben fan complex

(Harding et al., 1990).

The Faroe-Shetland Basin between the Faroe Islands and Shetland Islands developed in response to proto-Atlantic Ocean extension, and saw early Eocene deltaic (non-marine) sedimentation alongside periods of shallow marine deposition (Lamers and Carmichael,

1999; Stoker et al., 2013). Instances of lower Cenozoic turbiditic sands (along with commercial exploration successes in the Faroe-Shetland Basin) are primarily restricted to the Paleocene in both the Faroe-Shetland Basin and the Kangerlussuaq Basin of southern

335

East Greenland (Grant et al., 1999; Lamers and Carmichael, 1999; Larsen et al., 2006;

Egbeni et al., 2014), though latest Paleocene to late Eocene Faroe-Shetland Basin subsidence likely led to deposition of sand-rich submarine fans in the middle to late Eocene

(Lamers and Carmichael, 1999; Davies et al., 2004), and limited deposition of fans in the early Eocene is possible (Sørensen, 2003). In offshore western Ireland to the south, the

Porcupine Basin similarly represents a series of failed proto-Atlantic Ocean borderland rift basins (Croker and Klemperer, 1989; Moore and Shannon, 1992). The main basin contains some 6 km of Cretaceous to Recent sediments deposited during post-rift subsidence

(Croker and Klemperer, 1989; Moore and Shannon, 1992). Four Paleocene to Eocene southerly prograding deltaic sequences, each capped by marine transgressive successions, occur in the basin and result in deposition of delta front sandstones and submarine fans of this age (Moore and Shannon, 1992; Shannon, 1992).

From the Late Cretaceous through most of the Cenozoic, North-South convergence between the Iberian and Eurasian plates led to development of the West-East trending

Pyrenees mountain chain (Labaume et al., 1985). Development of foreland basins to the north and to the south of the Pyrenees accompanied this convergent tectonism (Labaume et al., 1985). In the Southern Pyrenean foreland basin, which developed from the late

Paleocene through Miocene, tectonic activity contributed to deposition of the lower and middle Eocene Hecho Group submarine fan and basin plain turbidites (Mutti, 1977, 1983;

Labaume et al., 1985). The Hecho Group turbidite system is a delta-fed clastic prism some

175 km long, 40-50 km wide, and some 3.5 km thick (Mutti, 1977, 1983). The lower Hecho

Group is interpreted to have been deposited during relative lowering of sea level (Mutti,

1983). Overall, instances of tectonically-mediated early Eocene turbidite deposition are

336 prevalent throughout the Southern and Northern Pyrenean foreland basins and associated basins, including the Ainsa, Pamplona, and Basque basins, and coast of the Bay of Biscay, though also of note is documented influence of Milankovitch cyclicity and climate on deposition (Gawenda et al., 1999; Payros et al., 2006; Pickering and Bayliss, 2009; Payros and Martínez‐Braceras, 2014). In the offshore Bay of Biscay and along the French Atlantic margin, deposition of turbidites is inferred to postdate Eocene tectonism in the region, meaning deposition began in the late Eocene (Bourillet et al., 2006).

Breakup of Pangaea in the earliest Jurassic led to the opening of the Atlantic Ocean Basin

(Dietz and Holden, 1970), with rifting of the west Iberian margin (northeast Atlantic) taking place from the Late Jurassic through Cretaceous (Alves et al., 2003). The rifted

Iberian margin existed as a subsiding passive margin until Eocene deformation as a result of the Pyrenean orogeny (see above), at which time blocks along the northwestern Iberian margin were uplifted by up to several kilometers (Boillot et al., 1979; Munoz et al., 2003).

One might infer that this tectonism and uplift contributed to offshore sedimentation in the region, however, data on lower Eocene sediments in the margin’s offshore region is scarce.

In the case of the Galicia Interior Basin (northwest Iberia) (Murillas et al., 1990),

Lusitanian Basin (central western Iberia) (Wilson et al., 1989), and Alentejo Basin

(southwest Iberia) (Alves et al., 2003), lower Eocene sediment seems simply undocumented or not present. However, in the Peniche Basin of offshore central western

Iberia, the Late Cretaceous through middle Eocene is inferred, based on seismic facies and well ties, to be comprised of hemipelagites and turbidites (Alves et al., 2006). As along the western Iberian margin, far-field effects of the Pyrenean collision manifested in platform

337 emergence, which resulted in deposition of sandy turbidites on the southern Iberian margin

(southeast Spain) from the Paleocene through Eocene (De Ruig et al., 1991).

Much Cretaceous and Cenozoic deposition along the northern Tethyan (Helvetic, or present-day southern European) margin was influenced by collisional processes and complex tectonism related to the Alpine orogeny (Schmid et al., 2008). In the western and central Alps, Cenomanian to lower Oligocene flysch was deposited (Trümpy, 1960).

Similarly, the eastern Alps contain flysch of this age, including examples of Cretaceous and Eocene flysch (Hesse, 1975; Faupl and Wagreich, 2000). Cretaceous to Oligocene flysch is also encountered in the Carpathians, and includes such examples as the upper

Paleocene to Eocene Ciezkowice Sandstone of Poland, which is interpreted to represent turbidites of a basin-floor fan (Dziadzio et al., 2006).

338

Appendix C-2.6. India

Figure C-2.6. Locations of Paleogene sedimentary deposits described for the Indian continental margin (early Eocene paleomap after Scotese, 2001). KF—Khojak submarine fan, BF—Bengal submarine fan, WI—western India (Kutch and Cambay basins), EI— eastern India (rift basins). Symbols as in Figure C-2.1.

339

Table C-2.6A. Location of lower Eocene turbidites associated with the Indian subcontinent, as well as descriptions of deposition during the neighboring time periods.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin major increase Bengal in same as submarine fan N/A fan inception? sedimentation middle Eocene Yes? Yes

Table C-2.6B. Indian locations surveyed but lacking lower Eocene turbidites, including descriptions of deposition during the Paleocene through late Eocene.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin Afghanistan (N Katawaz Basin), Khojak fan Khojak fan? ? ? ? Yes Yes Pakistan (Katawaz Basin) & Iran (Makran) N/A N/A N/A Khojak fan Yes Yes shallow marine W India (Kutch shale & Basin) limestone No No W India (Cambay Basin) black shale No No

no Eocene rock exposed onshore, and E India rift only postulated same as early same as early basins offshore Eocene Eocene No No

340

Much as today, Paleogene turbidite deposition associated with the Indian subcontinent was dominated by sediments shed off the uplifting Himalaya as a result of the India-Asia collision. The precursors of the two largest modern sedimentary systems on Earth, the

Indus and Bengal fan systems, saw their inception during this early Himalaya uplift

(Qayyum et al., 1996; Curray et al., 2003; Ingersoll et al., 2003).

In the northwest of India, deposition occurred predominantly in the form of the Khojak submarine fan, a Paleogene analogue for the modern Indus delta-fan complex (Qayyum et al., 1996, 1997b; Carter et al., 2010). The Katawaz Basin of Pakistan and the Makran region of Iran contain delta-flysch sediments preserving deposition via a paleo-Indus river, by which both the Katawaz Delta and Khojak submarine fan sediments were delivered to the

Neo-Tethys Ocean from the uplifting, eroding Himalaya (Qayyum et al., 1996, 1997b,

2001; Carter et al., 2010). The middle and upper sediments of the Khojak fan are assigned a late Eocene to early Miocene age, although upper Paleocene strata of the Khojak are found in the northern Katawaz Basin in Afghanistan, indicating the fan may have been active then (Qayyum et al., 1997a, 2001). Deposition of Himalaya-derived sediments in the Indus submarine fan of the Indian Ocean did not begin until the early Miocene (Davies et al., 1995; Qayyum et al., 1997b).

In the northeast, Bengal submarine fan sedimentation may have begun by the earliest

Eocene in response to uplift of the eastern Himalaya to the north and the Indo-Burman

Ranges to the east (Alam et al., 2003; Curray et al., 2003; Curray, 2014). Evidence for this is present in Indoburman flysch preserved as offscraped sediments in the Indoburman

Ranges (Curray et al., 2003). However, a major increase in Bengal fan deep-water

341 sedimentation began in the middle Eocene as a result of the collision of India with the

Burma and Tibetan blocks (Alam et al., 2003).

Aside from the above two cases of active margin deep-water sedimentation, there is a paucity of evidence for any turbidite deposition around the Indian subcontinent in the early

Eocene. Overall, Paleogene sedimentation rates in the Indian Ocean are inferred to have been quite low (Davies et al., 1995).

India’s west coast continental margin saw slow and long-lived passive margin subsidence from the Mesozoic breakup of Gondwana to the late Oligocene (Whiting et al., 1994). Rift basins developed along the western passive margin, but contain relatively thin Cenozoic sections (a few hundred meters) (Biswas, 1982). Furthermore, the early Eocene in these basins is characterized by marine transgression, and either shallow marine shale and limestone deposition (Kutch Basin) or black shale deposition (Cambay Basin) (Biswas,

1982).

A number of rift basins also developed post-Gondwana breakup along the eastern coast of

India (Sastri et al., 1981). However, basins on India’s east coast lack onshore exposure of

Eocene rock, while offshore sections are only speculated to contain marine Eocene deposits, and are inferred to contain much thicker sections of Neogene sediment (Sastri et al., 1973; Bastia, 2006).

342

Appendix C-2.7. North America

Figure C-2.7. Locations of Paleogene sedimentary deposits described for the North

American continental margin (early Eocene paleomap after Scotese, 2001). SJB—San

Joaquin Basin, SB—Sacramento Basin, BB—California borderlands basins, YC—Yager complex, TB—Tyee Basin, Wa—Washington, VQ—Vancouver and Queen Charlotte islands, GA—Gulf of Alaska, BMC—Beaufort-Mackenzie and Canada basins, SvB—

Sverdrup Basin, LS—Labrador Sea, JA— Jeanne d’Arc Basin, ScB—Scotian Basin and shelf, BH—Baltimore Canyon Trough and Hatteras Basin, GM—Gulf of Mexico, Cu—

Cuba, PB—Peralta Belt, Ba—Barbados, VB—Venezuela Basin, CB—Colombia Basin,

BeB—Belize Basin, MB—Mosquitia Basin, SFB—Sandino forearc basin, LB— Limón

Basin. Symbols as in Figure C-2.1.

343

Table C-2.7A. Locations of lower Eocene turbidites associated with the North American continent, as well as descriptions of deposition during the neighboring time periods.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin Cantua Sandstone (submarine fan encased in San Joaquin early Eocene Basin ? Lodo Shale) ? ? ? Yes Sacramento Basin N/A turbidites fewer turbidites fewer turbidites Yes Yes submarine fans common in early borderlands Cenozoic; basins basins have (Gualala, La thick early Honda, Sierra Cenozoic Madre, Santa marine Ynez basins ) sediment fill No Yes borderlands (La Honda Basin, Santa Cruz Butano same as early same as early Mountains) unconformity Sandstone Eocene Eocene No Yes borderlands (N nonmarine/shall Santa Lucia deepwater The Rocks same as early ow marine to Range) Lucia Mudstone Sandstone Eocene deep marine No Yes N California far fewer Yager complex turbidites? turbidites fewer turbidites fewer turbidites Yes Yes Oregon Tyee Basin volcanics turbidites fewer turbidites fewer turbidites Yes Yes same as same as n Gulf of AK submarine fan Paleocene Paleocene siliciclastics ? Yes N Gulf of Alaska (Kodiak Island) ? turbidites? turbidites? turbidites? ? Yes S Beaufort Sea (Beaufort- Mackenzie prodeltaic same as same as same as Basin) turbidites Paleocene Paleocene Paleocene? Yes No Beaufort Sea/Arctic Ocean (S same as same as same as Canada Basin) clastics Paleocene Paleocene Paleocene Yes? No waning W & S Jeanne deltaic & same as sediment/shale same as d’Arc Basin turbidites Paleocene ? middle Eocene Yes? No E & NE waning Jeanne d’Arc same as sediment/shale same as Basin submarine fans Paleocene ? middle Eocene ? No same as waning same as Gulf of Mexico turbidites Paleocene sediment middle Eocene Yes No* W Cuba (Los clastics/carbon Palacios Basin) ? ate ? ? No Yes same as central Cuba turbidites Paleocene ? 344 ? ? Yes SE Hispaniola (SE Peralta Belt) ? turbidites limestone ? No? Yes same as early same as early E Barbados ? turbidites Eocene Eocene No? Yes S Belize/E Guatemala same as same as same as Yes? (Belize Basin) turbidites Paleocene Paleocene? Paleocene? (tectonism) W Nicaragua (coastal & offshore volcanic Sandino same as early turbidites / same as early same as early forearc basin) Eocene? submarine fan Eocene? Eocene? ? Yes E Costa Rica volcanic same as carbonate/turbi same as (Limón Basin) turbidites Paleocene dites middle Eocene ? Yes Location Paleocene early Eocene middle Eocene late Eocene River Active margin Cantua Sandstone (submarine fan encased in San Joaquin early Eocene Basin ? Lodo Shale) ? ? ? Yes Sacramento Basin N/A turbidites fewer turbidites fewer turbidites Yes Yes submarine fans common in early borderlands Cenozoic; basins basins have (Gualala, La thick early Honda, Sierra Cenozoic Madre, Santa marine Ynez basins ) sediment fill No Yes borderlands (La Honda Basin, Santa Cruz Butano same as early same as early Mountains) unconformity Sandstone Eocene Eocene No Yes borderlands (N nonmarine/shall Santa Lucia deepwater The Rocks same as early ow marine to Range) Lucia Mudstone Sandstone Eocene deep marine No Yes N California far fewer Yager complex turbidites? turbidites fewer turbidites fewer turbidites Yes Yes Oregon Tyee Basin volcanics turbidites fewer turbidites fewer turbidites Yes Yes same as same as n Gulf of AK submarine fan Paleocene Paleocene siliciclastics ? Yes N Gulf of Alaska (Kodiak Island) ? turbidites? turbidites? turbidites? ? Yes S Beaufort Sea (Beaufort- Mackenzie prodeltaic same as same as same as Basin) turbidites Paleocene Paleocene Paleocene? Yes No Beaufort Sea/Arctic Ocean (S same as same as same as Canada Basin) clastics Paleocene Paleocene Paleocene Yes? No waning W & S Jeanne deltaic & same as sediment/shale same as d’Arc Basin turbidites Paleocene ? middle Eocene Yes? No E & NE waning Jeanne d’Arc same as sediment/shale same as Basin submarine fans Paleocene ? middle Eocene ? No same as waning same as Gulf of Mexico turbidites Paleocene sediment middle Eocene Yes No* W Cuba (Los clastics/carbon Palacios Basin) ? ate ? ? No Yes same as central Cuba turbidites Paleocene ? ? ? Yes SE Hispaniola (SE Peralta Belt) ? turbidites limestone ? No? Yes same as early same as early E Barbados ? turbidites Eocene Eocene No? Yes S Belize/E Guatemala same as same as same as Yes? (Belize Basin) turbidites Paleocene Paleocene? Paleocene? (tectonism) W Nicaragua (coastal & offshore volcanic Sandino same as early turbidites / same as early same as early forearc basin) Eocene? submarine fan Eocene? Eocene? ? Yes E Costa Rica volcanic same as carbonate/turbi same as (Limón Basin) turbidites Paleocene dites middle Eocene ? Yes

*indicates passive margin, but local tectonics

Table C-2.7B. North American locations surveyed but lacking lower Eocene turbidites, including descriptions of deposition during the Paleocene through late Eocene.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin volcanic/siltston volcanic/turbidit same as Washington volcanic e es middle Eocene Yes

Vancouver/Que little Eocene or en Charlotte Oligocene rock islands region characterized Yes S & central same as Gulf of Alaska N/A N/A submarine fan middle Eocene ? Yes early Cenozoic deposition is mostly nonmarine; early Cenozoic Bering Sea onward, flysch (Aleutian deposition Basin, N occurs in basins, etc.) Aleutian Basin ? Yes Arctic Canada (Sverdrup same as same as same as Basin) nonmarine Paleocene Paleocene Paleocene ? No* East Greenland rift margin ? ? clays ? ? No* W 345 Greenland/E Canadian shelf same as early same as early (Labrador Sea) ? mudstone Eocene Eocene No No Scotian Basin same as same as same as & Shelf mudstone Paleocene Paleocene Paleocene No? No Baltimore Canyon Trough/Hattera dearth of same as same as same as s Basin clastics Paleocene Paleocene Paleocene No? No

N-central Cuba foredeep basin ? carbonate ? ? No? Yes? Venezuela Basin (Caribbean pelagic Sea) ? sediment ? ? No No? central Colombia Basin (Caribbean pelagic Sea) ? sediment ? ? No No? NE Honduras/N Nicaragua (Mosquita same as same as same as Basin) red beds Paleocene Paleocene Paleocene Yes Yes Location Paleocene early Eocene middle Eocene late Eocene River Active margin volcanic/siltston volcanic/turbidit same as Washington volcanic e es middle Eocene Yes

Vancouver/Que little Eocene or en Charlotte Oligocene rock islands region characterized Yes S & central same as Gulf of Alaska N/A N/A submarine fan middle Eocene ? Yes early Cenozoic deposition is mostly nonmarine; early Cenozoic Bering Sea onward, flysch (Aleutian deposition Basin, N occurs in basins, etc.) Aleutian Basin ? Yes Arctic Canada (Sverdrup same as same as same as Basin) nonmarine Paleocene Paleocene Paleocene ? No* East Greenland rift margin ? ? clays ? ? No* W Greenland/E Canadian shelf same as early same as early (Labrador Sea) ? mudstone Eocene Eocene No No Scotian Basin same as same as same as & Shelf mudstone Paleocene Paleocene Paleocene No? No Baltimore Canyon Trough/Hattera dearth of same as same as same as s Basin clastics Paleocene Paleocene Paleocene No? No

*indicates passive margin, but local tectonics

346

Appendix C-2.7.1. Northern America

The Cordilleran margin, including the region between Baja California and the Gulf of

California in the south to Kodiak Island of Alaska in the north, was an Andean-type margin from the late Mesozoic to early Cenozoic (Dickinson, 1976). This subduction continued through to the end of the Eocene, when the Farallon Plate came into contact with the North

American Plate and convergence slowed in the late Eocene through Oligocene (Bartow,

1991). Motion along the modern San Andreas fault system initiated ~28 Ma (with suggestions of Paleogene slip along a San Andreas system or precursor fault system being controversial—and, recently demonstrated by Gooley et al., 2020, to be unnecessary), with significant implications for the tectonic development of western North America and the basins of the California margin (Atwater, 1989; Sharman et al., 2013; Gooley et al., 2020).

The Great Valley of California developed in the late Mesozoic and early Cenozoic as a forearc basin open to the Pacific Ocean to the west, and in the later Cenozoic as a transform margin basin (Dickinson and Seely, 1979; Ingersoll, 1979; Bartow, 1991). The valley is at present some 700 km long and up to 100 km wide and is flanked to the east by the Sierra

Nevada batholith and the west by the Coast Ranges (Bartow, 1991). Deposition in the Great

Valley forearc from the Late Cretaceous through early Paleogene was largely due to denudation of the high-standing Sierra Nevada (Mulch et al., 2006; Sharman et al., 2015), and deposition in the Great Valley forearc as well as forearc regions of the Baja California

Peninsular Ranges and the Oregon forearc during early to middle Eocene time was in large part due to input from five adjacent fluvial systems (Sharman et al., 2015). The Great

Valley of California is divided into the San Joaquin Basin to the south, and the Sacramento

Basin to the north (Bartow, 1991). The southern portion of the San Joaquin Basin

347 experienced late Paleocene to early Eocene transgression due to both overall sea level rise and basinal tilting (Bartow, 1991). A strong northward shoaling sequence characterizes the upper Paleocene to lower Eocene, while a strong regression-induced, basin-wide unconformity separates the lower Eocene and overlying Eocene sequence and marks dramatic lowering of sea level (Bartow, 1991). Maximum Cenozoic transgression in the basin was attained in the middle Eocene (Bartow, 1991). Early Eocene deposition in the basin includes local examples of deep-sea fan systems sourced predominantly from the

Sierra Nevada and/or Mojave (Dickinson et al., 1979; Sharman et al., 2013, 2015). One such example is the lenticular lower Eocene Cantua Sandstone, up to 1,000 m in thickness and fed by a proximal submarine canyon located in a tectonically active, structurally- controlled localized area (the “borderland” area described below) that sourced sediment from the Sierra Nevada (Graham and Berry, 1979; Sharman et al., 2015). In the Sacramento

Basin to the north, the Princeton Submarine Valley system developed in the Cenozoic

(middle Paleocene or earliest Eocene through Miocene) as one of the largest submarine valleys, modern or ancient (Redwine, 1972; Dumitru et al., 2013). This valley system is filled with lower Eocene (~51-49.4 Ma) sandstone, conglomerate, and mudstone of the

Capay Formation, which was deposited during transgression initiated in the Paleocene and lasting to the beginning of the middle Eocene, possibly as sediment input waned (Redwine,

1972; Graham, 1981; Dumitru et al., 2013). These marine sediments are primarily turbiditic in nature, and the southern end of the valley system likely includes at least one deep-sea fan (Redwine, 1972). Deposition of these and younger sediments is inferred to be tied to subaerial erosion by an ancestral Sacramento River (the Princeton River), with a significant amount of detritus likely sourced from Idaho batholith region highlands due to extensional

348 tectonic activity (Redwine, 1972; Dumitru et al., 2013). The Princeton River first delivered

Idaho Batholith sediments to the Sacramento Basin in the Ypresian (at ~53 Ma), and continued to do so until the late Eocene (~37 Ma) (Dumitru et al., 2013).

Early Cenozoic “borderland” basins of central California developed in response to tectonic activity (Nilsen and Clarke, 1975). These pull-apart basins are juxtaposed with adjacent upland source areas of Salinian block affinity (Nilsen and Clarke, 1975). These deep basins were therefore fed by proximally derived sediments at high rates, meaning areally restricted deep-sea fans are a common feature in the lower Cenozoic fill of these basins, and thick, deep marine sedimentary sections characterize the major early Cenozoic basins, including the Gualala, La Honda, Sierra Madre, and Santa Ynez basins (Nilsen and

Clarke, 1975). The borderlands basins were in some cases continuously fed by deep-sea fans from the Late Cretaceous through late Eocene, and in other cases over shorter time periods (Nilsen and Clarke, 1975). Notable cases of borderland basin deep-sea fan deposits include the lower to upper Eocene Butano Sandstone of the Santa Cruz Mountains La

Honda Basin (Nilsen and Simoni, 1973; Nilsen, 1985), as well as the late lower Eocene to middle Eocene The Rocks Sandstone of the northern Santa Lucia Range (Link and Nilsen,

1980).

The active trench of early Cenozoic offshore northern California influenced a succession of trench-slope and slope-basin sediments deposited in the Paleocene to Oligocene

(Atwater and Molnar, 1973; Underwood, 1983). This interval of sedimentation is dominated by deep-water sand-rich turbiditic deposition indicative of submarine canyons and channels and corresponding fans, with a general trend of sandier systems to the south

(Underwood, 1983; Bachman et al., 1984). These deep-water systems (the faulted and

349 folded marine sandstones and mudstones of the Yager complex) persisted in the Paleogene, and may have been present from the Paleocene to Oligocene, though are deemed to be likely early to late Eocene (~56-36 Ma) in age, with most deposition concentrated between

~53 and 49 Ma (Underwood, 1983; Dumitru et al., 2013). Much of this sediment is inferred to have been delivered from the Idaho batholith highlands via the Princeton River (Dumitru et al., 2013).

The Tyee Basin of western Oregon emerged as a new forearc basin system in the early

Eocene, and was quickly filled with 1.5 to 2 km of upper Ypresian to lower Lutetian (49.4-

46.5 Ma) sand-rich turbiditic sediments of the Tyee Formation (Chan and Dott, 1983;

Dumitru et al., 2013). Like the Capay Formation of the Princeton Submarine Valley system, these sandstones were predominantly derived from erosion of Idaho Batholith highlands (Dumitru et al., 2013). Sediment was delivered via the Tyee River from the late

Ypresian (~49 Ma) until the late Eocene (~36 Ma), and was deposited in a major delta to fan to basin plain system, though turbiditic sediment in the Tyee Basin may be as old as

~54 Ma (Chan and Dott, 1983; Dickinson et al., 1988; Dumitru et al., 2013; Santra et al.,

2013).

Paleocene to lower middle Eocene rocks of Washington are predominantly volcanic in nature, with lower Eocene siltstone also present, though deltaic and marine turbidite sands are present only beginning in the middle Eocene (Armentrout and Suek, 1985; Snavely,

1987; Snavely and Wells, 1996).

The shift from orthogonal convergence to strike-slip motion between the Pacific and North

American plates took place in the region of the present-day Vancouver and Queen

Charlotte islands, and the Gulf of Alaska, in the late Eocene (Rohr and Dietrich, 1992). 350

However, little Eocene or Oligocene rocks have been characterized in the Vancouver and

Queen Charlotte islands region, either onshore or offshore (Rohr and Dietrich, 1992). The

Eocene Tofino Basin landward of Vancouver Island may contain some (late) Eocene deep- water sediment deposited during a period of uplift and volcanism, though sand is predicted to be quite rare (Tiffin et al., 1972; Bustin, 1995). The Queen Charlotte Basin to the north is younger in age (Rohr and Dietrich, 1992).

In the southern and central Gulf of Alaska, Cenozoic sedimentation atop the Pacific Plate has been characterized by deposition of three sizeable submarine fan bodies (Stevenson and Embley, 1987). However, the oldest of these fans post-dates the early Eocene, as it is some 42-24 Myr in age (Lutetian at oldest), and may be markedly younger (Hamilton,

1967; Stevenson and Embley, 1987). The northern Gulf of Alaska is characterized by a complex Mesozoic and early Cenozoic tectonic history, with much Paleocene to Eocene sediment complexly deformed (Plafker, 1987). The Orca Group of the northern Gulf of

Alaska region (near Prince William Sound) represents a deep-water submarine fan complex older than fan complexes of the central and southern Gulf of Alaska (Plafker, 1987;

Stevenson and Embley, 1987). The Orca Group submarine fan complex was deposited as trench sediment in the late Paleocene through middle Eocene, and is overlain by upper

Eocene to Quaternary siliciclastic sediment deposited in subsiding shelf and slope basins

(Winkler, 1976; Plafker, 1987). In the northern Gulf of Alaska on Kodiak Island, the

Sitkalidak Formation of Eocene age also consists of deep-water turbidites deposited as trench sediment (Bradley et al., 2003). Eocene sediments from the northern Gulf of Alaska reflect deposition in a tectonically active trench setting, with sediment compositions

351 evolving from earlier magmatic-arc-derived sediment to subduction-complex-derived sediment (Bradley et al., 2003).

The Bering Sea Basin lies north of the Aleutian Trench, which since the Late Cretaceous has represented the boundary of northward-directed subduction of the Pacific Plate beneath the North American (Cooper et al., 1976), and north of the Aleutian Ridge which arose around the earliest Cenozoic (Scholl et al., 1975). Late Cretaceous to early Cenozoic deposition in the Bering Sea was predominantly terrigenous (nonmarine), both in the present-day deep-water Aleutian Basin toward the south (north of the Aleutian Ridge and south of the Bering Shelf) (Scholl et al., 1975) and the northern basins of the Bering Shelf

(Fisher et al., 1982). In the Aleutian Basin, flysch-like terrigenous deposits shed from the

Bering Shelf ponded north of the Aleutian Ridge in the early Cenozoic onward (Scholl et al., 1975), however, little is known about early Cenozoic deposition in general due to a paucity of drilling and deep cores (e.g., Cooper et al., 1987; Borreggine et al., 2017).

Deepening due to subsidence in the Miocene led to marine deposition throughout the region

(Scholl et al., 1975; Fisher et al., 1982).

In the Chukchi and Beaufort seas off the coast of northern Alaska, little information is available on deep-water sediments of Eocene age as very little petroleum exploration has occurred in the Arctic in general (Gautier et al., 2009). The continental shelf in this region resembles a rifted passive margin of Atlantic type, with rifting estimated at Cretaceous or earlier (Grantz et al., 1975, 1979; Lawver and Scotese, 1990). In the Chukchi and Beaufort seas, three offshore sedimentary basins (Hope, North Chukchi, and Camden basins) are identified and inferred to contain a kilometer to several kilometers of Paleogene marine and nonmarine clastic sediment (Grantz et al., 1975). In the southern Beaufort Sea, the

352

Beaufort-Mackenzie Basin stretches from the modern shelf edge landward, and contains prograding deltaic cycles deposited from the Paleocene to modern (Willumsen and Cote,

1982). Here, the Eocene Reindeer Delta fed into the basin, and resulted in deposition of prodelta sandy turbiditic sediments from the late Paleocene through middle to late Eocene

(Willumsen and Cote, 1982). Offshore, the southern Canada Basin of the Beaufort

Sea/Arctic Ocean is inferred to contain Cretaceous to Holocene clastics (Tailleur, 1970).

In Arctic Canada, the latest Cretaceous to Cenozoic Eurekan orogeny and accompanying tectonism meant deposition in the Sverdrup Basin shifted from some 260 Myr of continuous subsidence and accumulation to widespread nonmarine deposition from the

Late Cretaceous through Miocene and possible Pliocene (Balkwill, 1978). On the polar margin of Canada and Greenland, this notable tectonic influence on deposition persisted from 65 to 47 Ma (early Danian to earliest Lutetian) (Harrison et al., 1999).

Opening of the northern portion of the North Atlantic Ocean initiated in the early Eocene

(Talwani and Eldholm, 1977; Larsen, 1980; Hamann et al., 2005), and resultant rifting of

Greenland and Northern Europe resulted in widespread periods of magmatism and emplacement of flood basalts along the East Greenland rift margin from 62 to 47 Ma

(Paleocene through Ypresian) (Tegner et al., 1998). Post-volcanic onshore uplift paired with offshore subsidence and passive margin development led to deposition of an Eocene-

Oligocene succession offshore East Greenland that is characterized by a 500 m to 2 km thick megasequence of continuous, primarily parallel, high-amplitude reflectors (Larsen,

1980; Surlyk et al., 1981; Hamann et al., 2005). Ypresian terrigenous muds grade into

Lutetian pelagic oozes as a result of changing sediment source and decreasing sedimentation rates in a core hole from the Norwegian-Greenland Sea (though rates are

353 more stable at other sites) (Eldrett et al., 2004), while along the East Greenland margin maximum ages penetrated are middle Eocene massive and laminated silty clays and clays

(Berger and Jokat, 2008). Glaciogenic control on deposition on the margin is documented beginning in the late Eocene (Eldrett et al., 2007).

Rifting of Greenland from Baffin Island as a result of extended Labrador Sea spreading led to opening of Baffin Bay in the Eocene (Beh, 1975). Deposition in the newly formed Baffin

Bay ocean basin consisted of a maximum of 600 m of Upper Cretaceous and lower

Cenozoic shale and sand, but more than 6 km of upper Eocene and younger fluvio-deltaic sediments (Beh, 1975). To the south, opening of the Labrador Sea (spreading between

Greenland and Baffin Island in the north, Greenland and Labrador in the south) occurred at an earlier point, with the northern Labrador Sea opening in Paleocene/Eocene time and the southern Labrador Sea opening in Campanian/Maastrichtian to Eocene time (Gradstein and Srivastava, 1980). This opening corresponded with latest Cretaceous through Eocene transgression, with bathyal conditions prevailing in the Eocene and deposition of terrigenous fine-grained sediment prevailing until the Eocene/Oligocene boundary

(Gradstein and Srivastava, 1980). Overall, deposition along the West Greenland and East

Canadian (Labrador) shelf during the Eocene was characterized by deposition of mudstone

(McWhae et al., 1980; Schenk, 2011).

Though breakup of Pangaea and resultant opening of the Atlantic Ocean Basin began in the Early Jurassic (Dietz and Holden, 1970; Hames et al., 2000; Bartolini and Larson,

2001), rifting between the central Grand Banks and Iberia, and within the Jeanne d’Arc

Basin, continued until at least the Early Cretaceous (late Aptian to early Albian) (Driscoll et al., 1995). Thermal subsidence persisted from the end of the Early/beginning of the Late

354

Cretaceous onward, meaning eustatic sea-level rather than tectonism influenced development of stratigraphy during this period (De Silva, 1993). Paleocene fill of the

Jeanne d’Arc Basin consists of deltaic sands with associated distal turbidite deposits, however, sediment supply decreased by the Eocene as the modern shelf was established, and much of Eocene deposition consists of neritic shales of the Banquereau Formation

(Grant et al., 1986; McAlpine, 1990). Nonetheless, clastic-rich deposition did occur in the early Eocene, with deltaic or shelfal deposition of a regressive unit continuing from the

Paleocene into the Ypresian in the west and south of the basin (Deptuck et al., 2003). To the east and northeast, deposition of lower Eocene submarine fans on the slope to basin floor occurred much as in the early Paleocene (Deptuck et al., 2003).

Some thousand kilometers to the southwest, the Scotian Basin and Scotian Shelf are characterized by mudstone deposition for much of the Cenozoic (McIver, 1972; Grant et al., 1986). As in the Jeanne d’Arc Basin, the Banquereau Formation here comprises the

Paleocene through Pliocene and is made up mostly of mudstone (McIver, 1972; Grant et al., 1986).

It should be noted that understanding of western North Atlantic deep-water Eocene sedimentation (off the coast of the eastern United States and much of Canada) is generally limited by a lack of well/core penetrations of sufficient depth to sample older sediments.

This is largely due to a dearth of deep-water hydrocarbon exploration, which in other areas

(e.g., the Gulf of Mexico and the western South Atlantic) contributes to understanding of

Eocene and older strata.

Sedimentation in the western North Atlantic (off the coast of the United States) has often been influenced by tectonism in onshore source areas, namely, the Appalachian Mountains 355

(Poag and Sevon, 1989). The Appalachians formed during a protracted period of plate convergence and orogenic activity bounded by the Devonian Acadian orogeny (Bradley,

1983; McKerrow et al., 2000) and the Carboniferous to Permian Alleghanian orogeny

(Horton et al., 1989; Hatcher, 2002; Nance et al., 2010), the latter of which marked the collision of Laurussia and Gondwana and the resultant formation of Pangaea (Blakey and

Wong, 2003). Breakup of Pangaea in the earliest Jurassic led to the opening of the Atlantic

Ocean Basin, and formation of the western North Atlantic coastline (Dietz and Holden,

1970; Hames et al., 2000; Bartolini and Larson, 2001).

Here, a combination of uplift, intense weathering, and erosion has since resulted in three primary periods of sedimentation, namely, the Early-Middle Jurassic, the mid-Early

Cretaceous, and the Late Cenozoic (middle Miocene) (Poag and Sevon, 1989). However, overall rates of denudation in the Appalachians have significantly decreased from high

Paleozoic rates to low Mesozoic and Cenozoic rates concurrent with a cessation of major orogenic activity (Matmon et al., 2003). Nonetheless, Cenozoic shelf progradation associated with the western North Atlantic margin north of Cape Hatteras is inferred to have contributed to several kilometers of sediment thickness comprising an offshore prism of turbidites, slump deposits, and hemipelagic clays (Schlee et al., 1976). However, DSDP drilling data, including a deep penetration on the continental rise off the coast of New York state, suggest that the Cenozoic sediment of the western North Atlantic continental rise, slope, and abyssal plain is predominantly fine-grained, with very little sediment of petroleum reservoir quality (i.e., very little sand) (Mattick et al., 1978), though siliceous turbidites of pre-middle Eocene age are reported (Tucholke and Mountain, 1979).

Furthermore, sediment accumulation rates calculated for this region (primarily the

356 extensive Baltimore Canyon Trough and Hatteras Basin) for the Paleocene-early Miocene are a full order of magnitude lower than either the preceding (Late Cretaceous) or succeeding (middle Miocene to present) time periods, which, paired with high sea levels leading to trapping of sediment in middle-shelf depocenters, suggests a general dearth of deep-water clastic sedimentation (Poag and Sevon, 1989). This drop in sedimentation rates is attributed largely to a period of tectonic quiescence (and perhaps, low topography) in the

Appalachian Mountains from the Late Cretaceous through the early Miocene (Poag and

Sevon, 1989). Interestingly, increased temperature and rainfall of the Paleocene meant development of a rainforest in the eastern United States, which would have contributed to trapping of sediment in onshore environments (Poag and Sevon, 1989). Heavy rainfall and elevated temperatures likely meant chemical weathering dominated at this time (Poag and

Sevon, 1989).

South of Cape Hatteras, Cenozoic sediments off the east coast of Florida are inferred to be some kilometers thick, and of a predominantly terrigenous origin (Shipley et al., 1978).

Lower Cenozoic sediments occur as fine-grained deposits seaward of a prograding shelf margin (downlapping clinoforms) evident in seismic, though sufficient age resolution to distinguish the lower Eocene is lacking (Shipley et al., 1978). Age resolution is also lacking in cored locations, with penetration of upper Eocene sediments possible and presence of

Oligocene and Paleocene sediments only inferred (Locker and Laine, 1992).

The breakup of Pangaea resulted in the Late Triassic to Late Cretaceous (early Campanian) divergence of the North American and South American continents (Dietz and Holden,

1970; Pindell et al., 1988). During this time, the Gulf of Mexico opened (Pindell et al.,

1988). Starting in the Late Cretaceous to Paleocene with the shift from a carbonate to a

357 siliciclastic system (Mackey et al., 2012), the Gulf of Mexico has represented the sink for most sediment sourced from southern North America, from the Western Cordillera to the

Appalachian Mountains (Galloway, 2008; Blum and Pecha, 2014). During the Paleocene to middle Eocene, uplift in the Central and Southern Rocky Mountains and the Sierra

Madre Oriental due to the Laramide orogeny was responsible for the majority of lower

Cenozoic sediment sourcing and delivery to the Gulf of Mexico (Winker, 1982; Galloway et al., 2000, 2005). As today, the early Cenozoic northern Gulf of Mexico Basin received sediment primarily from major extrabasinal fluvial systems encompassing drainages that cover much of southern North America (Galloway, 1981; Winker, 1982; Galloway et al.,

2005; Blum and Pecha, 2014).

A major increase in sediment delivery to the Gulf of Mexico occurred during deposition of the lower Wilcox, likely due to drainage capture in the Sevier-Laramide structural province

(Sharman et al., 2017). However, rates of deposition and volume of sediment delivery in the Gulf of Mexico subsequently decreased significantly 53-51.8 Ma as a result of drainage closure and formation of lakes onshore (Galloway et al., 2011; Sharman et al., 2017; Zhang et al., 2018). Nonetheless, intensified tectonism and uplift onshore continued to contribute to appreciable volumes of sediment delivery, if at much lower rates, to the Gulf of Mexico

(Galloway et al., 2000, 2005, 2011). Simultaneously, elevated levels of CO2 during the early Eocene likely led to elevated rates of silica dissolution, and thus, higher rates of physical weathering and denudation in the Rocky Mountains (Smith et al., 2008). This perhaps contributed, in conjunction with major Laramide uplift (Winker, 1982) and continued activity of the Rio Grande and Colorado rivers (Galloway et al., 2011), to continued delivery of appreciable volumes of sediment to the Gulf of Mexico during the

358 early Eocene and deposition of the lower Eocene upper Wilcox (Galloway et al., 2011).

Furthermore, despite a decrease in deposition rates, the depositional framework for the lower Eocene upper Wilcox remained quite similar to that of upper Paleocene lower

Wilcox (Galloway et al., 2000). The western Gulf of Mexico was blanketed by a basin floor apron rich in deep-water turbidite sands (e.g., Zarra, 2007; Kane and Pontén, 2012) extending hundreds of kilometers offshore and similar to the lower Wilcox apron of the

Paleocene, though extending somewhat more eastward (Galloway et al., 2000; Meyer et al., 2005, 2007; Galloway, 2008). The eastern Gulf of Mexico remained sediment-starved, save for muddy turbidites of the Chicxulub plume near the Yucatan scarp (Galloway et al.,

2000; Galloway, 2008). The western apron consists of aprons sourced from major delta complex feeders, shelf feeders, canyon-sourced fans, and delta bypass deposits of significant volume, the latter of which are common to all major Paleogene and Neogene depositional episodes (Winker, 1982; Mancini and Tew, 1995; Galloway et al., 2000).

Major episodes of Cenozoic delta-fed sand-rich deposition are inferred to be controlled by sediment supply rather than sea level fluctuation (Winker, 1982). A decrease in sediment supply due to erosion and burial of the southern Rocky Mountains occurs in the middle to late Eocene, and a regional transgressive blanket of shelf mud (the Reklaw Shale) caps deposits of the upper Wilcox at about 49 Ma (Galloway et al., 2000; Galloway, 2008). This occurs in concert with a middle to late Eocene to Oligocene transition back into icehouse climates, which some authors argue is responsible for the observed decrease in delivery of sand-rich sediment to the Gulf of Mexico Basin floor due to a switch from more static

Paleocene-early Eocene river mouths (which, as mentioned above, are largely responsible for voluminous delivery of sand deposited as the Wilcox Formation) to highly transitory

359 river mouths (which leads to trapping of sand in shelfal environments) (Sweet and Blum,

2011).

360

Appendix C-2.7.2. Central America and The Caribbean

During the divergence of the North American and South American continents and resultant opening of the Gulf of Mexico, opening of a proto-Caribbean seaway also occurred (Pindell et al., 1988). Cretaceous convergence along what are now the Caribbean Plate’s northern and southern margins initiated eastward motion of the Caribbean Plate (Malfait and

Dinkelman, 1972; Burke et al., 1984). This Cenozoic movement gave rise in the Eocene to some 1000 km of eastward motion of the plate relative to North and South America

(Malfait and Dinkelman, 1972; Burke et al., 1984; Meschede and Frisch, 1998; Pindell and

Kennan, 2009). The Caribbean Plate continues relative eastward motion at present.

Just to the north of the Caribbean Plate sits Cuba, part of the Greater Antilles Arc that began to form in the Cretaceous following Pangaea breakup (Pindell and Kennan, 2009).

This arc originally formed on the northern edge of the Caribbean Plate as the North

American Plate subducted beneath it, but plate reorganizations and collision of the Cuban arc with the Bahamian platform and Florida in the middle to late Eocene (but perhaps starting as early as the Paleocene; Pindell et al., 1988; Bralower and Iturralde-Vinent, 1997) transferred a portion of the northwestern Caribbean Plate (including Cuba) to the North

American Plate, where Cuba now resides (Pindell and Kennan, 2009). Strike-slip motion along the northern boundary of the Caribbean plate contributed to basin development in western Cuba, particularly with regard to formation of the Los Palacios Basin (Brust et al.,

2011; Villegas-Martín et al., 2014). Furthermore, tectonic activity in the early Eocene influenced deposition within the Los Palacios Basin, with emergence of elevated areas to the northeast of the basin contributing siliciclastics to the basin fill (Brust et al., 2011;

Villegas-Martín et al., 2014). These lower Eocene clastics, described as the Capdevila

361

Formation, consist of mudstone and sandstone intercalated with carbonates that grade upward into sandstones, conglomerates, and some mudstones, and are interpreted to have been deposited by turbiditic currents in a bathyal marine environment (Brust et al., 2011;

Villegas-Martín et al., 2014). In the later early Eocene, rapid subsidence and tectonic stability led to deposition of hemipelagic marls of the Universidad Formation (Brust et al.,

2011). One study reports the occurrence of upper Paleocene to lower Eocene turbidites in the Vertientes Formation in central Cuba (Menéndez et al., 2011). This study reports slope- associated clastic and bioclastic sediments likely deposited in a bathyal environment, and postulates sedimentation was influenced by intensity of precipitation (Menéndez et al.,

2011). Deposits in north-central Cuba record deposition in a Paleogene foredeep basin, however, deposition in the early Eocene seems to consist mostly of limestones and calcareous breccias (Paso Abierto Formation), carbonate breccias (Embarcadero

Formation), and marls, marly limestones, and some mudstone and sandstone (Vigía

Formation), and limestone (Vega Formation) (Iturralde-Vinent et al., 2008).

Active margin processes similarly influenced deposition of lower Eocene sediments now exposed on the island of Hispaniola, also part of the Greater Antilles Arc (Dolan et al.,

1991). The Peralta Belt, which cuts across the island diagonally from northwest to southeast, is in the southeast composed of two basinal sequences, one of which is the lower to upper Eocene Peralta Group (Dolan et al., 1991). The lower Eocene Ventura Formation makes up the oldest basin fill of the Peralta Basin, and is composed almost exclusively of turbiditic sandstones, which are largely composed of volcaniclastics, and mudstones

(Dolan et al., 1991). This formation is overlain by limestones of the middle Eocene Jura

Formation (Dolan et al., 1991). Dolan et al. (1991) proposed that the lower Eocene

362 turbidites were sourced erosionally from the Greater Antilles Arc and deposited in a narrow trough oriented parallel to the arc. These authors proposed siliciclastic sediment was sourced during relative sea level lowstand, while overlying carbonate material was sourced during sea level highstand (Dolan, 1989; Dolan et al., 1991). Influence of an early Eocene highstand is not discussed by Dolan et al. (1991).

Along the southern and eastern boundaries of the Caribbean Plate, east of the Caribbean

Sea and at the western extent of the Atlantic, accretionary processes influence depositional processes of Barbados, and served to incorporate and expose subaerially deep sea Atlantic sediments (Pudsey and Reading, 1982; Speed and Larue, 1982). Lower Eocene to upper

Oligocene pelagic and hemipelagic sediment, with little influence of continental sediment, is juxtaposed, most likely due to tectonics, against a suite of poorly dated (but likely

Cretaceous to Eocene) clastic sediment derived from continental South America and accumulated as coarse turbidites and mudstone in a trench wedge or in subsea fans (Speed,

1981; Speed and Larue, 1982). Lower Eocene outcrop on eastern Barbados is comprised both of hemipelagic terrigenous turbidites (quartz sandstone, conglomerates, mudstone) of the lower Scotland Formation and likely deposited in a trench wedge or in submarine fans in deep water (Speed, 1981; Pudsey and Reading, 1982) and of hemipelagic sediments

(radiolarian earth, marl, quartz sandstone, mudstone) potentially representing some degree of turbidite activity in slope basins (Speed, 1981). Early Eocene (through late Oligocene) deposition evinced by outcrop also includes pelagic sediments from a deep oceanic environment (Speed, 1981). Deposition of the turbiditic Scotland Formation lasted into the late Eocene (Speed, 1981). Northeast of Barbados, just east of the overriding Caribbean

Plate, and sitting atop the westward-subducting South American Plate, evidence of roughly

363 equivalent terrigenous turbidites exists on the Tiburon Rise, however, this clastic sequence appears to be at oldest middle Eocene to late Oligocene in age (Dolan et al., 1989, 1990).

Importantly, this deposition, and perhaps the lower Eocene turbidite deposition of

Barbados (which was also perhaps originally sediment deposited on the Atlantic Ocean floor/South American Plate, albeit later scraped off and incorporated into Barbados;

Pudsey and Reading, 1982), was likely generated due to tectonic forces active on the northern margin of South America, rather than due to eustasy (Dolan et al., 1989, 1990).

In the Caribbean Sea to the south of Cuba and Hispaniola and to the west of Barbados, on the fringes of the Colombia Basin and throughout the Venezuela Basin, legs 4 and 15 of the Deep Sea Drilling Project drilled some nine sites that penetrated lower Eocene strata

(Bader et al., 1970; Edgar et al., 1973). Throughout these sites, and throughout the

Caribbean Sea itself (Ewing et al., 1967), lower Eocene sediments are characterized by quite uniform pelagic sedimentation, with an assortment of cherts, limestones, cherty carbonate, siliceous limestones, radiolaria-rich layers, and chalk encountered at numerous sites (Bader et al., 1970; Edgar et al., 1973). To the west within central Colombia Basin, chalk, chert, and siliceous clay similarly characterize the lower Eocene, while further west within the western Colombia Basin, the lower Eocene is characterized by hiatus, erosion, or condensed section (Bowland, 1993).

The Belize Basin of southern Belize and eastern Guatemala records latest Cretaceous through Eocene siliciclastic deposition (Ramanathan and Garcia, 1991; Schafhauser et al.,

2003). This interval, described as the Sepur/Toledo Formation, is characterized by sequences of shale and sandstone, as well as breccias, inferred to have been deposited by turbidity currents and debris flows of proximal deep-water submarine fans initiated in

364 response to tectonic activity in the adjacent Maya Mountains (Ramanathan and Garcia,

1991; Schafhauser et al., 2003).

In the northeastern portion of Honduras and the northern part of Nicaragua, the Mosquitia

Basin comprises a significant amount of coastal outcrop (Mills and Hugh, 1974; Mills and

Barton, 1996). In the coastal region, pull-apart basins formed here during the Cenozoic

(Mills and Barton, 1996). These basins were filled largely with sediment from the Patuca and Coco rivers, with Paleogene sediments composed predominantly of the Ahuas red beds

(Mills and Barton, 1996). This generally fits with the character of sediments throughout the Mosquitia, where red beds make up much of the Late Cretaceous through Pliocene in

Honduras and Late Cretaceous through Eocene in Nicaragua (Mills and Hugh, 1974).

In the late Cretaceous to Cenozoic Sandino forearc basin, which formed due to subduction of the Cocos and Nazca plates beneath the Caribbean Plate and lies at present day coastal and offshore western (Pacific) Nicaragua (Kumpulainen, 1995; Struss et al., 2008), the

Brito Formation of Eocene age (it should be noted that Struss et al. (2008) cite biostratigraphic literature assigning the outcropping Brito a middle to late Eocene age and well data assigning the Brito a latest Paleocene to early Oligocene age, while Ranero et al.

(2000) describe the onshore Brito as upper Paleocene to lower middle Eocene) crops out onshore and is made up of volcanic breccias, tuffs, limestones, and shales (Swain, 1966).

This sits atop Late Cretaceous turbiditic tuffaceous shale and sandstone of the Rivas

Formation, and is overlain by Oligocene massive tuffs, breccias, and shales of the

Masachapa Formation (Swain, 1966). Kumpulainen (1995) suggests the Rivas and Brito were deposited as basin plain and deep water fan deposits, Struss et al. (2008) describe the

Brito as submarine fan deposits, and Ranero et al. (2000) describes the Brito from onshore

365 outcrop and drillholes as volcaniclastic turbidite deposits. Onshore work in southwestern

Nicaragua and northwestern Costa Rica described the Brito as stacked channel-levee deposits of small, overlapping submarine fans (Struss et al., 2007). According to

Kumpulainen (1995), paleobathymetry data suggests abyssal depths for the Brito and

Masachapa to shallower depths for younger formations, while other authors suggest bathyal, though similarly deep-water, depths for the Brito and a similar overall shallowing- upward trend (Ranero et al., 2000). The Brito Formation is suggested to extend westward into the offshore portion of the Sandino forearc basin (Ranero et al., 2000).

In eastern Costa Rica, the Limón Basin spans the coastal plain and continental shelf

(Brandes et al., 2007). This basin sits atop the Caribbean Plate, and is a complex system composed of a northern back-arc and southern foreland basin region (Brandes et al., 2007).

The basin formed in the Late Cretaceous due to subduction of the Farallon Plate below the

Caribbean Plate, has likely been influenced by deformational events since the Paleocene, and is filled with Late Cretaceous to recent sediments (Brandes et al., 2007). This sediment fill records Paleocene to early Eocene deposition of the Tuís Formation, which is composed of coarse-grained volcaniclastic turbidite deposits, debris-flow deposits, tuffs, and lava flows (Bowland, 1993; Brandes et al., 2007). Authors propose the Tuís Formation formed as a prograding apron system (Brandes et al., 2007). These sediments are overlain by either shallow-marine carbonates of the Las Animas Formation or hemipelagic mudstones, turbidite deposits, and debris-flow deposits of the Senosri Formation, both of middle

Eocene to Oligocene age (Brandes et al., 2007).

Outcrop from eastern Panama reveals a succession of lower to middle Eocene abyssal volcanics mixed with radiolarian-rich sediments (Bandy and Casey, 1973; Bowland, 1993).

366

These deep-water deposits occur as tuffaceous sandstones, shales, and limestones collectively called the Morti Tuffs (Bandy and Casey, 1973; Bowland, 1993).

367

Appendix C-2.8. South America

Figure C-2.8. Locations of Paleogene sedimentary deposits described for the South

American continental margin (early Eocene paleomap after Scotese, 2001). MFB—

Maracaibo foreland basin, EVB—Eastern Venezuela Basin, GSB—Guyana-Suriname

Basin, MB—Mundaú Basin, SeB—Sergipe Basin, ESB— Espírito Santo Basin, CB—

Campos Basin, SB—Santos Basin, PEP—Punta del Este and Pelotas basins, CAB— offshore central Argentina basins, MaB—Malvinas Basin, AB— Austral Basin, PC—Peru-

Chile Trench, TB—Talara Basin, EF—Ecuador forearc regions, PCB—Pacific Coastal

Basin, MM—Middle Magdalena Valley and Eastern Cordillera. Symbols as in Figure C-

2.1.

368

Table C-2.8A. Locations of lower Eocene turbidites associated with the South American continent, as well as descriptions of deposition during the neighboring time periods.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin Venezuela (Maracaibo deltaic/shallow same as early foreland basin) ? mar Eocene ? Yes Yes glauconitic shale with thin Eastern shale/thin mix of early Venezuela deepwater sandstone and late glauconitic Basin shale interbedded Eocene limestone ? Yes

offshore Guyana- same as early same as early Suriname Basin ? submarine fan Eocene? Eocene? ? No Brazil (Sergipe same as same as same as Basin) turbidites Paleocene Paleocene Paleocene Yes No drastic Brazil (Mundaú same as reduction in same as Basin) turbidites Paleocene sediment middle Eocene Yes No Brazil (Espírito coarse regressive same as Santo Basin) ? turbidites megasequence middle Eocene No No* Brazil (Campos Basin) clastics turbidites clastics clastics No No* Brazil (N & central Santos clastics/ same as same as same as Basin) turbidites Paleocene Paleocene Paleocene Yes No* Uruguay (Punta del regressive Este & Pelotas transgressive fluvial-deltaic same as early same as early basins) sequence sediment Eocene Eocene Yes? No regressive regressive/shall same as Austral Basin turbidites ow water transgressive middle Eocene No? Yes regressive regressive/shall same as Malvinas Basin turbidites ow water transgressive middle Eocene No? Yes

*indicates passive margin, but local tectonics

369

Table C-2.8B. South American locations surveyed but lacking lower Eocene turbidites, including descriptions of deposition during the Paleocene through late Eocene.

Location Paleocene early Eocene middle Eocene late Eocene River Active margin onshore Guyana- estuarine same as same as same as Suriname Basin sandstone Paleocene Paleocene Paleocene Yes? (delta) No offshore central hemipelagic, Argentina low-energy same as same as same as basins deposition Paleocene Paleocene Paleocene No No Peru-Chile Trench ? ? ? ? No (no major) Yes Peru continental margin (10 no early or Mesozoic/Ceno middle Eocene zoic basins) sediment Yes Peru (Talara shallow same as Basin) marine/fluvial Paleocene deepwater ? ? Yes Ecuador marine forearc regions turbidites hiatus sediment ? ? Yes N Ecuador-S Panama unconformity/ (Pacific same as resumed Coastal Basin) unconformity Paleocene sedimentation ? Yes Middle Magdalena Valley, Eastern Cordillera, & E alluvial/coastal same as Llanos Basin plain/estuarine Paleocene ? ? ? Yes

370

Following the rifting of the South American and African continents and resultant opening of the Atlantic Ocean, the northern margin of South America existed as a passive margin.

This passive margin prevailed from the Late Jurassic/earliest Cretaceous until the onset of active tectonism in the Eocene (Erlich and Barrett, 1992). Strike-slip motion and compression initiated earliest (early-middle Eocene) in western Venezuela and northern

Colombia, and later (late Oligocene-middle Miocene) in eastern Venezuela and Trinidad

(Erlich and Barrett, 1992). Eastward movement of the Caribbean Plate along the northern margin of the South American Plate continues today.

The Maracaibo foreland basin of Venezuela records Eocene deposition dominated by the influence of sediment input from the proto-Maracaibo River and great delta system

(Talukdar et al., 1986; Xie et al., 2010). This fluviodeltaic influence paired with Paleocene-

Eocene foreland basin subsidence controlled stratigraphic evolution of the basin until the middle Eocene (Lugo and Mann, 1995; Escalona and Mann, 2006). Eustasy played little role during this period (Escalona and Mann, 2006). In the central Maracaibo Basin, the

Eocene is dominated by the Misoa Formation, which exists as a large (~180 km wide) clastic wedge sourced from the northward-flowing proto-Maracaibo River and major delta system (Talukdar et al., 1986; Xie et al., 2010). The lower to middle Eocene in the central

Maracaibo is characterized by a succession of aggradational fluviodeltaic sandstones overlain by retrogradational shallow-marine shales and sandstones (Escalona and Mann,

2006).

In the Eastern Venezuela Basin, a general trend of passive margin, shallow marine clastic deposition prevailed from the end of the Cretaceous through the Paleogene (Erlich and

Barrett, 1992). Deposition accompanied the onset of tectonism and formation of the

371 foreland basin, resulting in a succession of some 3 km of Eocene-Oligocene clastics within the basin (Erlich and Barrett, 1992). Much of the Eocene and younger sediment onshore is characterized by sandstones and shales, likely with shales making up most of the thickness

(Hedberg, 1950). In the southern part of the basin, Eocene-Oligocene uplift led to development of an erosional unconformity (Erlich and Barrett, 1992). A study found that the upper Paleocene-Eocene is absent onshore due to such an erosional unconformity, but that the upper Paleocene-Eocene is present as a thin unit of near constant thickness widespread on seismic lines (Di Croce, 1996). The upper Thanetian in offshore wells is characterized by shales representing deep (bathyal) water depths (Di Croce, 1996). This is overlain by lower Eocene glauconitic shale with interbedded thin beds of silty shale and fine-grained sandstone, which grades into upper Eocene glauconitic limestone (Di Croce,

1996).

At present, the eastern margin of South America lies in a passive margin setting. The breakup of Gondwana and resultant rifting of the South American and African continents in the Cretaceous created a rifted margin (and numerous rift basins) along South America’s eastern coast.

The Guyana-Suriname Basin spans coastal to ultradeep offshore French Guiana, Suriname,

Guyana, and eastern Venezuela. In the coastal area of the basin, a transgressive sequence gave way to a regressive sequence toward the end of the Paleocene, with shallowing water depths and ultimately an erosional unconformity characterizing the Paleocene-Eocene boundary (Wong, 1994). This unconformity is overlain by presumably Eocene sediments, which seem to record a period of transgression (Wong, 1994). Onshore, unconsolidated

Paleocene and Eocene sands provide significant petroleum reservoirs. Eocene sands here

372 were deposited in an estuarine delta setting (Toelsie and Goerdajal, 2013). The Paleocene reservoirs below were deposited in a fluvio-estuarine to coastal marine setting (Toelsie and

Goerdajal, 2013). Little granular data exist for offshore Guyana-Suriname, although a publication (Nemčok et al., 2016) and various reports discuss the presence of a deep-water

Eocene turbidite fan (referred to as “Eagle”) imageable in seismic. This fan sits stratigraphically above a basin-wide unconformity of Thanetian-Ypresian age (Nemčok et al., 2016).

Along the margin of eastern Brazil, deposition following Cretaceous rifting was dominated by continental sediments and volcanics in the Early Cretaceous, evaporites in the Aptian, carbonate deposition in the Albian to late Albian, and low-energy sediments in the late

Albian (Demercian et al., 1993; Karner and Driscoll, 1999). Overall, the Cenozoic

Brazilian continental margin is characterized by a regressive marine supersequence, with carbonate and siliciclastic deposition extending from the platform to deep-water (Mohriak,

2003), and deposition along eastern/southeastern Brazil (into the Espírito Santo, Campos, and Santos basins) driven by clastic sediments derived from the coastal mountain range of

Serra do Mar (Demercian et al., 1993). Major rivers have played a role in clastic delivery along the margin, as in the Mundaú and Santos basins (Karner and Driscoll, 1999; Modica and Brush, 2004). Overall, these forms of clastic sedimentation mean Eocene turbidites are found in several Brazilian basins, including the Campos, Santos, and Sergipe, and that

Paleocene-Eocene sedimentation rates along southeastern Brazil were not markedly lower than preceding or succeeding periods (Macgregor, 2012).

In the Mundaú and Sergipe basins of the northeastern and eastern Brazilian margins, fluvial systems exerted a marked control on the margins until the mid-Eocene, with corresponding

373 formation of canyon systems and turbidite deposition during periods of sea level fall

(Karner and Driscoll, 1999). The Sergipe Basin continues to be influenced by these systems, however, the turbidite system was drastically reduced in the mid-Eocene Mundaú

Basin, likely due to drainage capture of its feeder river, the Parnaíba (Karner and Driscoll,

1999).

In the Espírito Santo Basin, deposition along the eastern Brazilian margin in the late Albian to early Eocene is characterized by a transgressive succession (Bruhn and Walker, 1997).

Lower Eocene coarse-grained (sandy to conglomeratic) turbidites fill fault-controlled canyons in successions up to 400 m thick (Bruhn and Walker, 1997). These turbidites form a deepening-upward sequence of progressively narrower, thinner, and finer-grained deposits reflecting a period of sea level rise in the early Eocene (Bruhn and Walker, 1997).

Following this period of sea level rise, and starting in the early to middle Eocene, the sedimentary succession is characterized by a marine regressive megasequence (Bruhn and

Walker, 1997), consistent with overall Cenozoic regression off the Brazilian margin

(Mohriak, 2003).

Sedimentation in the post-Paleocene Campos Basin of Brazil is generally characterized by a regressive marine megasequence (Guardado et al., 1989). Thick packages of clastic sediments were deposited, with turbidite deposition controlled by salt movement and sea level fluctuations (Guardado et al., 1989). This deposition included turbidites of early

Eocene age (Mohriak et al., 1990). However, for the early and the middle Eocene, disconformities identified from continental shelf well data suggest periods of relative sea level rise (Contreras et al., 2010). This sea level rise led to landward migration of depocenters and basinward sediment starvation (Mohriak et al., 1990; Contreras et al.,

374

2010). On the slope, sandstones and shales exist, likely as products of gravity-driven processes (Contreras et al., 2010).

From the Late Cretaceous through the Oligocene, the ancestral Paraíba do Sul drainage characterized sedimentation in the northern and central Santos Basin off of Brazil (Modica and Brush, 2004). High clastic influx from this drainage system, driven by uplift and denudation of the Serra do Mar range in conjunction with high rainfall at the Eocene

Thermal Maximum, resulted in Eocene shelf progradation and turbidite deposition despite the global sea level high stand (that is, sedimentation rates outpaced accommodation)

(Modica and Brush, 2004; Berton and Vesely, 2016). In the northern Santos Basin, this pattern of progradation lasted from before the Paleocene through all of the Eocene, ending at the Eocene-Oligocene boundary due to a shift of the Paraíba do Sul to the north and into the Campos Basin (Berton and Vesely, 2016). In the southern Santos Basin, however, clastic input was effectively starved (Modica and Brush, 2004). One seismic line in this southern portion of the Santos Basin shows fluvio-deltaic sediments on the outer shelf- upper slope that grade into shales, marls, and sparse turbiditic sandstones on the slope

(Contreras et al., 2010).

Passive margin basins off the coast of Uruguay record a eustasy-driven regressive and transgressive sequence in the Paleocene and a regressive sequence in the Eocene-Oligocene

(Soto et al., 2011). This Eocene-Oligocene regressive sequence is made up of fluvio-deltaic sediments in the Punta del Este and Pelotas basins (Contreras et al., 2010; Soto et al.,

2011). In the Pelotas Basin, the upper Paleocene to lower Eocene is underlain by a shelf- wide erosional unconformity, above which sub-parallel reflectors on the upper slope dip toward the basin, but are absent on the shelf (Contreras et al., 2010). The character of

375 sediments in the ultra-deep-water Oriental del Plata Basin is unknown (Soto et al., 2011).

Sedimentation in the Cretaceous-Paleocene recorded a shift from turbidite-dominated deposition to bottom current-dominated (contourite) deposition (Creaser et al., 2017).

A number of underexplored basins lie off the eastern coast of Argentina. Off the coast of central Argentina (~41°S to 45°S), deposition during the Paleocene through Eocene was mostly hemipelagic and low-energy with little bottom water activity (Gruetzner et al.,

2012). This is the case for the eastern Colorado Basin, where Paleocene to early Eocene deposition is inferred to have taken place in a low-energy, deep-water setting (Bushnell et al., 2000). For this same period, however, the Argentine shelf is characterized by an unconformity between the Paleocene Pedro Luro Formation and the Eocene Elvira

Formation, and the Paleocene to early Eocene here is generally taken to be a time of transgression (Bushnell et al., 2000; Franke et al., 2006). In the onshore portion of the San

Jorge Basin, little Cenozoic data is available, however early Eocene to late Oligocene deposition seems to be nonmarine (Fitzgerald et al., 1990).

To the south of the passive margin rift basins of Brazil, Uruguay, and some of Argentina, uplift of the Patagonian and Fuegian Andes led to active margin, foreland basin deposition in the resultant Austral (Magallanes) and Malvinas basins (Olivero and Malumián, 2008).

This active margin setting resulted in quite different styles of sedimentation in comparison to the passive margin to the north, namely, much of the Cenozoic in the northern passive margin setting is characterized by an overall regressive succession, whereas in the southern active margin setting, most of the Cenozoic until the end of the Miocene is characterized by an overall transgressive succession.

376

In the Austral and Malvinas basins (throughout Tierra del Fuego), almost all sedimentation occurred in deep-water settings (Torres Carbonell and Olivero, 2012). Overall, the uppermost Cretaceous through Eocene Malvinas Basin is characterized by a sediment- starved interval (Galeazzi, 1998). However, in the Austral and Malvinas basins, the upper

Paleocene-lower Eocene Río Claro Group records an overall regressive megasequence, with deep-water turbiditic deposition in the Paleocene transitioning into shallower, shelfal deposition in the early Eocene (Olivero and Malumián, 2008). This shelfal deposition was relatively short-lived, with shallow-water conditions only occurring in the Lutetian

(Olivero and Malumián, 2008; Torres Carbonell and Olivero, 2012). Starting in the late middle Eocene and lasting through the late Eocene, deposition of the La Despedida Group represented an extended transgression (Olivero and Malumián, 2008), and deep-water settings persisted through the Miocene (Olivero and Malumián, 2008; Torres Carbonell and Olivero, 2012). Subsurface data from the distal foredeep of the Austral Basin show an

Eocene succession that ranges from 100 m thick in the north to >1,000 m thick in the south

(Torres Carbonell and Olivero, 2012). This succession is mostly a mud-rich clastic wedge, with sparsely interbedded ~tens of meters thick glauconitic sandstones (Torres Carbonell and Olivero, 2012). In the distal foredeep of the Malvinas Basin, the Eocene is composed of sandstones, mudstones, and carbonates, with thicknesses in excess of 500 m to the south

(Torres Carbonell and Olivero, 2012). The Eocene succession between the Austral and

Malvinas basins is composed of greater than 500 m of mudstones with interbedded sandstones (Torres Carbonell and Olivero, 2012). In outcrop on the Atlantic coast of Tierra del Fuego, the Eocene succession is much sandier than seen in subsurface (Torres

Carbonell and Olivero, 2012). The Ypresian-early Lutesian interval is characterized by

377 proximal foredeep deposition consisting of deep-water turbidites and shelfal deposits

(Torres Carbonell and Olivero, 2012). The turbidite successions occur in thicknesses in excess of 200-450 m (Torres Carbonell and Olivero, 2012).

As in the Austral and Malvinas basins to the east (Atlantic) of the southernmost South

American continent, so too is the Cenozoic history of the western (Pacific) margin of South

America influenced by active tectonism and Andean uplift. Through most of the Mesozoic, extensional and neutral tectonic regimes characterized the margin (Horton, 2018), but in the Late Cretaceous, a transition to compressional tectonics occurred along the margin, giving rise to the Andean orogenic belt (Hoorn et al., 2010; Ramos, 2010; Horton, 2018).

Rates of uplift were initially relatively slow, with much Andean uplift and acceleration of uplift rates occurring since the late middle Miocene (Gregory-Wodzicki, 2000; Hoorn et al., 2010). The Andean belt at present spans the entirety of the South American continent, from Argentina and Chile in the south to Colombia and Venezuela in the north. The mountain belt is a continental divide with large drainages and long rivers to the east, and small drainages and short rivers to the west (Horton, 2018).

The Peru-Chile Trench along the western margin of South America provides a sink for sediment delivered from the small western drainages and rivers of Chile and Peru (Horton,

2018), and possibly has been such a sink since initiation of Andean orogenesis.

Sedimentation west of the trench is pelagic with sediment composed of fine muds, much of which is calcareous ooze (Zen, 1959; Scholl et al., 1968). This pelagic sediment overlies basalt of the Nazca Plate, and is some 130-150 m thick west of the trench (Scholl et al.,

1968). The trench itself is filled with some amount of turbidite deposits (up to hundreds of meters), although previous workers hypothesize the bulk of this sediment must be Miocene

378 or younger (Scholl et al., 1968, 1970), and that much of this sediment was sourced during glaciated periods of low sea level (Scholl et al., 1970). East and landward of the trench, sediment is sandier still, though perhaps this too is mostly younger sediment reflecting the

Miocene-Pleistocene increase in Andean uplift rates (Zen, 1959; Scholl et al., 1968). At present, Peru-Chile Trench submarine fan formation and turbidite deposition along the western South American margin occurs at point sources, namely submarine canyons and gullies (Thornburg and Kulm, 1987), providing a mechanism for deposition along the active margin in the past (Prince et al., 1974) (however, it should be noted that deposition may have been more limited in the past, given lower rates of uplift, and thus, lower erosion onshore). There is little evidence for early Eocene turbidite deposition off the western coast of Chile.

Inland of the Peru-Chile Trench, the continental margin of Peru contains ten Mesozoic and

Cenozoic forearc basins (Dunbar et al., 1990). However, of these, only the Talara Basin is known to contain lower and middle Eocene sediments (Travis et al., 1976; Dunbar et al.,

1990). Eocene deposits within the Talara contain deltaic, fluvial, and turbidite sands

(Travis et al., 1976). However, deep-water deposition in the basin is restricted to the late middle Eocene (Fildani et al., 2008). Early Eocene deposition in the Talara is marginal marine to fluvial (Marsaglia and Carozzi, 1990; Fildani et al., 2008). In fact, deposition for the duration of the Paleocene through to late middle Eocene is comprised of shallow- marine mudstone deposition with interspersed periods of fluvio-deltaic to shallow-marine coarse siliciclastic sediment deposition (Fildani et al., 2008).

In southern Ecuador, the early Eocene is marked by a sedimentary hiatus likely reflecting regional tectonically-driven emergence, with no recorded formations of early Eocene age

379

(Jaillard et al., 1995). By contrast, the late Paleocene was characterized by coarse turbiditic deposition following a period of deformation and formation of an early fore-arc or slope basin, while the late Ypresian through early Lutetian is manifested as a period of subsidence, transgression, and marine deposition (Jaillard et al., 1995). To the north and in central Ecuador, as in the south, much of the late Cretaceous through Eocene geologic history and resultant sedimentation is dominated by the accretion of numerous oceanic terranes, with little to no early Eocene deep-water sedimentation recorded by the rocks of these regions (Alava and Jaillard, 2005; Marcaillou and Collot, 2008). Throughout the forearc regions of Ecuador, an early Eocene sedimentary hiatus is evident, with marine sedimentation not beginning until the middle Eocene (Nygren, 1950; Jaillard et al., 2000).

Off the coast of northern Ecuador and southern Colombia, between the trench and coastline, a similar story plays out in the southern portion of the Pacific Frontal Basin

(Bueno Salazar, 1989; Marcaillou and Collot, 2008) or Pacific Shelf Basin (Bueno Salazar and Govea, 1974). Here too, accretion of an oceanic terrane to the South American margin prior to the middle Eocene is suspected to be responsible for a hiatus, evident as a regional unconformity in seismic (Marcaillou and Collot, 2008). This hiatus is seen in outcrop analogues from the Borbón Basin of Coastal Ecuador and the Tumaco Basin of the onshore southern Pacific Coastal Basin in Western Colombia (Bueno Salazar and Govea, 1974;

Bueno Salazar, 1989; Evans and Whittaker, 1982; Marcaillou and Collot, 2008; Borrero et al., 2012). In fact, the entire Pacific Coastal Basin, which stretches from offshore and onshore northernmost Ecuador to southernmost Panama (i.e., all of coastal Colombia, from the Colombian Trench offshore to the Western Cordillera onshore) and which here includes the Pacific Frontal Basin, is generally characterized by a Paleocene-middle Eocene

380 regional unconformity associated with tectonic movement, with deposition only resuming in the middle Eocene (Bueno Salazar and Govea, 1974; Bueno Salazar, 1989). Throughout the area of the basin, middle Cretaceous to Paleocene rocks are unconformably overlain by middle Eocene and younger rocks (Bueno Salazar and Govea, 1974; Bueno Salazar, 1989).

During the late Maastrichtian to early Eocene, the present-day Middle Magdalena Valley,

Eastern Cordillera, and eastern Llanos Basin lay in a foreland basin setting associated with accretion of the Western Cordillera (Cooper et al., 1995), which bounds the easternmost extent of the aforementioned Pacific Coastal Basin. The foreland basin megasequence deposited during this time consists of coal-rich alluvial plain, coastal plain, and estuarine deposits (Cooper et al., 1995).

381

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APPENDIX D

SUPPLEMENTARY MATERIAL FOR CHAPTER 6

APPENDIX D-1. DETAILED DESCRIPTION OF UNCONFORMITIES

APPENDIX D-1.1. Africa

APPENDIX D-1.2. Antarctica

APPENDIX D-1.3. Asia

APPENDIX D-1.4. Australia/Oceania

APPENDIX D-1.5. Europe

APPENDIX D-1.6. North America

APPENDIX D-1.6.1. Northern America

APPENDIX D-1.6.2. Central America and The Caribbean

APPENDIX D-1.7. South America

APPENDIX D-2. REFERENCES CITED

437

APPENDIX D-1. DETAILED DESCRIPTION OF UNCONFORMITIES

APPENDIX D-1.1. Africa

Since Gondwana breakup, opening of the Atlantic Ocean, and respective rifting apart of

South America and Africa, the west African margin has predominantly existed as a passive margin (Rabinowitz and LaBrecque, 1979; Lavier et al., 2001).

An exception is northern Morocco, where Cenozoic inversion of Triassic to Jurassic extensional systems, and general tectonic activity has influenced development of the northwestern African margin’s geologic features as a result of African-Eurasian plate convergence (e.g., Hafid, 2000). In southern Morocco where passive margin conditions did prevail, the Oligocene section is missing in the onshore Tarfaya-Laayoune-Dakhla Basin

(Leprêtre et al., 2015). On the onshore section, shelf, and upper rise of the Aaiun Basin of

Western Sahara, the Oligocene section is likewise missing (von Rad and Einsele, 1980), and is also absent along the Mauritanian continental margin (Vear, 2005), though an exception may be the Oligocene-recent Nouakchott delta of central Mauritania (Davison,

2005). These hiatuses are attributed to Eocene-Oligocene regression (Arthur et al., 1979), and potentially represent periods of onshore and shelfal erosion, sediment bypass to deep basins, and deepwater siliciclastic deposition. Major Oligocene unconformities are also present along the margins of Senegal, The Gambia, Guinea-Bissau, Liberia, Côte d'Ivoire,

Ghana, Togo, and Benin (Brownfield and Charpentier, 2003, 2006).

Major early Oligocene unconformities similarly observed in offshore Ghana (Séranne,

1999; Brownfield and Charpentier, 2006), offshore Gabon (Brink, 1974; Teisserenc and

438

Villemin, 1989; Rasmussen, 1997; Séranne, 1999), offshore Congo (Séranne, 1999; Da

Costa et al., 2001; Broucke et al., 2004), and offshore Angola (Brice et al., 1984; Séranne,

1999) have been interpreted as a response to the Eocene-Oligocene climatic transition and eustatic lowering of sea level, whereby depositional systems along West Africa underwent a transition from aggrading Aptian to Eocene carbonate systems to a prograding Oligocene to a present day clastic wedge characterized by elevated influx of terrigenous sediment

(Séranne, 1999).

The modern Niger delta became active, depositing continental sands, deltaic and fluvial facies, and some transitional marine facies, as early as the early Eocene (Short and Stäuble,

1967; Reijers, 2011). The delta saw major expansion starting in the late Eocene to early

Oligocene, prograding some 2 km every 2 Myr along various delta lobes down the continental slope and onto the ocean floor (Burke, 1972; Reijers, 2011). It was also during the late Eocene or early Oligocene that the Opuama channel complex developed as an extensive network of deep, turbidite-cut marine channels (Petters, 1984; Udo and

Ekweozor, 1990), which likely served to deliver clastics to the deep ocean.

The Cenozoic sequence in the Orange Basin off southwestern Africa is relatively thin

(Hirsch et al., 2007). A seismic study noted an Eocene-Oligocene unconformity between two phases of progradational siliciclastic deposition (de Vera et al., 2010), perhaps similar in nature to the unconformities characteristic of the northwestern and western African margin, and likewise perhaps suggesting shelfal erosion and sediment bypass to deepwater basins. However, it should be noted that arid climates during the early Oligocene are inferred to have resulted in drastic reductions in denudation rates onshore, and thus, in

439 reduced delivery of sediment to the deepwater realm along the southwest African margin

(Séranne and Anka, 2005).

Like the west African margin, the southeastern to eastern margin of Africa has existed as a passive margin since the Late Cretaceous or earlier breakup of Gondwanaland (Norton and Sclater, 1979).

In the Southern Mozambique Basin, the middle Eocene to Oligocene interval is predominantly characterized by widespread carbonate deposition (Said et al., 2015). To the north, in and outboard of the Zambezi Delta, however, the Eocene-Oligocene boundary is characterized by an unconformity and the Oligocene is characterized by a nearly six-fold increase in sediment flux to the deep marine realm as compared to the sediment-starved

Paleocene to Eocene interval (Walford et al., 2005). Authors note that Oligocene increases in sediment flux to the Zambezi Delta could have been a response to global climate change, but also note that the increase may be more reasonably attributed to rapid local or regional uplift in Southern Africa (Walford et al., 2005; Mahanjane et al., 2014). Similarly, the

Rovuma Basin of northern Mozambique is characterized by a widespread Eocene-

Oligocene unconformity, and progradation of the Rovuma delta complex in the Oligocene to Miocene (Key et al., 2008; Mahanjane et al., 2014). Here too, authors postulate that deltaic progradation is connected to onshore uplift (Mahanjane et al., 2014). Onshore uplift is also cited as the causal mechanism for Paleocene to Oligocene deposition of turbidite systems in the Tanzania Coastal Basin, where a locally erosive, down-cutting surface is noted at the beginning of the Oligocene (Sansom, 2018). Further north, in the Lamu Basin of southeast Kenya, deposition through the Eocene and Oligocene is described as having occurred in fluvial, deltaic and delta-front, and restricted-shelf settings, whereby authors

440 do not note distinct differences between late Eocene and early Oligocene deposition

(Nyagah, 1995; Zongying et al., 2013).

Along the Somalian coast, authors note that lower to middle Oligocene sediments are apparently absent, with upper Oligocene to lower Miocene sediments overlying upper

Eocene marls (Bosellini et al., 1987). Bosellini et al. (1987) tentatively attribute this unconformity to a dramatic fall in sea level at 30 Ma as in Vail et al. (1977), potentially representative of the climatic changes at the Eocene-Oligocene boundary.

In the modern Nile Delta region, the upper Eocene and lower Oligocene is characterized by shale, with deltaic deposition by an ancestral Nile not beginning until the late Oligocene

(Sestini, 1989). In northern Egypt (onshore), no unconformity is evident at the Eocene-

Oligocene boundary, with late Eocene deposition characterized by sandstone and siltstone and early Oligocene deposition characterized by shale and marl (Swezey, 2009).

Across most all of the Saharan region, however, a major unconformity is present at the end of the Eocene/beginning of the Oligocene (Swezey, 2009). This includes Eocene-

Oligocene unconformities recorded in northeastern Libya, coastal central Libya, offshore northwestern Libya, the Chotts Basin of Algeria and Tunisia, some regions in central

Algeria, the Tindouf-Ouarzazate Basin of Morocco, the Iullemmeden Basin of Niger, the southeastern margin of the Taoudenni Basin of Mali, and the Mauritania-Senegal Basin

(Swezey, 2009). Swezey (2009) notes that the regional extent of this unconformity is most consistent with the Eocene-Oligocene boundary eustatic fall in sea level, but also mentions that tectonic activity throughout much of the Sahara also began in the late Eocene in response to African-Eurasian plate convergence.

441

In offshore northwestern Libya, the Eocene-Oligocene boundary unconformity is preceded by middle to upper Eocene siltstone and shale, and is overlain by Oligocene to lower

Miocene limestone, dolomite, and shale (Swezey, 2009).

442

APPENDIX D-1.2. Antarctica

Antarctica has occupied high latitudes since at the Cretaceous or earlier, during which time it was connected with Australia (Kennett, 1977). Beginning in the early Eocene, Australia separated and drifted toward lower latitudes, and the margins rimming the Antarctic continent have existed as passive rifted margins since that time (Kennett, 1977).

Along the Weddell Sea continental margin, the character of seismic data is distinctly different between interpreted preglacial sediment and glacial sediment (Kuvaas and

Kristoffersen, 1991). This transition is characterized by an interpreted Albian-early

Oligocene unconformity, and is overlain by a submarine fan complex (the Crary Fan) interpreted to be Oligocene to recent in age, and interpreted to have grown in response to major climatic change and glaciation beginning at the Eocene-Oligocene boundary

(Kuvaas and Kristoffersen, 1991).

Similarly, in Prydz Bay, the character of seismic data is distinctly different between interpreted preglacial and glacial sediment (Kuvaas and Leitchenkov, 1992). As with the

Weddell Sea sediments, this shift is interpreted to indicate the initiation of turbidite deposition as a result of onset of glaciation in the late Eocene to early Oligocene (Kuvaas and Leitchenkov, 1992). Here, the post-Eocene-Oligocene boundary turbidites make up a succession of sediment up to 2.2 km in thickness (Kuvaas and Leitchenkov, 1992).

Rapid and voluminous Oligocene to Miocene deep-water turbidite deposition interpreted from seismic data along the East Antarcic continental margin is similarly inferred to have been directly produced by glacial activity and growth of the East Antarctic Ice Sheet, i.e.,

443 as a direct result of global climatic changes at the Eocene-Oligocene boundary (Close,

2010).

Overall, it should be noted that interglacial times since the late Eocene are characterized by little to no clastic sedimentation on Antarctic shelfs, with progradational shelf sequences, including fan systems, almost exclusively related to glacial activity (Cooper et al., 1991; Kuvaas and Kristofferson, 1991).

444

APPENDIX D-1.3. Asia

Since the Mesozoic and throughout the Cenozoic, convergence between the Eurasian Plate and the northward-moving African, Arabian, and Indian plates has resulted in the formation of the extensive Alpine-Himalayan orogenic belt (Dewey and Bird, 1970).

The Thrace Basin of northwestern Turkey evolved as an intermontane molasse basin from

Eocene to present (Siyako and Huvaz, 2007). Here, the Eocene-Oligocene boundary is marked by a transition from shallow and deep marine facies (including carbonates and proximal to distal turbidites) to terrestrial pro-delta facies, indicating progradation (Siyako and Huvaz, 2007).

Most central Anatolian basins of Turkey contain erosional periods coincident with the

Eocene-Oligocene boundary, and marking a shift from predominantly marine to nonmarine deposition (Görür et al., 1998). For instance, the Çankırı and Tuzgölü basins contain upper

Eocene-lower Oligocene conglomeratic strata overlying lower Eocene strata, the Şarkışla and Sivas basins contain upper Oligocene strata overlying lower Eocene strata, and the

Refahiye and Yıldızeli basins contain late Eocene-early Oligocene unconformities (Görür et al., 1998).

To the south, the Levantine Basin off the coast of Turkey, Syria, Lebanon, Israel, and Egypt contains a depositional hiatus marked by an abrupt transition from middle Eocene carbonate to Oligocene and Miocene siliciclastics (Gardosh and Druckman, 2006).

The Arabian Plate existed as part of the passive northern margin of Gondwana from the

Triassic through Early Cretaceous, while the northeastern portion of the plate has been influenced by development of the Zagros fold-thrust system since the Early Cretaceous

445

(Alsharhan and Nairn, 1995; Alavi, 2004). Deposition around much of the Arabian Plate is characterized by a major unconformity separating middle Eocene carbonates from

Oligocene to lower Miocene conglomerates and sandstones in South Iraq, Kuwait, Saudi

Arabia, and parts of Qatar (Alsharhan and Nairn, 1995). Alsharhan and Nairn (1995) attribute this unconformity to major late Eocene-early Oligocene regression. In Yemen, this unconformity is also evident, as evinced by a late Eocene to late Oligocene hiatus in carbonate deposition (Brannan et al., 1997). In the United Arab Emirates, this unconformity is not evident, with deposition of carbonates and marls continuous through the late Eocene to Oligocene (Alsharhan and Nairn, 1995).

The Early Cretaceous to recent Zagros fold-thrust belt constitutes a portion of the Alpine-

Himalayan orogenic belt and spans about 2000 kilometers of southeastern Turkey, northern

Syria and Iraq, and western and southern Iran (Alavi, 2004). In the Iranian portion of the fold-thrust system, the late Eocene to early Oligocene is marked by an unconformity separating middle to upper Eocene retrogradational carbonates and as flysch-system siliciclastics from Oligocene carbonates (Koop and Stoneley, 1982; Alavi, 2004).

However, it should be noted that Alavi (2004) attributed development of this unconformity to tectonic activity.

Since the Paleogene, siliciclastic deposition associated with the Indian subcontinent has predominantly consisted of sediments derived from the actively uplifting Himalaya in response to Indian-Eurasian plate convergence. Early Himalaya uplift was responsible for the inception of the two largest modern sedimentary systems on Earth: the Indus and the

Bengal fans (Qayyum et al., 1996; Curray et al., 2003; Ingersoll et al., 2003).

446

The proto-Indus fan – referred to as the Khojak fan – actively fed into the Neo-Tethys

Ocean Katawaz Basin of the northwest margin of the Indian subcontinent during the

Paleogene (Qayyum et al., 1996, 1997b; Carter et al., 2010). The majority of Khojak fan sections from the Katawaz Basin as exposed in Pakistan are of middle and late Eocene to early Miocene age (Qayyum et al., 1996; Qayyum et al., 1997a; Qayyum et al., 2001; Carter et al., 2010), however, Khojak strata in the Katawaz Basin as exposed in Afghanistan and in the Makran region of Iran have yielded ages as old as late Paleocene (Qayyam et al.,

1997b; Carter et al., 2010). Inception of the Indus fan, and correspondent delivery of

Himalaya-derived sediment into the Indian Ocean, was not initiated until the early Miocene

(Davies et al., 1995; Qayyum et al., 1997b).

The western margin of India has existed as a passive margin since the Mesozoic breakup of Gondwana, and experienced slow, long-lived subsidence from the Mesozoic to the late

Oligocene (Biswas, 1982, 1987; Whiting et al., 1994). Numerous rift basins developed along this margin, though most of these only contain a few hundred meters of Cenozoic sediment (Biswas, 1982). The Eocene-Oligocene boundary in the Kutch Basin is characterized by an unconformity attributed to regression, with shallow marine Oligocene limestone and shale unconformably overlying shallow marine middle Eocene limestone and minor shale (Biswas, 1982, 1987). In the Cambay Basin, no unconformity is noted, with the upper Eocene to Oligocene consisting of marine and deltaic shale, sandstone, and limestone deposited during oscillatory transgression/regression cycles (Biswas, 1982,

1987).

The eastern margin of India has also existed as a passive margin since Gondwana breakup, and similarly saw development of numerous post-breakup rift basins (Sastri et al., 1981).

447

In Cauvery Basin, the Eocene-Oligocene section is some 1.5 km thick and is made up of marine shale, sandstone, and minor limestone (Sastri et al., 1973). No end-Eocene unconformity is noted in the Cauvery Basin, however, the Krishna Godavari Basin does contain an end-Eocene unconformity that separates Eocene carbonates and volcanics from

“pre-Miocene” claystone and sandstone (Sastri et al., 1973).

The Bengal submarine fan of the northeast Indian margin may have seen inception as early as the earliest Eocene, however, the primary influx of voluminous deepwater sediment occurred during the middle Eocene to early Miocene in response to collision of the Indian subcontinent with the Burma and Tibetan blocks (Alam et al., 2003; Curray et al., 2003;

Curray, 2014). A late Eocene-early Oligocene unconformity – attributed to basin-wide marine regression – is evident in the Bengal Basin foredeep (Alam, 1989), as well as in the basin stable shelf province (Alam et al., 2003).

The Andaman Islands in the Bay of Bengal lie at the interface of the Indian Plate subducting beneath the Australian Plate along the Java trench (Curray, 2005). Here, the so-called

Andaman Flysch Group was deposited in an active margin forearc setting during the

Oligocene (and perhaps as early as the late Eocene) as an inferred response to a tectonically-sourced sediment supply (Chakraborty and Pal, 2001; Bandopadhyay and

Ghosh, 2015).

There are over 60 Cenozoic basins in Thailand, but most of these developed beginning in the late Oligocene (Polachan et al., 1991). These basins are predominantly filled with upper

Oligocene and younger nonmarine sediment (Polachan et al., 1991).

448

In the South China Sea region, the timing of rifting remains controversial, but may have begun as early as the latest Cretaceous or the Paleocene, with sea-floor spreading as early as the early and middle Eocene (Taylor and Hayes, 1983; Ru and Pigott, 1986). Most syn- rift deposition throughout the basin is characterized by nonmarine facies. Syn-rift sediments of offshore southern and central Vietnam consist of nonmarine sediments of possible Eocene to Oligocene age (Lee and Watkins, 1998; Lee et al., 2001b), while syn- rift sediments off northern Vietnam and southern China consist of Eocene nonmarine

(largely lacustrine) sediments grading upward into upper Oligocene marginal marine sediments (Chen et al., 1993). Syn-rift sediments in the Taiwan region consist of upper

Paleocene to upper Eocene largely nonmarine sediments (Lin et al., 2003). The “Oligocene breakup unconformity” is ubiquitous around the Taiwan region, and marks a period of missing stratigraphic section spanning at least 37 to 30 Ma (Lin et al., 2003; Huang et al.,

2012). Similarly, Wu (1994) notes a hiatus between Eocene and late Oligocene to early

Miocene sedimentary environments across the South China Sea region, and suggests the presence of a regional 40 to 32 Ma uplift event or sea-level drop. Holloway (1982) notes a hiatus in the Palawan region of the Philippines, though this hiatus consists of missing middle to upper Oligocene sediment, rather than upper Eocene and lower Oligocene sediment.

Intracontinental basins in eastern China and basins of China’s eastern marginal seas developed as extensional basins in the Cenozoic (Ren et al., 2002). These rift basins are predominantly filled with nonmarine (primarily alluvial and lacustrine) sediments, and in basins including the Songliao, Bohaiwan, Pearl River Mouth, Qiongdongnan, and

Yinggerhai basins, signal of the Eocene-Oligocene transition does not seem to be preserved

449 in the form of notable unconformities, though in all of these basins aside from the Songliao, an unconformity of middle and/or late Eocene age is noted (Ren et al., 2002). The Ulleung rift basin of the Sea of Japan opened beginning in the early Oligocene, and most sedimentary fill is of Miocene or younger age (Chough and Barg, 1987; Lee et al., 2001a).

Similarly, the Kurile rift basin of the Okhotsk Sea opened in the early Oligocene to middle

Miocene, and contain upper Oligocene and younger strata (Baranov et al., 2002; Karp et al., 2006).

The Arctic Ocean and its Amerasia and Eurasia basins are rimmed by continental margins most likely rifted sometime in the late Mesozoic or early Cenozoic (Alvey et al., 2004;

Franke et al., 2004 and references therein). The East Siberian Shelf and Laptev Shelf of this region are virtually unexplored, while the deepwater Amerasia basin – which includes the Canada and Makarov basins – is even more poorly understood (Franke et al., 2004).

Along the East Siberian Shelf, parallel reflectors observed in seismic data at water depths in excess of 1 km may represent turbiditic sandstones and clays of Late Cretaceous to

Miocene age outboard of shelfal clinoforms inferred to represent a deltaic sequence of

Paleogene to early Miocene age (Sekretov, 2001). Along the Laptev Shelf, rift basins are filled with a thick succession of sediment of inferred Late Cretaceous to recent age

(Drachev et al., 1998; Franke and Hinz, 2005).

450

APPENDIX D-1.4. Australia/Oceania

In Borneo, initiation and rifting of the Kutai Basin is poorly constrained but is generally inferred to have occurred in the middle Eocene (Moss et al., 1997; Moss and Chambers,

1999). Regional subsidence occurred throughout the basin in the late Eocene, and a sag phase of deposition — characterized by limited coarse clastic input into the basin — dominated from the late Eocene to the late Oligocene (Moss and Chambers, 1999). Late

Eocene through early Oligocene deposition was apparently continuous, with no major unconformity noted for this time period (Moss et al., 1997; Moss and Chambers, 1999).

Throughout Papua New Guinea, an “Oligocene disconformity” is noted, which spans at least the early Oligocene, and may locally encompass the late Eocene through early

Miocene (Pigram et al., 1989; Pigram and Symonds, 1991; van Ufford and Cloos, 2005).

Van Ufford and Cloos (2005) reinterpret this unconformity – which they note has been documented across much of the Australian continental shelf – as a direct product of the dramatic sea-level fall at the Eocene/Oligocene boundary. Overall, Paleocene to recent deposition in the Papua region is dominated by shallow marine carbonates (Home et al.,

1990; van Ufford and Cloos, 2005), and van Ufford and Cloos (2005) note that the only siliciclastic unit of appreciable thickness in the western and central Papua region is a quartz-rich sandstone unit of some 10-100-m thickness that overlies the Oligocene disconformity, and which they interpret to be a basal transgressive unit.

In the New Caledonia Basin, which opened some time in the Cretaceous or Paleocene and lay in an active margin setting in the Eocene and Oligocene, the Eocene-Oligocene unconformity is regionwide (Collot et al., 2008), and is also present in the adjacent Fairway

Basin (Exon et al., 2007; Nouzé et al., 2009). 451

All around the New Zealand margin, which largely existed as a passive margin at the

Eocene-Oligocene boundary, the so-called Marshall Paraconformity has been recognized, though this feature has been attributed to mid-Oligocene sea-level fall rather than the sea- level fall coinciding with the Eocene-Oligocene boundary (e.g., Fulthorpe et al., 1996). In southern New Zealand’s Fiordland, presence of an early Oligocene submarine fan complex has been noted, though here a phase of tectonically active flysch-style deposition is invoked

(Carter and Lindqvist, 1975).

Most of the rifted Australian continental margin has existed as a passive margin since the

Cretaceous (e.g., von Rad and Exon, 1983; Seidl et al., 1996), though in the northeast of

Australia rifting progressed through the early Cenozoic (Davies et al., 1989), and in the

Tasman region, Australia-Antarctica rifting began in the Late Cretaceous (Exon et al.,

2004).

At the Eocene-Oligocene boundary, the Tasmanian Gateway between Australia and

Antarctica opened (Exon et al., 2004; Stickley et al., 2004). This opening is in fact cited by some as being partly responsible for climatic cooling at the Eocene-Oligocene boundary

(Stickley et al., 2004; Bijl et al., 2013). Notably, late Eocene-early Oligocene hiatuses are evident across the Tasman region, including the west Tasmanian margin’s Strahan Sub- basin, the South Tasman Rise, the East Tasman Plateau, and the Tasman Basin (Exon et al., 2004). Paleocene to Eocene sediment in the region includes predominantly silty claystone and some clastics, while Oligocene and younger sediment mostly consists of carbonate (Exon et al., 2004).

452

Throughout Southern Australia’s passive margin Eucla Basin, an early Oligocene unconformity is observed, separating middle-upper Eocene carbonate and upper Eocene siliciclastics from Miocene carbonate (Hou et al., 2006).

Along Australia’s passive western margin, most Cenozoic sedimentation has consisted of carbonates (Quilty, 1977; Apthorpe, 1988). Off Australia’s northwest shelf, the Late

Cretaceous to present has been dominated by limestone and marl deposition (Apthorpe,

1988; Butcher, 1990; Baillie et al., 1994). Notably, a widespread end-Eocene/beginning of the Oligocene unconformity is evident across the entire northwest shelf region, and spans at least the early Oligocene and often all of the Oligocene and parts of the early Miocene

(Quilty, 1977; Apthorpe, 1988).

The northeast of Australia was subject to Late Cretaceous through early Cenozoic rifting, but was tectonically stable in the middle Cenozoic (Davies et al., 1989). Nonetheless, here too, a widespread late Eocene-early Oligocene unconformity is observed, including in the

Gulf of Papua, on the Queensland Plateau, and in Capricorn Basin (Davies et al., 1989).

453

APPENDIX D-1.5. Europe

Throughout Europe, the opening of the Atlantic Ocean (Pitman and Talwani, 1972) and the failed rifting of the North Sea led to a major transgressive phase during the late

Paleocene and early Eocene, however, the middle Eocene to early Oligocene is characterized by a general regressive phase (Hardenbol et al., 1998).

Along the west Svalbard margin off of the Svalbard mainland, rifting led to development of a number of basins in the early to middle Oligocene (Steel et al., 1985). Some of these basins are stated to contain lower Oligocene or even upper Eocene siliciclastic sediment, though it is noted that deposition of these successions was in direct response to an intensive tectonic regime (Livšic, 1992), perhaps in part a product of earliest Oligocene shifts in relative plate motions in the region (e.g., Faleide et al., 1991).

South of the Svalbard mainland, the Barents Sea likewise evolved in a transform setting through Late Cretaceous to Paleogene rifting (Faleide et al., 1993, 2008). Throughout much of the Barents Sea region – including the Hammerfest, Nordkapp, South Barents, and

Timan-Pechora basins and the Bjarmeland Platform – most sediment of Cenozoic age is missing, with Quaternary sediment resting unconformably upon Cretaceous rock

(Henriksen et al., 2011). This absence is attributed to uplift and glacial erosion (Henriksen et al., 2011). Nonetheless, the western Barents Sea does contain thick successions of lower

Eocene to Oligocene sediment, particularly in the Tromsø Basin (Spencer et al., 1984;

Henriksen et al., 2011). However, along this western margin, Ryseth et al. (2003) note a minor time-stratigraphic break and marine shallowing at the Eocene-Oligocene boundary in the Sørvestsnaget Basin of the southwestern Barents Sea, which they mention may be attributable to tectonism accompanying plate rearrangements. 454

In the Norwegian Sea, tectonism and substantial plate rearrangements are also invoked, and generally correlated temporally with the Eocene-Oligocene boundary, while the succeeding Oligocene was characterized by tectonic quiescence (Henriksen et al., 2005).

Throughout the Norwegian Sea, and including the Møre and Vøring basins, a widespread unconformity is present at the Eocene-Oligocene boundary (Eidvin et al., 1998, 2014;

Henriksen et al., 2005). Generally, lower Oligocene to Pliocene deltaic sand, as well as marine claystone, sits atop lower to middle Eocene marine claystone (Eidvin et al., 1998,

2014; Henriksen et al., 2005).

As in the Norwegian Sea, the North Sea is characterized by a widespread and prominent unconformity at the Eocene-Oligocene boundary (Jordt et al., 1995; Huuse and Clausen,

2001; Stoker et al., 2005; Anell et al., 2012; Eidvin et al., 2014). This unconformity has been attributed to tectonic uplift of Fennoscandia (Jordt et al., 1995), however, later work suggested that the unconformity was instead primarily attributable to global climatic and glacioeustatic changes (Clausen, 1998; Clausen et al., 1999; Huuse, 2002). Due to flooding and general Eocene tectonic quiescence in the North Sea, most of the Eocene interval is characterized by diminished sediment supply, including in particular a dearth of coarse- grained siliciclastics (Anell et al., 2012). Directly following the Eocene-Oligocene boundary, however, there was a dramatic increase in sediment supply delivered from southern Norway, with major sand influx into the North Sea, as well as some amount of erosion of the basin floor (Huuse and Clausen, 2001; Rundberg and Eidvin, 2005).

The Faroe-Shetland region and offshore western Ireland’s Rockall Trough region, as with the European Atlantic margin in general, developed in response to extension associated with the opening of the Atlantic (Stoker et al., 2005). The regions experienced widespread

455 influence of tectonism in the Eocene and Oligocene, including late Eocene to Oligocene subsidence as well as compressional activity (Andersen et al., 2000; Stoker et al., 2005;

Ritchie et al., 2008). Throughout the Faroe-Shetland and Rockall Trough regions, as well as throughout the Porcupine Basin of offshore western Ireland, an unconformity is present at the Eocene-Oligocene boundary (Croker and Klemperer, 1989; Davies et al., 2004;

Stoker et al., 2005). The Eocene through Oligocene succession in the Faroe-Shetland Basin is characterized by claystone, with sand bodies restricted to fans of middle Eocene age

(Davies et al., 2004; Shoulders et al., 2007). In the Porcupine Basin, the Eocene is characterized by sandy sediments deposited as deltaic sequences, whereas the Oligocene generally consists of deep marine shale (Croker and Klemperer, 1989; Shannon, 1992).

The Eocene-Oligocene boundary unconformity in the Porcupine Basin involved significant amounts of erosion and channeling (Croker and Klemperer, 1989).

The Pyrenees Mountains developed as a West-East trending orogenic belt in response to

Iberian-Eurasian plate convergence that lasted from the Late Cretaceous through the majority of the Cenozoic (Labaume et al., 1985). In the southern Pyrenees, syntectonic unconformities of Eocene to Oligocene age have been identified and associated with numerous conglomeratic deposits (Riba, 1976). Ages of these unconformities are not precisely constrained, and the mechanism for these unconformities is interpreted to be orogenic activity, nonetheless, Riba (1976) states that the age of the middle unconformity is likely late Eocene (thus, potentially coincident with the Eocene-Oligocene boundary).

Similarly, erosion and high-volume conglomeratic deposition in the central Pyrenees is associated with a period of inferred rapid late Eocene to early Oligocene exhumation

(Fillon et al., 2013). The Aquitaine Basin of the northern Pyrenees and the offshore Bay of

456

Biscay, including the offshore Parentis sub-basin, contains Paleocene to Eocene shales and sandstones separated from Oligocene shales by a latest Eocene unconformity (Biteau et al.,

2006; Ferrer et al., 2012).

Earliest Jurassic breakup and rifting of Pangaea marked the beginning of the opening of the Atlantic Ocean basin (Dietz and Holden, 1970). Rifting of the west Iberian margin of the northeast Atlantic occurred from the Late Jurassic through Cretaceous (Alves et al.,

2003). This rifted margin existed as a subsiding passive margin until the Pyrenean orogeny

(see above) resulted in Eocene tectonism and significant uplift along the northwestern

Iberian margin (Boillot et al., 1979; Munoz et al., 2003). Cenozoic sedimentary infill of basins along the western Iberian margin is relatively poorly understood, with stratigraphic studies of the region traditionally relying on seismic data, and thus, understanding of the

Eocene-Oligocene transition as might be manifest in sediments is not documented for basins including the Galicia Interior Basin of northwest Iberia (Murillas et al., 1990) and the Alentejo Basin of southwest Iberia (Alves et al., 2003). There is, however, an unconformity documented at the Eocene-Oligocene boundary in the Lusitanian Basin of central western Iberia, where it separates Paleocene to upper Eocene dolomite from upper

Oligocene to recent marine sandstone (Pinheiro et al., 1996; Alves et al., 2006). Along the southern Iberian margin, as well as along the Catalan margin and the northwest

Mediterranean Sea, evidence of an Eocene-Oligocene boundary unconformity is present, including an Eocene-Oligocene unconformity in the Alicante Trough of the Betic

Cordillera (Guerrera et al., 2006), an Eocene-Oligocene unconformity separating Eocene lacustrine carbonates from Oligocene marine facies on the island of Mallorca (Ramos et al., 2001), and the unconformable juxtaposition of upper Eocene or lower Oligocene and

457 younger continental sediments atop pre-Cenozoic basement rocks along the Catalan margin

(Roca et al., 1999).

Along the northern Tethyan (i.e., Helvetic or present-day southern European) margin,

Cretaceous through Cenozoic deposition was largely controlled by tectonism related to the

Alpine orogeny (Schmid et al., 2008). This is manifest in widespread flysch deposited throughout the Cretaceous through Oligocene (e.g., Trümpy, 1960), thus potentially overwhelming any sea-level related record of the Eocene-Oligocene transition. Offshore of this margin, marine back-arc basins of the Mediterranean Sea have developed mostly since the Oligocene, and thus, stratigraphic record of the Eocene-Oligocene boundary is generally lacking (Rosenbaum et al., 2002). Nonetheless, record of this boundary does occur. For instance, in the Ionian forearc basin of southern Italy, lower Oligocene conglomerate lies unconformably atop the basement complex (Cavazza and Ingersoll,

2005).

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APPENDIX D-1.6. North America

APPENDIX D-1.6.1. Northern America

The Cordilleran margin, spanning from Baja California and the Gulf of California in the south to Alaska’s Kodiak Island in the north, was an Andean-type subduction margin from the late Mesozoic through the end of the Eocene (Dickinson, 1976). This convergence slowed in the late Eocene through Oligocene (Bartow, 1991).

Basins along much of the California continental margin record an Eocene-Oligocene unconformity and a general transition from deep-marine Eocene sedimentation to predominantly non-marine Oligocene sedimentation (e.g., a shallowing of several km)

(Nilsen and Clarke, 1975; Graham, 1976; Bartow, 1978; Nilsen, 1984; Loomis and Ingle,

1994). Authors generally agree that this unconformity and the shift from marine to nonmarine deposition in California is attributable to Eocene-Oligocene tectonic events (e.g.,

Graham, 1976; Nilsen, 1984; Loomis and Ingle, 1994), though workers also point out that the unconformity and shift in deposition is at least partly attributable to climatic change beginning in the late Eocene and resulting in colder, more arid Oligocene climatic conditions and major sea-level fall (Bartow, 1978; Ingle, 1981b; Nilsen, 1984). The

Eocene-Oligocene unconformity is manifested in basins of southern California, including the Santa Maria, Santa Ynez, Soledad, and Los Angeles basins, whereby Oligocene redbeds unconformably overly lower Cenozoic deep-marine strata (Bartow, 1978; Ingle,

1981a; Nilsen, 1984). The unconformity is also present in basins of Salinian block affinity, including in the northern Santa Lucia Range, where the non-marine lower Oligocene Berry

Formation unconformably to conformably overlies the upper Eocene The Rocks Sandstone or unconformably overlies older Cenozoic rocks or crystalline basement rocks (Link and 459

Nilsen, 1980). Similarly, Chipping (1972) notes that the region encompassing the southern

Santa Lucia Range, the La Panza Range, and the Sierra Madre Range was emergent in the

Oligocene, meaning much of the upper Eocene is missing and meaning Oligocene and

Miocene strata unconformably overly Late Cretaceous to late middle Eocene strata. Further north, the Gualala Basin of contentious Salinian block affinity records a middle Eocene to late Oligocene unconformity separating deep-water turbidites of the Paleocene–middle

Eocene German Rancho Formation from the upper Oligocene–lower Miocene Iversen

Basalt (Nilsen and Clarke, 1975; Loomis and Ingle, 1994). Loomis and Ingle (1994) note that they believe the Gualala Basin unconformity to be attributable to tectonism rather than a fall in sea level.

Further inland, the Great Valley of California evolved as a forearc basin in the late

Mesozoic and early Cenozoic, and as a transform margin basin in the later Cenozoic

(Dickinson and Seely, 1979; Ingersoll, 1979; Bartow, 1991). Throughout the San Joaquin and Sacramento basins that comprise the Great Valley, a regional unconformity separates upper Eocene and/or lower Oligocene strata (the Kreyenhagen and Tumey formations) from strata of Oligocene or Miocene age (Anderson and Pack, 1915; Milam, 1985; Moxon,

1990; Hosford Scheirer and Magoon, 2007; Johnson and Graham, 2007). Milam (1985) attributes the unconformity at the top of the Kreyenhagen Formation to subaerial erosion caused by early to middle Oligocene tectonic uplift, however, Johnson and Graham (2007) speculate that the late Eocene or early Oligocene internal unconformity within the Tumey

Formation may be the result of rapid sea-level fall.

Non-marine conditions persisted throughout the Eocene and Oligocene in the northern

Sierra Nevada inland of the northern Great Valley (Cassel et al., 2009; Cassel and Graham,

460

2011). Nonetheless, a widespread erosional unconformity occurs at the top of the Eocene, separating Eocene fluvial deposits from interbedded Oligocene volcanics and fluvial deposits (Cassel et al., 2009; Cassel and Graham, 2011).

Coastal northern California is generally dominated by accreted terranes containing deep marine rocks of Eocene and older age (e.g., the Franciscan Coastal belt and Yager complex)

(Evitt and Pierce, 1975; Bachman, 1978; McLaughlin et al., 1982; Dumitru et al., 2013).

An exception is the King Range terrane, which contains rocks of Cretaceous through

Miocene age (McLaughlin et al., 1982). Detailed stratigraphy of the King Range terrane, including potential occurrence of an Eocene-Oligocene unconformity, is not discussed by these workers. In the onshore Eel River Basin, the Eocene and older Yager complex constitutes basement and is separated from the youngest (Miocene) basin fill by an unconformity (Ogle, 1953; McLaughlin et al., 1982; Clarke, 1992).

The Tyee forearc basin of western Oregon is filled with several kilometers of predominantly Eocene-age sediment (Chan and Dott, 1983; Dumitru et al., 2013). Upper middle to upper Eocene strata were predominantly deposited in deltaic and shoreface environments, with deltaic to fluvial conditions prevailing locally in the late Eocene (Ryu et al., 1992). Throughout the Tyee Basin, the upper contact of these middle to upper Eocene strata has been removed due to erosion, and throughout the Oligocene and Miocene, calcalkaline volcanism of the Western Cascade arc persisted (Ryu et al., 1992). An exception is the Coos Bay Basin at the southwest boundary of the Tyee Basin. Here, upper

Eocene to upper Miocene deep-marine, shallow-marine, and fluvial sandstone and mudstone were deposited (Baldwin and Beaulieu, 1973; Ryu et al., 1992). Deposition is generally interpreted to have occurred conformably between the upper Eocene rocks of

461

Coos Bay and the upper Eocene to lower Oligocene Bastendorff and Tunnel Point formations until ~33 Ma (Baldwin and Beaulieu, 1973; Prothero and Donohoo, 2001).

However, some workers have mentioned the possibility of an as-of-yet unidentified ~2.5 m.y. hiatus separating the Bastendorff Formation from the Tunnel Point Formation (i.e., the possibility of an Eocene-Oligocene unconformity) (Prothero and Donohoo, 2001).

Regardless, the lower Oligocene (or possibly middle Oligocene; Prothero and Donohoo,

2001) Tunnel Point Formation itself is overlain by an unconformity separating the Tunnel

Point from beds of Miocene age (Baldwin and Beaulieu, 1973).

As in Oregon, Paleogene volcanism occurred throughout much of Washington (Lipman et al., 1972; Wells et al., 1984). While some regions contain relatively continuous sections of

Eocene through Oligocene marine sediment, others preserve evidence for an Eocene-

Oligocene unconformity, including areas of the Olympic Peninsula, where at least some upper Eocene and most lower and middle Oligocene strata are missing, and the Bellingham area, where most of the Oligocene is missing (Armentrout et al., 1983; Snavely, 1987;

Snavely and Kvenvolden, 1989). Some authors have speculated that missing strata of latest

Eocene to middle Oligocene age in the Olympic Peninsula area may have been subducted during the late Oligocene (Snavely and Kvenvolden, 1989).

During the late Eocene, interaction between the Pacific and North American plates in the present-day Vancouver and Queen Charlotte islands areas and the Gulf of Alaska shifted from orthogonal convergence to strike-slip motion (Rohr and Dietrich, 1992). Few rocks of Eocene and Oligocene age have been described in either the onshore or offshore regions of the Vancouver and Queen Charlotte islands (Rohr and Dietrich, 1992), nonetheless, the

Tofino Basin of offshore Vancouver Island does contain some marine sediment

462

(predominantly siltstone and shale) of Eocene–Oligocene age (as encountered in the

“Pluto” well), and some rock of this age crops out on Vancouver Island itself (Shouldice,

1971, 1973; Tiffin et al., 1972). Shouldice (1971, 1973) notes that while complete detail is lacking, Tofino Basin seems to have been characterized by extensive early and middle

Eocene submarine volcanism followed by a period of uplift and then subsidence in the late

Eocene, and transgression throughout the Oligocene and early Miocene. Notably,

Shouldice (1971, 1973) includes a stratigraphic cross section displaying an unconformity separating Eocene marine strata from Oligocene marine strata in the Tofino Basin, but does not discuss this unconformity in detail, while Stewart (1976) notes an Eocene-Oligocene unconformity preserved on western Vancouver Island. By contrast, the Queen Charlotte

Basin just to the north contains sediment no older than Miocene in age (Shouldice, 1971,

1973; Rohr and Dietrich, 1992).

Cenozoic deep-water deposition atop the Pacific Plate in the southern and central Gulf of

Alaska has consisted of the emplacement of three large submarine fans, the oldest of which

(the Zodiac fan of the Aleutian Abyssal Plain) is estimated to be late Eocene to late

Oligocene in age, or ~42–24 Ma (Hamilton, 1967; Creager et al., 1973; Stevenson et al.,

1983; Stevenson and Embley, 1987). Deep-water turbidite deposition here is inferred to have been relatively continuous from the late Eocene through Oligocene, with no unconformities mentioned (Creager et al., 1973; Stevenson and Embley, 1987), though it should be noted that recovery of the DSDP core upon which these interpretations are based was characterized by only ~30% recovery (Creager et al., 1973). Stevenson et al. (1983) postulate that the volume of sediment contained within the fan requires that an integrated drainage of substantial size (at least 500,000 km2) was active during this time.

463

The Alaskan onshore at the northern edge of the Gulf of Alaska contains record of an

Eocene-Oligocene unconformity. For instance, the Malaspina district has an unconformity of approximate late Eocene age, whereby non-marine to shallow marine strata of the

Eocene Kulthieth Formation are topped by an unconformity (Plafker, 1971, 1987), and similarly, the nearby Yakataga district preserves an Eocene-Oligocene unconformity atop the Kulthieth Formation (Plafker, 1971; Stewart, 1976). On Kodiak and Trinity islands at the northern edge of the Gulf of Alaska and just outboard of the Alaska Peninsula, an unconformity occurs at the Eocene-Oligocene boundary and separates the Eocene

Sitkalidak Formation from the Oligocene Sitkinak Formation (Plafker, 1971; Stewart,

1976; Moore and Allwardt, 1980). The intensely deformed Sitkalidak Formation is characterized by interbedded deep-marine sandstone and shale making up a turbidite succession up to several km thick, while the overlying, less deformed Sitkinak Formation is generally shelfal marine to non-marine, in some areas containing basal conglomerate overlain by interbedded marine sandstone and shale comprising a turbidite succession, and in others containing fluvial conglomerate and sandstone interbedded with coal-bearing strata (Nilsen and Moore, 1979). Moore and Allwardt (1980) note that the unconformity is attributable to deformation and uplift of the Sitkalidak Formation prior to deposition of the

Sitkinak Formation. Just inland of Kodiak Island, an Eocene-Oligocene unconformity is also document in the Shelikof Strait and Cook Inlet (Stewart, 1976).

Stratigraphy of the Bering Sea’s deep-water Aleutian Basin remains poorly constrained, with Miocene and older stratigraphy particularly poorly understood (Cooper et al., 1979;

Martin et al., 2019).

464

The Bering shelf includes five large sedimentary basins—the Bristol, St. George, Norton,

Navarin, and Anadyr basins (Marlow et al., 1994). An Eocene-Oligocene unconformity is apparent in Bristol Basin, which spans part of the Alaska Peninsula as well as Bristol Bay

(Galloway, 1974). This unconformity separates the Paleocene–Eocene Tolstoi Formation from the Oligocene Stepovak Formation (Galloway, 1974), though it is noted that in the onshore Alaska Peninsula, it is often difficult to distinguish these two units, both of which are characterized by marine coarse volcaniclastics and black siltstones (Burk, 1965). In the

Norton Basin, an undated Paleogene unconformity is documented, with sediment of Late

Cretaceous and Paleogene age inferred to be predominantly nonmarine and sediment of

Neogene age inferred to be marine (Fisher et al., 1982). An unconformity is inferred (on the basis of well log, lithological, geochemical, and paleontological data) to separate the marine Eocene and Oligocene sections in the Navarin Basin, though it is noted that geophysical evidence for such an unconformity is lacking (Turner et al., 1984). In the

Anadyr Basin, upper Eocene and Oligocene sediment is described as nonmarine by Fisher et al. (1982), though McLean (1979) described nonmarine as well as marine and paralic sediment of Late Cretaceous to Oligocene age. McLean (1979) also note post-early Eocene erosion, however, documentation of more detailed stratigraphy, including possible unconformities, is lacking.

In Hope Basin of the Chukchi Sea to the north, ages and lithologies are unknown, but most

Cenozoic fill is inferred to be nonmarine clastic sediment (Grantz et al., 1975; Tolson,

1987). Throughout the Chukchi and Beaufort seas off northern Alaska, the Hope, North

Chukchi, and Camden basins contain poorly constrained stratigraphy, but are inferred to contain marine and nonmarine clastic rocks of Cretaceous and Cenozoic age (Grantz et al.,

465

1975). Sufficient detail is currently unavailable to assess whether or not an Eocene-

Oligocene boundary unconformity may be expressed in these basins — recent seismic acquisition on the Chukchi Shelf (United States portion) provided a more detailed understanding of Paleozoic and Mesozoic stratigraphy in the region, though discussion of the Cenozoic section was restricted to the identification of a base Cenozoic horizon in the seismic data, and mention of the shelf’s post-rift Lower Cretaceous to Cenozoic Brookian megasequence evident from existing well data from the shelf (Kumar et al., 2011). In the

Mackenzie Basin of the southern Beaufort Sea, an upper Paleocene to Eocene delta sequence is separated from an Oligocene delta sequence by thick, widespread deposits of shale (Willumsen and Cote, 1982). Willumsen and Cote (1982) make no mention of an

Eocene-Oligocene boundary unconformity (instead mentioning a major unconformity of early to middle Miocene age), but Harrison et al. (1999) not an unconformity of locally angular character in the Caribou Hills and Mackenzie Delta areas that occurs around the time of the early Oligocene. In general, Harrison et al. (1999) describe the early Oligocene as a time during which present-day off-shelf areas saw widespread lowstand deposition.

The Sverdrup Basin of Arctic Canada was characterized by widespread nonmarine deposition from the Late Cretaceous through at least the Miocene (Balkwill, 1978). On the polar margin of Canada and Greenland, a distinction transition occurs from ~65–47 Ma dominance of tectonic influence on deposition to ~47 Ma to present dominance of global climatic influence on deposition, though the Eurekan Orogeny was active throughout the region at least until the end of the Eocene (Harrison et al., 1999). It is inferred that the late middle and late Eocene were characterized by widespread deformation, uplift, and erosion

(Harrison et al., 1999). In the onshore realm of eastern Arctic Canada and northern

466

Greenland, it is noted that, almost ubiquitously, the entirety of the Neogene record is missing (Harrison et al., 2011).

On the continental shelves of the Labrador Sea, three major unconformities are identified in the Cretaceous and Paleogene based on seismic and well data, including an unconformity of early Oligocene age that is noted to be coincident with the end of the Eurekan Orogeny

(Whittaker et al., 1997; Harrison et al., 2011). For instance, a distinct seismic marker of estimated Eocene-Oligocene boundary age occurs in the Melville Bay area of offshore northwest Greenland, flanking Baffin Bay — although in this particular area, well control is lacking (Whittaker et al., 1997). Harrison et al. (1999) note that, using these particular seismic indicators from the Melville Bay region, it has been interpreted that the lower

Oligocene and younger strata are separated by a significant angular unconformity from the underlying, presumably Eocene-age structures. In the Baffin Bay in general, sea-floor spreading is inferred to have ceased at about the time of the Eocene-Oligocene boundary

(the end of an Eocene syn-drift phase), followed by a shift to simple progradational sedimentation (the start of a Neogene post-drift phase) that delivered over 6 km of sediment to the basin during the Neogene (Whittaker et al., 1997; Harrison et al., 2011).

In the Grand Banks region of Newfoundland, the Jeanne d’Arc Basin and adjacent margin contain Paleocene to Pliocene fill and Paleocene to Oligocene-Miocene fill, respectively, that consists almost entirely of uninterrupted shales of the Banquereau Formation (Grant et al., 1986; McAlpine, 1990). Similarly, in the Scotian Shelf and Scotian Basin, mudstone of the Banquereau Formation makes up Paleocene to Pliocene fill (McIver, 1972; Grant et al., 1986).

467

Generally, understanding of western North Atlantic (i.e., eastern margin of the United

States and much of Canada) Paleogene deep-water deposition is limited due to a lack of exploration wells in this region (i.e., a lack of drilling to sufficient depth to investigate

Paleogene successions).

Along the New Jersey margin, the Cretaceous–Eocene section consists of flat-lying strata overlain by prominent clinoform geometries of Oligocene to recent age (Katz and Miller,

1996; Miller et al., 1998; Steckler et al., 1999). Eocene sediments consist primarily of clastic sediment-starved deep-water carbonates topped by a hiatus separating upper Eocene sediment from overlying upper Oligocene coarse siliciclastics, with the change in depositional regime speculated to have resulted from changes in onshore uplift and denudation rates, sea-level fall, and/or climatic change (Katz and Miller, 1996; Miller et al., 1998; Steckler et al., 1999). To the south, in the region including the Baltimore Canyon

Trough and Hatteras Basin, Paleocene–early Miocene sedimentation rates were an order of magnitude lower than preceding or succeeding time periods, and suggest a dearth of deep- water clastic deposition during this time, which has been attributed to a lack of tectonic activity in the onshore Appalachian region (Poag and Sevon, 1989). Along the continental margin from North Carolina to Florida, including the Blake Plateau region, a major hiatus of approximate mid-Oligocene age is present (as with the major mid-Oligocene hiatus in the New Jersey transect; Miller et al., 1985), but no mention is made of an unconformity at the Eocene-Oligocene boundary (Shipley et al., 1978; Vail et al., 1980; Locker and

Laine, 1992).

In the Gulf of Mexico, the late middle Eocene to late Eocene was a renewed period of major clastic deposition due to uplift of the Mexican Cordillera, and was characterized by

468 significant coastal, deltaic, and overall shelfal progradation, with only the far eastern reaches of the basin starved of sediment (Galloway et al., 2000). Toward the latest Eocene, transgression resulted in regional deposition of the thin Moodys Branch Formation (e.g.,

Elder and Hansen, 1981; Galloway et al., 2000), which was then followed in the latest

Eocene by the Jackson episode of clastic sedimentation along the western and northwestern margins of the Gulf, and mud-rich progradation in the central Gulf (Galloway et al., 2000).

This was followed by a period of exceptionally voluminous clastic sediment supply characterizing the Oligocene (Galloway et al., 2000). Record of an Eocene-Oligocene boundary unconformity in clastic sections of the basin is not evident (Galloway et al.,

2000).

469

APPENDIX D-1.6.2. Central America and The Caribbean

The Caribbean Plate has experienced eastward movement relative to the North America and South America throughout the Cenozoic (Malfait and Dinkelman, 1972; Burke et al.,

1984; Meschede and Frisch, 1998; Pindell and Kennan, 2009).

Along the northern boundary of the Caribbean Plate (i.e., the North America–Caribbean plate interface) sits the Greater Antilles volcanic arc, which includes present-day Cuba and

Hispaniola (Pindell and Kennan, 2009). In northern Hispaniola, regionally extensive

Paleocene to early Eocene sedimentation within a forearc basin setting (and bounded by subducting Atlantic oceanic crust to the north and the Hispaniola volcanic arc to the south) consisted of fine-grained deep-water hemipelagic turbidites, and was terminated by basin- wide folding and uplift during the middle Eocene (de Zoeten and Mann, 1999). This was followed by late Eocene to early Miocene deposition in elongate basin settings, with deposits consisting of several kilometers of siliciclastic turbidites (de Zoeten and Mann,

1999). In one of the elongate basin settings most proximal to the arc (located in the present- day Dominican Republic), the middle Eocene is unconformably overlain by lower

Oligocene strata of the Tabera Formation, with the upper Eocene section or the uppermost middle Eocene through upper Eocene section entirely missing, depending on locality (de

Zoeten and Mann, 1999).

Along the eastern and southern boundaries of the Caribbean Plate, deep-marine Atlantic strata exposed on Barbados via accretionary processes include lower Eocene to upper

Oligocene pelagic and hemipelagic sediment, and a paucity of continental sediment (Speed,

1981; Speed and Larue, 1982).

470

In Belize, atop the Southern Yucatan Block, Late Cretaceous carbonates are overlain by

Paleocene to (lower?) Eocene turbidite sandstone and shale of the Sepur/Toledo Formation, while younger strata are absent (Schafhauser et al., 2003).

In northeastern Honduras and northern Nicaragua, sedimentary fill of coastal basins

(including the Mosquitia Basin) from the Late Cretaceous through most of the Cenozoic consists of nonmarine sediment, especially red beds (Mills and Hugh, 1974; Mills and

Barton, 1996).

Along Nicaragua’s western (Pacific) margin, the Sandino forearc basin fill includes Eocene to Oligocene clastics deposited at inferred bathyal depths, however, previous workers do not mention the presence of an unconformity of Eocene-Oligocene age here (Kumpulainen,

1995; Ranero et al., 2000).

Along Costa Rica’s Caribbean margin, the middle Eocene to Oligocene of the Limón Basin contains shallow-marine carbonates (the Las Animas Formation) deposited atop structural highs and coeval hemipelagic mudstones, calcareous turbidites, and carbonate debris-flow sediments (the Senosri Formation) deposited in basinal areas (Brandes et al., 2007). An

Eocene-Oligocene boundary unconformity is not recorded here (Brandes et al., 2007).

In eastern Panama, a major hiatus spans most of the late Eocene through early Oligocene, with the deep-water Morti Tuffs overlain by middle Oligocene and younger tuff and limestone (Bandy and Casey, 1973; Bowland, 1993).

471

APPENDIX D-1.7. South America

Active tectonism initiated in the Eocene along the northern margin of South America, which prior to that time — since the Late Jurassic/earliest Cretaceous — had existed as a passive margin (Erlich and Barrett, 1992). Initiation of tectonism — both strike-slip motion and compression — generally progressed from the west to the east (from northern

Colombia and western Venezuela to eastern Venezuela and Trinidad), with eastward movement of the Caribbean Plate relative to the northern margin of the South American

Plate still active at present day (Erlich and Barrett, 1992).

Oblique collision caused subsidence and increasing accommodation through the early and much of the middle Eocene in the Maracaibo Basin of western Venezuela, but this tectonic loading waned by the end of the middle Eocene and gave way to a period of isostatic rebound that resulted in the emergence of most of the basin in the late Eocene and

Oligocene (Escalona and Mann, 2006).

In the Eastern Venezuela Basin of northeastern Venezuela, Late Cretaceous through

Paleogene deposition generally consisted of shallow marine clastics (Erlich and Barrett,

1992). Uplift during the Eocene and Oligocene resulted in erosion across the southern portions of the basin, with the Cretaceous through Oligocene sections missing furthest inland, and at least the Oligocene missing throughout the rest of the onshore (Erlich and

Barrett, 1992; Di Croce et al., 1999). Offshore, the upper Paleocene through Eocene passive margin sequence consists of a thin condensed section of relatively uniform thickness, and including an Eocene section generally composed of glauconitic shale grading upward into reworked glauconitic limestone, which is overlain by a lower

Oligocene unit consisting of fossiliferous glauconitic shale grading upward into glauconitic 472 limestone interbedded with fine- and medium-grained calcareous sandstone and containing a highly condensed transgressive section with a maximum flooding surface at ~35 Ma (Di

Croce et al., 1999). The offshore passive margin Oligocene sequence is separated from overlying lower Miocene and younger tectonostratigraphic sequence by a regional unconformity (Di Croce et al., 1999).

South America’s eastern continental margin has existed as a rifted passive margin since the

Cretaceous separation of the South American and African landmasses post-Gondwana breakup.

In the Guyana-Suriname Basin, which includes coastal to ultradeep eastern Venezuela to

French Guiana, record of an Eocene-Oligocene boundary unconformity is absent, with the basin-ward sections consisting of clays and shales of Eocene and Oligocene age (Nemčok et al., 2016).

Along the Brazilian continental margin, Cenozoic deposition is generally characterized by a regressive marine supersequence including both carbonate and siliciclastic deposition

(Bruhn and Walker, 1997; Mohriak, 2003).

Stratigraphic development of the Mundaú Basin has been significantly influenced by the activity of the Parnaíba River, however, this influence — and clastic deposition in the basin

— was drastically reduced in the middle Eocene, possibly due to drainage capture of the

Parnaíba by the São Francisco River (Karner and Driscoll, 1999). Evidence of an unconformity at the Eocene-Oligocene boundary is lacking in the Mundaú Basin, and in the greater Ceará Basin in general, with unconformities instead present in the middle

473

Eocene and middle Oligocene (Leopoldino Oliveira et al., 2020; Maia de Almeida et al.,

2020).

The Sergipe-Alagoas Basin of Brazil’s eastern margin has experienced a general marine regressive period since the Late Cretaceous (Campos Neto et al., 2007). As in the Mundaú

Basin to the north, development of the Sergipe-Alagoas Basin has been largely influenced by fluvial activity, but, unlike in the Mundaú Basin, the fluvial activity (by the São

Francisco River) influencing deposition in the Sergipe-Alagoas Basin was active throughout the development of the basin, with an increase in sediment flux and turbidite deposition starting in the middle Eocene likely due to drainage capture of the Parnaíba

(Karner and Driscoll, 1999). Furthermore, both the Sergipe and the Alagoas sub-basins contain a notable unconformity at the Eocene-Oligocene boundary (Karner and Driscoll,

1999; Campos Neto et al., 2007).

In the Espírito Santo Basin, tectonism has been active in the Cenozoic, including uplift of coastal ranges and salt tectonism in the offshore realm, and including a peak in activity in the Eocene to early Oligocene (Gamboa et al., 2010). As with much of the Brazilian margin, an erosional unconformity of early through middle Eocene age occurs throughout the basin (Gamboa et al., 2010). From the middle Eocene to Oligocene, deposition in the basin was controlled to a significant degree by mass wasting and the emplacement of mass transport deposits (which was likely promoted by high Eocene sedimentation rates; e.g.,

Bruhn and Walker, 1997), which caused margin progradation during this time (Gamboa et al., 2010). An erosional unconformity of approximate latest Eocene age is recorded in the basin, and separates Cretaceous to upper Eocene clayey turbidites and pelagic ooze from overlying uppermost Eocene/lower Oligocene to middle Miocene coarse-grained mass

474 transport deposits, submarine channel and overbank deposits, and some sand-rich turbidites

(Fiduk et al., 2004; Alves et al., 2009).

Campos Basin has experienced a general regressive pattern of deposition since the late

Paleocene (Contreras et al., 2010). The upper Paleocene to Eocene section consists of intercalated sandstones and carbonates on the inner shelf and wedges of turbidite sandstone and shale on the slope (Contreras et al., 2010). This interval is separated from overlying

Oligocene and Miocene cleaner, better-sorted turbidite sandstones and shales by a distinct

Eocene-Oligocene unconformity (at 33.9 Ma) that is explicitly linked to eustatic sea-level fall (Contreras et al., 2010).

In the Santos Basin, a distinct shift in depositional style occurs at the Eocene-Oligocene boundary due to capture and diversion of the Paraíba do Sul River out of the northern

Santos Basin and into the Campos Basin to the north (Karner and Driscoll, 1999; Modica and Brush, 2004; Berton and Vesely, 2016). Deposition in northern Santos Basin in the middle and late Eocene was characterized by sand-rich delta systems that deposited widespread sand-rich turbidites and debris flows across the northern basin, but the Paraíba do Sul drainage capture cut off this clastic influx, leading to transgression and sediment starvation in the Oligocene (Modica and Brush, 2004; Berton and Vesely, 2016). This is reflected in the observation that the modern shelf of the northern Santos Basin is more than

50 km landward of the position of the late Eocene shelf (Modica and Brush, 2004), and some authors record an Eocene-Oligocene unconformity coincident with this drainage capture and landward-shift of depocenters (Moreira et al., 2007; Berton and Vesely, 2016).

In the southern Santos Basin, the Eocene and the early Oligocene was generally a time of sediment-starvation, characterized by deposition of deep-water shale and calcareous shale

475 intercalated with fine- and medium-grained sandstone, before initiation of mud-rich clastic influx in the late Oligocene (Modica and Brush, 2004; Contreras et al., 2010).

With regard to the offshore basins of Uruguay (Punta del Este, Pelotas, and Oriental del

Plata basins), the upper Eocene to lower Oligocene sections are generally described by a common group of authors as a regressive sequence (Soto et al., 2011; Conti et al., 2017;

Morales et al., 2017). These authors do not discuss the presence of an Eocene-Oligocene boundary unconformity. Contreras et al. (2010) do note a ~30–33.9 Ma erosional unconformity on the shelf and correlative conformity basinward in the Pelotas Basin, though this timing may be best aligned with the early/late Oligocene boundary unconformity of Miller et al. (1985).

Along the Argentine continental margin, widespread erosion at the Eocene-Oligocene boundary of Paleocene–Eocene hemipelagic sediment, followed by deposition of a sizeable sediment drift in the Oligocene and early Miocene, has specifically been attributed to initiation of strong Antarctic bottom water circulation coeval with global cooling and opening of the Drake Passage (Hinz et al., 1999; Hernández-Molina et al., 2010; Gruetzner et al., 2012). An unconformity separating the Paleocene to Eocene section from the upper

Oligocene to Miocene and younger section is noted in the Colorado Basin (Bushnell et al.,

2000).

In the Fuegian Andes — in the thrust and fold belt region of the Austral and Malvinas basins — a major unconformity separates thick marine transgressive deposits of the upper middle Eocene to upper Eocene from the overlying deep-marine lower Oligocene to middle lower Miocene mudstones and sandstones and middle to upper Miocene shallow marine to fluvio-deltaic deposits (Olivero and Malumián, 2008). This unconformity is inferred to 476 represent the final stages of fold belt compression, and is succeeded by a significant deepening of the Austral Basin’s foredeep (Olivero and Malumián, 2008).

Along South America’s western margin, the Peru-Chile Trench has likely acted as a sink for sediment derived from Andean orogenesis since compressional activity began along the margin in the Late Cretaceous (Hoorn et al., 2010; Ramos, 2010; Horton, 2018). However, much of the coarser-grained fill (as much as hundreds of m) in the trench is inferred to be of Miocene or younger age (Scholl et al., 1968, 1970), while yet coarser-grained sediment lies landward of the trench, though this too may primarily consist of younger sediment produced during Miocene–Pliocene increases in the rate of onshore uplift (Zen, 1959;

Scholl et al., 1968). Seaward of the trench, up to 150 m of pelagic sediment sits atop the oceanic crust (Scholl et al., 1968). Generally, trench and adjacent deep-marine sediment is poorly characterized, and sufficient information is lacking on whether an Eocene-

Oligocene boundary unconformity might be expressed in coarser-grained sequences here.

On Peru’s continental margin, numerous Mesozoic and Cenozoic foreland basins occur, including the Talara Basin, whose youngest fill is of late Eocene age, as well as the Sechura and Pisco basins, which contain marine sediment ranging from middle and late Eocene to recent in age (Travis et al., 1976; Dunbar et al., 1990). In both the Sechura and Pisco basins, an unconformity spans most of the Oligocene (Travis et al., 1976; Dunbar et al., 1990).

In coastal Ecuador, uplift-driven emergence occurred in the late Eocene, lasting until the latest Oligocene (Jaillard et al., 1995, 2000). Thus, much of the Oligocene section is missing throughout this region (Jaillard et al., 1995, 2000).

477

An Eocene-Oligocene erosional unconformity is present in the onshore Borbón Basin of coastal Ecuador, as well as in the onshore Pacific Coastal Basin in western Colombia and the offshore Manglares Basin (which makes up the southern portion of the Pacific Frontal

Basin ) off northern Ecuador and southern Colombia (Marcaillou and Collot, 2008).

Marcaillou and Collot (2008) infer that this unconformity marks the Incaic phase of compression that was active in the late Eocene and early Oligocene. These authors note

(based on seismic data) a marked change in the character of sediment in the Manglares

Basin, from poorly-bedded clastic units underlying the unconformity to well-bedded, presumably coarser sediment above the unconformity, and also discuss previous work in the Borbón Basin, where bathyal to out shelf deposition of forearc sediment occurred pre- unconformity and claystone and fine-grained sandstone deposition occurred after, and in the Pacific Coastal Basin, where siliceous limestones, calcareous sandstones, and shales underly the unconformity and conglomerates overly the unconformity (Marcaillou and

Collot, 2008). In the Middle Magdalena Valley, Eastern Cordillera, and Llanos Basin to the east of the Pacific Coastal Basin, Eocene and Oligocene deposition was nonmarine

(Cooper et al., 1995).

478

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